Tropical Geomorphology

Tropical Geomorphology

Although similar geomorphic processes take place in other regions, in the tropics these processes operate at different rates and with varying intensities. Tropical geomorphology therefore provides many new insights regarding geomorphic processes. This textbook describes both the humid and the arid tropics. It provides thoroughly up-to-date concepts and relevant case studies, and emphasises the importance of geomorphology in the management and sustainable development of the tropical environment, including climate change scenarios. The text is supported by a large number of illustrations, including satellite images. Student exercises accompany each chapter. The book highlights three areas: • Geology, landforms and geomorphic processes in the humid and arid tropics • Source-to-sink passage of water and sediment from the mountains to the sea • Anthropogenic alteration of natural geomorphic rates and processes, including climate change. Tropical Geomorphology is an ideal textbook for any course on tropical geomorphology or the tropical environment, and is also invaluable as a reference text for researchers and environmental managers in the tropics. Avijit Gupta is a Honorary Principal Fellow at the University of Wollongong, Australia and a Visiting Scientist at the Centre for Remote Imaging, Sensing and Processing, National University of Singapore. He was educated at Presidency College, Kolkata, and Johns Hopkins University. He has held university positions in India, the USA, Singapore and the UK. His research interests focus on fluvial geomorphology in the tropics, rivers with high-magnitude floods, large rivers, urban geomorphology and the application of remote sensing in geomorphology. Dr Gupta has served as a Committee Member of the International Geographical Union Commission on Measurement, Theory and Applications in Geomorphology (COMTAG) and the International Association of Geomorphologists (IAG). He is currently the Chair of the IAG Working Group on the Effect of Climate Change on Large Rivers and Deltas. He is a member of the American Geophysical Union, the Association of American Geographers, the Society of Sedimentary Geologists and the International Association of Hydrological Sciences. He is also a corresponding member of the Académie Royale des Sciences d’Outre-Mer, Belgium. He is on the editorial board of two journals on geomorphology. Dr Gupta has written eight books and over seventy research papers. He recently edited The Physical Geography of Southeast Asia (2005) and Large Rivers:€Geomorphology and Management (2007).

Tropical Geomorphology Avijit Gupta Centre for Remote Imaging, Sensing and Processing National University of Singapore and School of Earth and Environmental Sciences University of Wollongong, Australia

cambridge university press

Cambridge, New York, Melbourne, Madrid, Cape Town, Singapore, São Paulo, Delhi, Tokyo, Mexico City Cambridge University Press The Edinburgh Building, Cambridge CB2 8RU, UK Published in the United States of America by Cambridge University Press, New York www.cambridge.org Information on this title:€www.cambridge.org/9780521879903 © Avijit Gupta 2011 This publication is in copyright. Subject to statutory exception and to the provisions of relevant collective licensing agreements, no reproduction of any part may take place without the written permission of Cambridge University Press. First published 2011 Printed in the United Kingdom at the University Press, Cambridge A catalogue record for this publication is available from the British Library Library of Congress Cataloging in Publication data Gupta, Avijit. Tropical geomorphology / Avijit Gupta. p.â•… cm. Includes bibliographical references and index. ISBN 978-0-521-87990-3 (hardback) 1.╇ Geomorphology–Tropics.â•… I.╇ Title. GB446.G87 2011 551.410913–dc22 2011011810 ISBN 978-0-521-87990-3 Hardback Cambridge University Press has no responsibility for the persistence or accuracy of URLs for external or third-party internet websites referred to in this publication, and does not guarantee that any content on such websites is, or will remain, accurate or appropriate.

In Memory of Reds

Contents

Preface



page xiii

Part I╇ The tropical environment

1╇ Introduction 1.1 Geomorphology in the tropics 1.2 Traditional tropical geomorphology 1.3 Modernisation of tropical geomorphology 1.4 Structure of tropical geomorphology 1.5 Structure of the book

2 Geological framework of the tropical lands 2.1 Introduction 2.2 A brief introduction to plate tectonics 2.3 Major landforms across the tropics 2.4 Interrelationships:€plate tectonics, landforms, erosion and sediment production Questions

3 Tropical hydrology 3.1 The tropical climate:€a brief review 3.2 Temperature 3.3 Wind circulation 3.4 Precipitation 3.5 Tropical disturbances 3.6 Miscellaneous factors 3.7 Water balance 3.8 Climate and geomorphology in the tropics Questions

4 Erosion and land cover in the tropics 4.1 Erosion from tropical rainfall 4.2 Distribution of natural vegetation in the tropics 4.3 Tropical rain forests 4.4 Tropical deciduous forests, grasslands and deserts vii

3 3 7 8 9 11 13 13 13 20 27 29 31 31 31 32 34 37 43 44 47 48 49 49 52 52 55

Contents

viii

4.5 Anthropogenic alteration of the tropical vegetation Questions



57 58

Part II╇ Process geomorphology in the tropics

5 Weathering in the tropics 5.1 Introduction 5.2 Sub-processes of weathering:€a brief review 5.3 Products of weathering 5.4 Weathering and vertical zonation 5.5 Pans and crusts 5.6 Effects of weathering 5.7 Tropics and weathering Questions

6 Slopes:€forms and processes 6.1 Properties of a slope 6.2 Mass movement on hillslopes 6.3 Running water on hillslopes 6.4 Storage and transfer of surficial material on tropical slopes 6.5 A general description of tropical slopes Questions

7 Rivers in the tropics 7.1 Components of a river system 7.2 Water in river channels 7.3 Sediment in river channels 7.4 Channel geometry 7.5 Channel network and nodes 7.6 River systems of the humid tropics Questions

8 Alluvial valleys 8.1 Fluvial depositional environment 8.2 The alluvial valley 8.3 The channel alluvium 8.4 Bars 8.5 Floodplain 8.6 Terrace 8.7 Valley margins 8.8 Sediment transfer along the valley axis Questions

61 61 63 69 75 77 78 80 81 82 82 84 93 94 99 99 101 101 101 104 108 119 120 127 129 129 130 132 135 136 139 140 141 141

Contents

ix

9 Large rivers in the tropics 9.1 Introduction 9.2 Characteristics of a large river 9.3 The Amazon 9.4 The Zambezi 9.5 The Ganga–Brahmaputra system 9.6 The Mekong 9.7 The importance of major tropical rivers Questions

10 The tropical coasts 10.1 Introduction 10.2 Types of coast 10.3 Moving water:€tides, waves and currents 10.4 Rocky coasts 10.5 Non-rocky coast 10.6 Coastal sand dunes 10.7 Coastal tropics 10.8 Coral reefs 10.9 Tropical coasts and time Questions

11 Deltas in the tropics 11.1 Introduction 11.2 Distribution of deltas in the tropics 11.3 Age and evolution of deltas 11.4 Delta morphology 11.5 Delta sediments and sedimentary structures 11.6 The Ganga–Brahmaputra Delta:€a case study 11.7 Deltas in the tropics: a summary Questions

12 The arid tropics 12.1 Arid areas 12.2 Geological characteristics of arid lands 12.3 Arid hydrology 12.4 Arid landforms 12.5 The rock desert 12.6 Running water in arid lands 12.7 Aeolian geomorphology of sandy areas 12.8 Conclusion Questions

143 143 143 147 152 155 164 169 169 170 170 171 172 179 181 186 186 189 191 194 195 195 196 197 201 202 203 207 208 209 209 210 211 215 215 216 223 230 230

Contents

x

13 Tropical highlands 13.1 Importance of highlands 13.2 Glaciation in tropical mountains 13.3 Mechanics of mountain glaciation 13.4 Glacial forms and processes 13.5 Slopes and valley floors in high mountains 13.6 Rivers in the tropical mountains 13.7 Sediment from tropical mountains 13.8 Conclusion Questions

14 Volcanic landforms 14.1 Introduction 14.2 Types of volcano and the related landscape 14.3 Lava and pyroclastic deposits 14.4 Volcaniclastic flows: debris avalanches and flows 14.5 Landscape on flood basalts 14.6 Conclusion Questions

15 Tropical karst 15.1 Introduction to karst 15.2 Karst in the tropics:€the geographical distribution 15.3 Karst hydrology 15.4 Dissolution of karst rocks 15.5 Karst landforms 15.6 Karst in the tropics 15.7 Tropical karst as an environment Questions

16 Quaternary in the tropics 16.1 Introduction 16.2 History and structure of the Quaternary 16.3 Quaternary glaciation in the tropics 16.4 Climate change 16.5 Sea-level change 16.6 The Ganga River system:€Quaternary adjustments 16.7 Quaternary changes around the Sunda Shelf 16.8 Conclusion Questions

232 232 234 236 240 246 247 249 251 252 254 254 254 257 262 266 268 268 270 270 271 272 273 274 281 284 286 287 287 288 290 292 295 298 299 302 303

Contents

xi



Part III╇ Anthropogenic changes

17 Anthropogenic alteration of geomorphic processes in the tropics 17.1 The beginning 17.2 Deforestation, land use changes and rural migration 17.3 Temporal and seasonal patterns of sediment transport 17.4 Spatial transfer of sediment 17.5 Impoundments along rivers and their effects 17.6 Application of geomorphology towards a better environment Questions

18 Urban geomorphology in the tropics 18.1 Introduction to urban geomorphology 18.2 Urbanisation in developing countries 18.3 Three examples of geomorphic hazards and their amelioration 18.4 The general nature of urban geomorphological problems 18.5 Geomorphology and urban management Questions

19 The future with climate change 19.1 Climate change and the future 19.2 A robust prediction of the effects of climate change in the tropics 19.3 Geomorphological adjustments in the tropics from climate and sea-level changes 19.4 The noise effect of anthropogenic changes 19.5 Tropical geomorphology in the near future Questions References Index Colour plates appear between pages 212 and 213.

307 307 311 314 315 319 323 324 325 325 328 330 337 338 341 342 342 343 345 346 347 348 349 374

Preface

This is an introduction to a very large part of the world’s surface with rich and varied landforms. The tropics include high mountain ranges, major rivers, ancient surfaces, large alluvial plains and deltas, arid landscapes, and wonderful examples of volcanic landforms. The tropical oceanic coverage is huge and it influences the world’s climate. It is surprising that, in spite of a recent spurt in case studies, our knowledge regarding the geomorphology of the tropics remains limited and that case studies from the tropics have hardly been used for generalisation and theory construction. This lacuna is fascinating, especially as all world maps on sedimentation rates indicate that huge amounts of sediment are pouring into the oceans from certain parts of the tropics, as a result of events happening inland. No single template can exist for tropical geomorphology given the wide variations in regional geology, climate, and land cover. A major part of the tropics carries old subdued landscapes that have evolved since the Gondwana era, whereas other parts, including active plate boundaries and large alluvial plains, are much younger and may record a rapid rate of erosion and sedimentation. The original land cover is changing drastically and the current rates of geomorphic processes are no longer natural everywhere. The old images of a chemical-weathering driven, deep regolith-covered landscape of large plains and isolated hills are only partly correct. This book is an attempt to present the tropics in their rich and varied reality. This objective has determined the selection and arrangement of topics included for discussion and the level at which they have been covered. The book updates the concept of tropical geomorphology in stressing the increasingly important anthropogenic alterations of the landscape. The book ends with an attempt to look into the future of the tropics, given current modifications such as climate change. The approach of the book also defines the expectations from its users. It is written primarily as an advanced undergraduate textbook and assumes that the readers already have a basic background in physical geography or geology. It makes no further demands apart from an interest in the tropics. Two aspects of the book should be mentioned. It is well illustrated, as a book in geomorphology should be, and the illustrations are of equal importance to the text. The illustrations include several high-resolution satellite images from IKONOS. Given the 1 m resolution of these images, they are a wonderful tool for future geomorphological studies, a tool which is probably not as well used as it should be. We now even have commercial satellites capable of providing images at 50 cm resolution. The second aspect of the book is the set of questions at the end of the chapters. These are a mixed batch; some for problem solving, others encouraging the reader to think in more detail beyond the text. Like the illustrations, these questions complement the text. xiii

xiv

Preface

The formulation of this book started decades ago, when I was a student in Presidency College, Kolkata and had trouble matching what I read in books with the landscape around me or even with the landscape displayed on the topographical sheets of India. I am immensely grateful to the late M. G. Wolman for allowing me to do my PhD fieldwork in Jamaica which was a liberating educational experience. A long stay in Singapore permitted field access to Southeast Asia. My various friends and colleagues completed my education by allowing me to work with them in the field in various parts of the tropics. I am indebted to all of them. The introductory chapter has benefited tremendously from the comments of Professor Wolman. Parts of the book were also read and commented on by Liew Soo Chin, JeanClaude Thouret, Richard Corlett, and Anthea Fraser Gupta.The Centre for Remote Imaging, Sensing and Processing (CRISP), National University of Singapore very kindly allowed me to use satellite images from their archives. I have also been permitted to use illustrations from a number of publications, which are acknowledged in specific places. I should acknowledge the kindness of Lee Li Kheng in drafting the figures and of David Appleyard in transforming my photographs into publishable material. I am grateful to Jean Rollinson for her careful copy-editing and to Laura Clark of Cambridge University Press for guiding me through the production stage. Avijit Gupta

Part I

The Tropical Environment

1

Introduction

The tropics are my element, and I have never been so constantly healthy as in the last two years. Alexander von Humboldt

1.1╇ Geomorphology in the tropics In tropical geomorphology we are constantly surprised by new discoveries. Currently, we have a limited understanding of the geomorphic processes, landforms and sediment in the tropics. Furthermore, the rapid, ongoing anthropogenic development in the tropics continues to modify landforms and operating processes and change the rates of erosion and sedimentation from the expected natural norms. Geomorphology in the tropics provides twin opportunities to discover new facts and to apply such information to managing the environment for a sustainable future. Tropical geomorphology thus has a tendency to look forward rather than look back exclusively at past landforms. The tropics are in essence a climatic region, although the only shared meteorological component across the belt of low latitudes is high temperature. Considerable climatic variations exist across the tropical zone:€the most impressive of which is the variation in rainfall. The annual total, the seasonal pattern and occasional synoptic disturbances all vary across the tropics. The Amazon lowlands, the Rift Valley of Africa, Raub al Khali of the Arabian Peninsula, the Ganga–Brahmaputra Delta, the wetlands of eastern Sumatra and a considerable part of the Red Heart of Australia are all areas of low elevation in the tropics, but they exhibit huge differences in rainfall. The tropics can be divided into two primary units based on annual rainfall:€ the humid tropics and the arid topics. The transition between the two can be sharp (for example, where an orographic barrier prevails), or gradational (with a subhumid zone in between). About half of the tropical land surface is humid, with the annual rainfall exceeding annual evapotranspiration. The rest is subhumid or arid. Certain climatic characteristics, such as high temperature, high intensity of rainfall and high potential evapotranspiration are generally associated with the tropics but do not occur with the same intensity everywhere. Formally, the tropics can be defined as an area of radiative surplus at the Earth–Â� atmosphere interface, bounded by anticyclonic circulations near the 30° north and south latitudes (Reynolds, 1985). The margin of the tropics is best perceived in a pragmatic fashion as a fluctuating boundary, between 30 and 35 degrees of latitude. It is an area 3

4



Introduction

Fig. 1.1 Major pressure belts and wind systems of the Earth. Shaded area approximately indicates tropics and subtropics. The

vertical circulation of air is shown diagrammatically towards the right of the figure over which the Hadley Cell operates, and the three-dimensional pattern of air movement shows a distinct separation between the Hadley and the Ferrel Cells at these latitudes (Fig. 1.1). The two traditional latitude markers, identified as Tropic of Cancer and Tropic of Capricorn, are not effective boundaries. Thus, the tropics, as defined in this book, also include areas conventionally identified as the subtropics. Nearly 60 per cent of the total surface area of our planet lies between the 35° N and S Â�latitudes. The tropical oceanic expanse is huge and plays an important role in Â�influencing the climate of the world (Graham et al., 1994). In spite of such extensive coverage, case studies from the tropics have contributed very little to mainstream theories in geomorphology. Like climate, landforms and operating geomorphic processes are not the same across the tropics. The tropics are an assemblage of active tectonic belts, ancient cratons, alluvial valleys and subsiding deltas (Fig. 1.2). The early tropical geomorphologists did not always recognise such wide-ranging geologic variations, putting too much emphasis on the hot and humid climate as the prime controlling factor. As a result, the tropics used to be perceived as a set of climo-morphogenetic landforms, where physical features are primarily controlled by the ambient climate. The characteristics of the landscape were generally explained by assuming that they evolved in a hot and humid location over a very long period of time. The arrival of the theory of plate tectonics finally destroyed such concepts.

Geomorphology in the tropics

5

(a)

(c)



(b)

(d)

Fig. 1.2 Varieties of tropical landforms:€(a) Western Himalaya Mountain (high tectonic mountains); (b) arid Central Australia

(arid landscape on ancient craton); (c) Auranga River, India (seasonal river on Gondwana rocks). From Gupta and Dutt, 1989. (www.borntraeger-cramer.de); (d) Mekong Delta (river-dominated part of a Holocene delta). Photographs: A. Gupta

Landmasses that are now in the tropics were once part of a single large continent on Earth, known as Pangaea, before its break-up about 200 million years ago. The present tropical landmasses (Australia, part of Southeast Asia, the Indian subcontinent, Africa, South America) and a number of islands of various dimensions together with the cold Antarctica constituted the southern part of Pangaea, known as Gondwana or Gondwanaland (Box 1.1). The present physical landforms reflect the entire geological history from Pangaea. This includes the break-up, the nature of drifting of the separated landmasses towards the equator, the geomorphic processes that began to operate once they reached the tropical zone, and the current highly active anthropogenic modification of both landforms and processes. Traditionally, the Indian Peninsula was viewed as an excellent example of an assemblage of tropical landforms that had developed over a long period of time in a warm humid climate. The peninsula, however, had a near-polar location before the disintegration of Pangaea. It then drifted through a wide range of latitudes to reach its present position and collided with the Eurasian Plate to form the Himalaya Mountains and fuse with the Asian landmass (Fig. 1.3). The present monsoon system and operating geomorphic processes developed mostly after the collision. Landforms and subsurface sediment in the tropics were also modified by climate and sea-level changes during the Pleistocene. Like elsewhere on Earth, landforms in the tropics are multifactorial in origin. The history of the landmasses goes back in time beyond Pangaea, but we will start our narrative from this particular supercontinent.

6



Box 1.1

Introduction

Gondwana/Gondwanaland From the Upper Palaeozoic to the Cretaceous, the entire continental landmass of the Earth formed a single continent called Pangaea. The southern half of Pangaea is collectively known as Gondwana or Gondwanaland, both names being used interchangeably. Gondwana was located south of the equator and a very large part of it stretched to high southern latitudes near the South Pole, which, as expected, was a very cold location. As Pangaea broke up, Gondwana disintegrated into several units which moved away from each other. These fragments of Gondwana are now known as Antarctica, Australia, India, part of western Asia, Africa, Madagascar and South America. Except for Antarctica, all land masses drifted towards the equator and to warmer locations. The evidence of their former connections lies in the similarity of the rocks of the PermoCarboniferous age found on all these continents. At the base of such rocks is a glacial sediment, clay with boulders, the lithified form of which is called a tillite. This occurs on all continents that were part of Gondwana, indicating the former cold location and presence of glacial ice. The rocks also contain a characteristic cold-weather flora called Glossopteris (from the Greek for tongue-like, as the leaves of the flora resembled long tongues) and other associated flora such as Gangamopteris. Glossopteris is found in rocks of this age on the continents that constituted Gondwana, thereby indicating their past union and the shared glacial climate. From the Upper Carboniferous to the Jurassic, similar climatic conditions and rock types occurred on the continents that formed part of Gondwana. Such rocks were studied in central India by H. B. Medlicott, who first used Gondwanaland as a term in 1872 in an unpublished report (Krishnan, 1982). The name is derived from a part of central India which carries the diagnostic geology and used to be the kingdom of the Gonds. The tillites were followed by thick deposits of fluviatile or lacustrine origin, in places deposited in structural basins and bearing coal seams. These in turn were followed by other sedimentary rocks but without coal. Volcanic lavas are found on top of the Gondwana rocks in a number of places. The similarity of geology in locations now widely separated from each other indicates the existence of a single landmass which later disintegrated. The present physical features reflect geological history. For example, extensive areas in Australia, the Indian Peninsula, Africa and South America are inherited from Gondwana. As a result these areas are underlain by old, hard, generally metamorphosed rocks on which weathering, soil formation and erosional processes operate very slowly. These are the cratons. Elsewhere the edges of the continents may reflect the history of their passage and collision with other plates. The Himalayan range was formed following the collision of India with Eurasia. The Andes was formed by subduction of the Cocos, Nazca and Antarctic Plates underneath the South American Plate. These active margins now form tectonic mountains which are the source areas of big rivers that drain the continents. The rivers flow from high active margins across the continent towards the continental margin that is tectonically passive. These are quiet places where the big rivers deposit their sediment and build deltas. The Amazon, the biggest of all, is an excellent example. The present river came into existence after the formation of the Andes to flow eastward along a structural low and to build a huge delta on the passive side of the continent into the South Atlantic Ocean. In contrast, Africa has not drifted much and its margins are passive, with tectonic mountains occurring only at the extreme northern and southern ends. Africa, therefore, is a continent with old surfaces at various levels. The present landscape on any continent thus reflects its past history inherited from Gondwana, marginal modifications that reflect the nature of its movement since the break-up of Pangaea, and other happenings over time such as large eruptions of basalt (western India) or the formation of rift valleys (East Africa).

7



Fig. 1.3

Traditional tropical geomorphology

Drifting of the Indian plate. Adapted from US Geological Survey figure

1.2╇ Traditional tropical geomorphology We know very little about the geomorphological knowledge that existed before the arrival of the Western maritime powers in the tropics. There must have been some, given the widespread successful utilisation of water as a resource, often with designed structures. After the establishment of global maritime sailing routes, the tropics were visited from the last decade of the eighteenth century with astonishment, enthusiasm and insight by Western scientific explorers including Alexander von Humboldt, Alfred Russel Wallace and Charles Darwin. This revolutionised primarily the biological sciences, but remarkable observations on regional geology and geomorphology were also registered. We may recall Humboldt’s work in tropical South America or Wallace’s description of the landforms of Southeast Asia. In 1859, the year On the Origin of Species was published, Darwin was honoured by the Royal Society for his research, and the citation listed his contribution to the geology of the Andes. We should also recognise the work of German geologists in various parts of the tropics: Sapper, Passarge, Bornhardt and others, and that of Dana (1850) from the United States in the Pacific. Dana provided a very early discussion on fluvial processes in the tropics. Later, as part of colonial governance, land- and soil-related information was collected, processed and published in a number of countries. Some of these are now classics, such as reports by Buchanan on Indian laterite, and Mohr and Van Baren on tropical soils (Buchanan, 1807; Mohr and Van Baren, 1954). Topographical maps and official

8

Introduction

geological reports also started to appear. All these later became very useful for studying local geomorphology. Up to the second half of the twentieth century, or even later, the teaching of geomorphology in the tropical countries was carried out using standard textbooks written for European or North American students (Gupta, 1993). These did not reflect the ambient landscape and local or regional examples were rarely furnished. Large Andean landslides, coral reefs or tropical karsts were mentioned, but the common reference to tropical geomorphology was the climo-morphogenetic regions of Peltier (1950) or De Martonne (1951) or Büdel (1982). Krynine’s (1936) approach was different; he studied the relationship between geomorphology and sedimentology in the humid tropics. The general emphasis was on landforms, and climatic characteristics were used to explain their nature and distribution. For years, geomorphological studies in the tropics were driven by two concepts:€the climoÂ�morphogenetic region and the Davisian cycle of erosion. The typical tropical landscape was perceived as a stable erosion surface, studded with low hills (called inselbergs or bornhardts) and underlain by soil and deeply weathered rock. Geomorphology was a description of landforms.

1.3╇ Modernisation of tropical geomorphology Tropical geomorphology started to lose its essentially climate-based approach in the second half of the twentieth century. The writings of L. C. King (1951, 1962) on the South African scenery led the practitioners to think in terms of rock types and geological history as important explanations for the ambient landscape. About this time, a number of case studies started to appear that were more in line with process-based approaches prevalent in other parts of the world. Early examples include Ruxton and Berry (1957) on granite weathering in Hong Kong, Fournier (1960) on erosion rates, Simonett (1967) on earthquake-generated landslides that eroded the mountains of New Guinea, and Coleman’s study (1969) on the morphology and sedimentation of the Brahmaputra River. These papers highlighted rationality, methodology and excitement in tropical geomorphology. Case studies became easily available from this time, due to the publication of several regional collections (Jennings and Mabbutt, 1967; Davies and Williams, 1978; Dardis and Moon, 1988; Warner, 1988). Prior to these, regional studies were difficult to find as they were published in local journals, commonly with a limited circulation. All this overlapped with the diffusion of modern mainstream concepts, techniques and textbooks (Horton, 1945; Strahler, 1952a, 1964; Leopold et al., 1964; Young, 1972) to the practitioners in the tropics. From the 1970s, such diffusion, along with the presence of process-oriented geomorphologists in the tropics, led to a tremendous increase in the quantity and quality of papers in tropical geomorphology. Textbooks in process geomorphology with a tropical slant started to appear (Douglas, 1977; Faniran and Jeje, 1983; Thomas, 1994), and tropical case studies also were mentioned in mainstream textbooks (Schumm, 1977). The new research indicated that the same geomorphic processes operate in the tropics as elsewhere, but they operate at different rates and with varying intensities (Selby, 1993). Research publications were generally in English

9

Structure of tropical geomorphology

but, given the large number of tropical countries, research and institutional reports were also published in a number of other languages. This at times constituted linguistic demands on the tropical geomorphologist. For example, a researcher on the lower Amazon Basin is handicapped without knowledge of Portuguese. The diffusion of ongoing research across the tropics was limited before electronic communication became common. The free distribution of the Tropical Geomorphology Newsletter, published from 1986 to 1996, used to help. As computer-based searches became common, the problem of inadequate diffusion of knowledge largely disappeared.

1.4╇ Structure of tropical geomorphology Tropical geomorphology highlights three areas: 1. geology, landforms and geomorphic processes across the tropics 2. the passage of water and sediment from the mountains to the coast, mainly via river systems:€a large volume of moisture is in circulation over the humid tropics 3. anthropogenic alteration of the natural rates and processes, associated environmental degradation, and related geomorphic principles for better environmental management. We can break these down to a list of specific topics that should be studied (Table 1.1). As this table indicates, a combination of common and exotic factors characterises tropical geomorphology. It is necessary to emphasise that although climate acts as an important control on landforms and processes in both humid and arid tropics, the operating processes and the characteristic landforms at a particular location reflect a diversity of causes. Gardner et al. (1987), writing on locations in Central America and the Caribbean, focused on three selected topics among a diversity of regional geomorphic processes and landforms:€ karst; alluvial fans; and tectonism along convergent plate margins. This was justified, as these three are the most characteristic of the region, representing the dominance of lithology, a rapid fluvial depositional process and neotectonics. Elsewhere in the Caribbean, other factors such as hurricanes or volcanoes dominate local geomorphology. The fascination and challenge of present-day geomorphology in the tropics lies in the recognition of such regional diversity. Geomorphological rates are high in the humid tropics, especially in high-relief areas (Fig. 1.4). A number of world maps of suspended sediment yield have been published since the 1980s (Milliman and Meade, 1983; Walling and Webb, 1987; Milliman and Syvitski, 1992). All such maps show elevated rates for most of the humid tropics. Working with a data set of 280 rivers, Milliman and Syvitski found very high sediment rates for South Asia and high oceanic islands, and identified basin relief as a very important factor behind the high sediment yield. They interpreted relief as a surrogate for tectonism and proposed that ‘the entire tectonic milieu of fractured and brecciated rocks, oversteepened slopes, seismic and volcanic activity … promotes the large sediment yields from active orogenic belts’ (Milliman and Syvitski, 1992:€539–540). The high sediment yield is also due to intense tropical rain falling on such mountains, for at least part of the year.

Introduction

10

Table 1.1╇A summary description of tropical geomorphology Topic

Description

Major controls in geomorphology

Location of tectonic belts, volcanoes, cratons, alluvial valleys, deltas, etc. as determined by plate tectonics Wind pattern and rainfall systems (especially tropical storms) Distribution of vegetation cover Deforestation, agricultural expansion, urbanisation and channel controls Tropical weathering, and its effect on slope material and river load Mass movements on tropical slopes Rivers, a number of which are seasonal and prone to flooding Glacial, glacio-fluvial and fluvial processes operating on high mountain slopes Fluvial and aeolian processes in the arid tropics Coastal processes, presence of mangroves, salt marshes and coral reefs Tectonic movements and volcanism Pleistocene glaciations of the tropical mountains Climate change Sea-level changes affecting coasts and lower river reaches Common anthropogenic changes Global warming and climate change

Major operating processes; same as in other parts of the world, but different in rates and relative importance

Quaternary inheritance

Present and future changes

The natural rates of erosion, however, are not necessarily in operation everywhere. Parts of the tropics have a long history of intense human occupation leading to modification of the natural landscape. This process has accelerated since about 1950 and the present landforms and processes may be natural only in certain areas. For example, destruction of the tropical rain forest has increased the annual sediment yield in Southeast Asia from less than 102 tkm−2 to more than 103 tkm−2 (Gupta and Krishnan, 1994; Gupta, 2005a). Transfer and storage of this excessive sediment sequentially affect hillslopes, gullies, rivers and coastal forms. Satellite images, as on the cover of this book, show sediment plumes at the mouth of rivers whose basins are undergoing significant anthropogenic changes (Gupta and Krishnan, 1994). The high rate of erosion and sedimentation therefore is a function of relief, climate and land use − all three. Destruction of vegetation cover and accelerated instability of slopes lead to excessive sediment transfer and storage in the river systems. In contrast, thousands of large dams globally block the passage of sediment downstream, the total world volume rising to about 50 km3 of sediment each year (Mahmood, 1987). We need to keep all such modifications

11



Structure of the book

Fig. 1.4 Change of state of landforms. Note according to his estimate, a combination of high relief and climate leads to

approximately the same amount of change in landforms in ten years in the Himalaya Mountains as has occurred since the end of the last glacial advance in Western Europe. The graphs for New Zealand compare the role of the third factor in accelerating geomorphic processes:€land use. From Selby (1993). By permission of Oxford University Press

in mind. Environmental problems in the developing countries that constitute most of the tropics involve land and water degradation to a large extent. The recognition and solution of such problems depend on geomorphological knowledge of the area. Current tropical geomorphology therefore combines scientific research with the application of its findings for the betterment of people. Such a linkage will become crucial with climate change that is expected to alter the rate and intensity of operating geomorphic processes. In one of the crime fictions of Ellery Queen (1934), a detective was described as a prophet looking backwards. A tropical geomorphologist frequently needs to be that in order to explain a particular landform. At the same time, the geomorphologist also needs to operate in the conventional prophetic fashion, predicting the future determined by anthropogenic alterations including climate change. It is a challenging but wonderful occupation.

1.5╇ Structure of the book This book is divided into three sections. The first section introduces the physical background of the tropics against which the operating geomorphological processes are subsequently presented. An introduction to the geological framework of the tropics and landforms (Chapter 2) is followed by a discussion of tropical hydrology (Chapter 3), and an account of vegetation cover and land use including a brief introduction to anthropogenic modification of the land and water (Chapter 4). Part 2, dealing with geomorphic processes, starts with weathering that prepares slope material and river sediment for subsequent soil formation and removal (Chapter 5). This is followed by chapters on slopes and slope processes (Chapter 6) and tropical rivers (Chapter 7–9). The coastal environment is discussed next:€first as a general case (Chapter 10) and

12

Introduction

then regarding the special case of tropical deltas (Chapter 11). Hydrology, river systems and aeolian processes of the arid tropics are presented and compared with the forms and processes of the humid tropics in Chapter 12. The next chapter describes the tropical mountains, both glaciated and non-glaciated, as sources of water and sediment to rivers and the storage and transfer of both along river systems. Chapter 14 is an exposition on volcanic landforms in the tropics as a special case and Chapter 15 is on tropical limestone geomorphology. The section on geomorphic processes ends with a brief discussion on Quaternary inheritance and its inertia effects (Chapter 16). The third part of the book deals with anthropogenic influences of various kinds:€application of geomorphology in the management of the tropical environment, the urban landscape and the possible future scenarios due to climate change (Chapters 17–19). Although the chapters are structured separately for easy comprehension, the processes discussed individually usually come together to shape the land and to determine the passage of water and sediment from the higher parts of a drainage basin to the sea. In large parts of the humid tropics, this transfer is naturally efficient as mapped by Milliman and Syvitski (1992) and other researchers. Anthropogenic modifications and climate change accelerate such rates, and comprehension of such changes rises out of our understanding of tropical landforms and processes.

2

Geological framework of the tropical lands

Civilization exists by geological consent, subject to change without notice. Will Durant

2.1╇ Introduction A little more than a century ago, William Morris Davis described the form of the land as a function of three variables:€structure, process and time (Davis, 1899). The statement is equally true for tropical landforms. Geology is usually the primary determinant of land:€its elevation, steepness and stability. Understanding of tropical landforms should start with an introduction to the geological framework. The distribution of tectonic mountains that produce large quantities of sediment; stable areas underlain by old rocks (cratons) that produce very little; and large river valleys where the sediment is temporarily stored while being transported to the sea are all determined by plate tectonics. A familiarity with the distribution of plates and their boundaries in the tropics is necessary to understand regional landforms and operating processes.

2.2╇ A brief introduction to plate tectonics The theory of plate tectonics started to develop in the 1960s. The surface of the Earth is divided into several three-dimensional solid bodies called plates that drift slowly, usually at the rate of several centimetres a year. It has been pointed out that this is the rate at which our fingernails grow. The top surface of a plate may form the ocean floor or a continent or a combination of both. Everything on a plate moves with it. The plates come in various sizes. The largest plates are seven in number:€the Australian–Indian Plate, the Antarctic Plate, the Eurasian Plate, the African Plate, the Pacific Plate, the North American Plate and the South American Plate. Plates of smaller size include examples such as the Philippine Plate, the Caribbean Plate and the Nazca Plate (Fig. 2.1). The plates are able to move because temperature and pressure increase with depth inside the Earth. At a depth of about 120 km the temperature is approximately 1350°C. This softens the ambient material, allowing rigid plates to move slowly through it. The rigid layer of rocks at the surface is known as the lithosphere; and the softer part of the upper mantle through which the plates move is the asthenosphere. 13



Fig. 2.1

Map of the tropical latitudes, showing the major tectonic plates

A brief introduction to plate tectonics

15

(a)

(b)

(c)

(d)



Fig. 2.2

Types of plate boundary. (a) Divergent boundary; (b) Convergent boundary between an oceanic and a continental plate; (c) Convergent boundary between two continental plates; (d) Transform fault boundary There are three types of inter-plate boundary (also called margins) (Fig. 2.2):€divergent boundary (where two plates move away from each other), convergent boundary (where two plates move towards each other and ultimately collide) and transform fault boundary (where two plates move past each other). In certain places, three plates meet at a point; this special case is known as the triple junction.

16

Geological framework of the tropical lands

2.2.1╇ Divergent boundary Here two plates move away from each other, allowing an opening to develop on the lithospheric surface between them. The African Rift Valley is an example where plates begin to split. With time, plates on either side of a new rift move away from the break to form a linear sea. The Red Sea is a good example of a linear sea. Eventually, after tens of millions of years, the plates move away far enough to transform a linear sea to an ocean that fills the wide depressed space between the plates, e.g. the Atlantic. Divergent boundaries below oceans are marked by a mid-ocean ridge, actually two parallel ridges separated by a linear depression where lava emerges from below the crust. New molten rock material emerges in the opening from underneath, solidifies, and forms new crust at the edge of the plates as they diverge. Thus new ocean floors are formed and the emerging new material usually solidifies as the volcanic rock basalt. The floor of a rift valley, a linear sea or an ocean is therefore made of basalt of a comparatively young age. The emergence of new material from below is known as lava eruption. Lavas are usually very fluid and erupt easily without any explosive activity. The standard examples of divergent boundaries are submerged volcanic ridges below the oceans. Part of the ridge may emerge above water to form volcanic islands. For example, the mid-Atlantic Ridge runs north–south through the middle of the Atlantic Ocean and locally emerges to form islands such as Iceland and Ascension.

2.2.2╇ Convergent boundary In this case, two plates move towards each other leading to a collision. The convergence can happen between plates of comparable or different densities. In the first instance, as between two continental plates, mostly made of granitic rocks, the former ocean between them is squeezed out before the plates collide. The ocean bottom sediments crumple but do not disappear underneath the crust. Instead they and the ocean floor rocks are folded into mountains, along with emplacement of granitic rocks below and through them. A mountain is thus formed between the two plates, which are now sutured together. The Himalaya is an excellent example. In this process, volcanic rocks of the old sea floor, marine sediments and fossils, and other types of sedimentary rock are folded, squeezed, and thrust over each other. As a result, limestones formed under shallow marine conditions have ended up high on Mount Everest. The thrusts may give rise to large regional faults. Seismic activities may be common. Lines of old collisions and joining (sutures) can be identified even when geologically very old mountains, once very high, have been eroded down to a landscape of small hills and plateaus. In contrast, when convergence takes place between two plates of unequal density, e.g. one oceanic and one continental, the heavy one (oceanic) slides and sinks below the light one (continental) in the form of a wedge. This process is known as subduction. Where subduction takes place, the floor of the ocean is pulled down, forming an ocean-floor trench. Such trenches form the deepest parts of oceans. As the subducted plate sinks, it is stressed because of the friction between the plates, and the episodic release from such friction generates earthquakes. The heat generated from the friction also melts rocks. This molten material, originating from the sinking plate and rising because of the lower density of the

17

A brief introduction to plate tectonics

melt, passes through the top plate. As the hot melt rises, it reacts with the silica-rich continental plate, and the chemical composition of the melt is altered before it emerges on the surface or is stored below. The addition of silica to the molten material makes these melts viscous (sticky) and less fluid than basalt. As they rise, cool and solidify into rigid rocks, the viscosity causes the rocks to stick together to form steep-sided volcanoes. The more silica and alkali there is in their composition, the steeper the sides. The mouth of the volcano (crater) may even solidify for a time, leading to a build-up of pressure underneath, as the melt rises with steam and various kinds of gas:€CO2, SO2, H2S, etc. Eventually, parts of the volcano burst open, breaking the rocks into hot little fragments. This is known as a pyroclastic (fiery fragments) eruption and can be very destructive. Many varieties of volcanic rock can be formed due to different scales of mixing between the materials of the two plates, but the common ‘mixed’ rock is andesite (named after the Andes Mountains, which were formed by subduction). At times, the melt does not emerge on the surface but cools and solidifies inside the crust to form other varieties of rock. If it is silica-rich and coarse-grained from slow cooling deep inside the crust, it is usually granite or a variation of it. Granite cools to form three-dimensional bodies, arching up the surface to form huge, steep-sided, round-topped hills and mountains. The continental plates are granitic in composition. Plate convergence with subduction thus leads to deep offshore trenches, frequent earthquakes and a line of mountains with explosive volcanoes. For example, the Pacific Ocean is ringed by subduction trenches and a line of volcanoes, popularly described as the ‘ring of fire’. Subduction controls the topography. A deep trench, the Java Trench, lies offshore southwest and west of Sumatra, Indonesia where the Indian–Australian Plate subducts below the Sunda Plate. In the Java Trench, the Indian Ocean floor drops to 6000 m below sea level. A line of islands occurs towards Sumatra, parallel to the trench, formed by an accumulation of sediment scraped off the colliding plates and rising magma. This is the fore-arc ridge associated with subduction. These islands are separated from Sumatra by a narrow sea, the fore-arc basin, where the sea floor is pulled down by the effect of subduction. Inland, within tens of kilometres of the west coast of Sumatra rises the long range of the Barisan Mountains with a number of volcanoes, forming the spine of Sumatra and classified as a volcanic arc. East and northeast of this slightly curved mountain range is the low eastern plain of Sumatra, filling a back-arc basin by sediments derived from the Barisan Mountains. The back-arc basin reaches the stable Sunda Plate to the east (Fig. 2.3). Plate convergence and subduction not only explain the topography of the region, but also determine that the region will be affected by earthquakes, volcanic eruptions and tsunamis. The large tsunami of 26 December 2004 eroded back the Aceh coast of northwest Sumatra for about 500 m, removing almost all the beaches, low sand dunes and wetlands behind the beaches. Within a couple of years, new beaches, dune and wetlands were built resembling the pre-tsunami coast (Liew et al., 2008). Large and episodic alterations on the landscape may thus be modified over time by common geomorphic processes. The appearance of the landscape at a point in time depends on how long and how successfully the common geomorphic processes have been operating after the last large episodic alteration like a volcanic eruption or an earthquake.

18



Fig. 2.3

Geological framework of the tropical lands

Subduction-related topography of Sumatra and the islands and ocean to its southwest. See text for explanation New oceanic crust is created at divergent boundaries, but the old oceanic crust is destroyed by subduction at convergent margins. Thus the total crust of the Earth maintains its finite area and, as T. J. Wilson remarked, no part of the floor of ocean is older than 160 million years, whereas the age of the planet is 4600 million.

2.2.3╇ Transform fault boundary Two plates slipping past each other without prominent divergence or convergence give rise to this kind of boundary. Ragged plate edges tend to lock, leading to a build-up of stress

19

A brief introduction to plate tectonics

that is relieved by episodic tectonic movements. Thus strike-slip faults occur near the surface and commonly with a broad shear zone underneath. Most examples of transform fault boundaries occur between oceanic plates, but two land examples with destructive earthquakes have acquired notoriety. The first is the San Andreas Fault in California between the Pacific Plate to the west and the North American Plate to the east. The second example is the North Anatolian Fault in Turkey between the Eurasian and Anatolian Plates. A good example within the tropics is the movement of the North American and Caribbean Plates past each other in opposite directions giving rise to earthquakes, faults and other structural lineaments and landslides. Landforms on the islands of the Greater Antilles display such effects. The magnitude 7 earthquake of 12 January 2010 in Haiti occurred at this transform plate boundary at a shallow depth of 13 km, and along the Enriquillo–Plantain Garden Fault System that runs through southern Haiti, passing near Port-au-Prince (Showstack, 2010).

2.2.4╇ Plate tectonics and landforms Plate tectonics determines the basic template over which geomorphic processes like rivers, winds and slope failures operate. In general, active earth movements do not tend to occur in the interior of plates where the old rocks are. Such stable areas inside continents are called cratons. Parts of cratons that are underlain by metamorphic rocks are shields. If floored by sedimentary rocks, they are called platforms. These are areas where old rocks are close to the surface and the relief is usually not high; the surface is commonly in the form of low, eroded plateaus or hills. Such areas are difficult to weather or erode because of the underlying old, hard, often metamorphosed rocks, and not much sediment is derived from these areas. In contrast, convergent plate boundaries are active areas where tectonism, often manifested in earthquakes, volcanism and mountain building, maintains a high relief and a steep gradient. Simonett, working in Papua New Guinea, estimated that the erosion rate of the Â�earthquake-prone Toricelli Mountains (where frequent landslides are linked with tectonics) is about 100 cm per 1000 years. In comparison, the denudation rate is 22 cm per 1000 years in the neighbouring Betwani Mountains which are less affected by earthquakes (Simonett, 1967). Steep tectonic mountains are prone to slope failures, contributing a large proportion of sediment that travels down major rivers whose headwaters originate in mountains associated with active plate margins. Slope failures in the tropical mountains tend to contribute most of the river sediment, irrespective of the size of the river. Larsen and Torres Sánchez (1992) calculated that 81 per cent of the total sediment transported out of the small Mameyes Basin in eastern Puerto Rico comes via mass movements. The stressed rocks in these mountains adjust by faulting, folding and flowage. Thus steep and high mountains are subjected to frequent slope failures along plains of structural weakness and, furthermore, if such mountains form orographic barriers against moisturebearing winds, floods tend to be common. Even a slight tilting of a river slope increases its energy, causing it to deepen its channel and flow through a gorge. Explosive volcanic flows that emerge from the craters of convergent zone volcanoes end as sediment flows along

20

Geological framework of the tropical lands

river channels draining the lower slopes. As a result, convergent zones are erosion-prone and produce large volumes of sediment. Divergent plate boundaries are not common on the continents, but where they occur, rift valleys are formed with possible volcanic eruptions, as in East Africa. Transform boundaries give rise to strike-slip faults which disrupt pre-existing rivers and ridges giving rise to various types of offset. A linear trough occurs along the fault. As a result, small depressions such as sag ponds and stream diversions are formed along the linear feature. The vertical barriers created by the fault tend to disrupt former drainage lines creating scarps, springs and interrupted river systems. Small ridges create obstacles across drainage lines. These are called shutter ridges.

2.3╇ Major landforms across the tropics A variety of landforms occur across the tropics, depending on their past history. Parts of the tropics consist of older rocks, inherited from the Gondwana. These areas usually form uplands or basins at various levels, generally made of old metamorphosed rocks that are difficult to erode. The Brazilian Shield is a good example. Parts of the old uplands are formed by flood basalts (the Deccan Plateau of India) and granitic bodies (the Main Range of Malay Peninsula) which having undergone considerable weathering and erosion have given rise to flat-top plateaus and steep rounded hills respectively. The movement of plates after the fragmentation of Pangaea gave rise to convergent plate boundaries, which on the active side of moving continental plates created high mountains such as the Andes and the Himalaya. Large structural downwarps have been filled with sediment to form wide plains such as the Amazon and the Ganga valleys. Sediment derived from the higher parts of the continents has built coastal plains over the shallower parts of the continental shelf. A higher variation of landforms occurs at a smaller scale. The structure and rock types therefore induce a range of landforms across the tropics which, given the ambient climate, have undergone and are undergoing appropriate geomorphic processes to produce the tropical landforms we are familiar with. We will go through a very brief account of the geographical distribution of large-scale landforms in the tropics as an aidemémoire for the subsequent material in the book. This is a quick summary and not a detailed description. The active side of South America (the west) is dominated by the spectacular Andes Mountains, formed at the convergent plate boundary between the South American Plate to the east and the Nazca, Cocos and Antarctica plates to the west. The northern and central parts of this long mountain chain fall within the tropical latitudes. The building of the Andes, however, happened over a long period of time starting from the Mesozoic and includes folding, thrusting, crustal thickening, uplift, block faulting, granitic intrusions and volcanic eruptions. The history of mountain building differs along the length of the Andes. The Andes is a long and high mountainous region which includes several lines of ranges, not all of which are continuous, and with the higher peaks in the central region. From west to east, the first line of ranges is the Coastal Cordillera, which rises from a narrow coastal



21

Major landforms across the tropics

Fig. 2.4

Major morphological divisions of tropical South America. From Gansser 1973. By permission of the Geological Society and A. Gansser plain bordering the Pacific. Eastward, this range is succeeded by the Western Cordillera. The Central Cordillera lies towards the north, and the line of the Eastern Cordillera, which comprises elevated but narrow ranges, marks the eastern Andes. A large structural basin between the Western and Eastern Cordilleras has formed a high plateau, the Altiplano. Granitic intrusions and volcanic activities form high peaks. Hundreds of volcanic peaks rise to 5000–7000 m in the high Andes. This is a landscape of high elevation and relief, steep slopes and deep valleys, and tectonically shattered rocks and volcanic material producing a huge amount of sediment that is drained eastward by the drainage systems of the Amazon and the Orinoco. The derived sediment filled a foreland basin created by the topographic load developed by crustal thickening and the eastward extension of the central Andes (Horton and Decelles, 2001). The large alluvial plains slope gently away eastward from the mountains (Fig. 2.4). Beyond the foreland, the Amazon and the Orinoco flow through huge lowlands; the Amazon between the ancient rocks of the Guyana and Brazilian Shields (Fig. 2.5) and the Orinoco between the ancient rocks of the Guyana Shield and the younger ones of the eastern Andes and the mountains of northern Venezuela. It is the large-scale

22



Fig. 2.5

Geological framework of the tropical lands

The old landforms on the low-relief Brazilian Shield with inselbergs. Photograph: A. Gupta structural features and the past geological history that created the landforms and also determined the nature and volume of sediment drained by the two large rivers to the Atlantic Ocean. Northwards, the narrow landmass of Central America is marked by a central mountain system, alluvial fans, small coastal plains and karst. Similar features continue northwards into Mexico, dominated by steep mountains and intermontane basins. Coastal plains are generally limited in area except towards the Yucatan Peninsula, where an excellent development of a karst landscape has happened. Episodic high-magnitude events are common and important for regional geomorphology. For example, a large mudslide following heavy rains devastated hillslopes in southern Mexico on 28 September 2010. The Caribbean islands of the Greater Antilles (Cuba, Hispaniola, Jamaica and Puerto Rico) have central highlands with steep slopes, large areas of karst landscape and coastal plains of varying size. The islands of the Lesser Antilles are smaller and either dominated by a volcano (Martinique) or karst landscape (Barbados). Both features are common. The geomorphology of the Caribbean Islands also needs to respond to episodic events of landslides and floods following the arrival of tropical storms and Atlantic hurricanes. The landforms and tectonics of Africa are different. Africa is distinguished by a mean elevation which is higher than that of other continents. A very large part of the continent consists of high extensive plateaus and large basins. Since the Palaeozoic, Africa has been controlled by crustal extension and rifting. Significant compressional movements have occurred only at the extreme north (the Atlas Mountains) and the extreme south (the Cape Fold Belt). Southern Africa forms the higher part of the continent. A number of very large escarpments mark the edge of the African highlands, e.g. the Great Escarpment of Southern Africa. Highlands above such escarpments are generally characterised by low relief. As a result, internal or centripetal drainage operates over parts of Africa and a low volume of sediment reaches the seas. The aridity of the North (Sahara) and the Southwest (Kalahari, Namib) is also responsible for a low sediment transfer. A considerable volume of sediment originating from the volcanic and rifted

23

Major landforms across the tropics

highlands of East Africa used to travel down the Nile, but anthropogenic impoundments have ended that. Repeated rifting episodes attributed high-relief landforms to parts of Africa (Lambiase, 1989). Two major periods of uplift and rifting should be recognised: a Jurassic–Cretaceous rifting associated with the fragmentation of Gondwana (180–130 million years) and a rifting phase that started 35–25 million years ago and is still continuing (Summerfield, 1991). The East African Rift System has given rise to steep slopes and deep lakes such as Lake Tanganyika. Not all the rifts are associated with volcanism (Lambiase, 1989) but some are. The eastern branch of the Late-Cenozoic East African Rift System and parts of several older rifts are associated with volcanism. Volcanism is not always associated with rifting, as the Neogene volcanic rocks of the Tibesti and Ahaggr Plateaus demonstrate. Parts of West Asia that fall within the tropical zone display a variety of landforms due to the presence of active plate boundaries and variations in annual precipitation. Three active plate margins affect this area. The Arabian Plate is moving away from the African Plate along the Red Sea and Gulf of Aden. The Arabian Plate is also moving northwards and colliding with the Eurasian and Turkish Plates, giving rise to various mountains and basins of West Asia. River systems drain relatively humid regions, but elsewhere huge areas are arid, forming rock and sand deserts and coastal salt flats. South Asia consists of three large physiographic zones:€the Himalaya Mountains, the Indian Peninsula and the alluvial Indo-Gangetic Plain between the two. The Himalaya Mountains consist of a series of asymmetrical ranges with steeper southern slopes, separated by discontinuous elongated valleys. The mountains bend at both the eastern and western ends, where the rivers Brahmaputra and Indus flow through in deep gorges. Two high peaks occur at these bends, the Namche Barwa near to the Brahmaputra and the Nanga Parbat next to the Indus. Between the two peaks the 250–400 km wide mountains extend for 2500 km. The parallel ranges of the Himalaya increase in elevation from south to north; the main range, known as the Great Himalaya or the Himadri, exceeds 6500 m in elevation and includes the highest peak in the world, Mount Everest (8847 m). The Himalaya can be divided into several units according to lithology, tectonics and physiography. The units are divided by regional faults and huge thrusts with varying tectonic activity (Valdiya, 2002). The southernmost range, the 900–1500 m Siwalik Hills are separated from the Indo-Gangetic Plain to the south by the Himalayan Frontal Fault (HFF). The Siwalik Hills are made of Late Tertiary to Early Quaternary sedimentary rocks known as molasses. The next range to the north, the 500–2500 m high Lesser Himalaya, is separated from the Siwalik Hills by the Main Boundary Thrust (MBT). The Lesser Himalaya is made of Middle and Late Proterozoic to early Cambrian sedimentary rocks, huge overthrusts of Early Proterozoic metamorphic rocks from the north and granitic intrusions. The 6500–7000 m high Great Himalaya lies to the north, separated from the Lesser Himalaya by the Main Central Thrust (MCT). It is a complex of high-grade metamorphic rocks and mid-Tertiary granites. India broke away from the part of Pangaea now known as Madagascar about 86–87€Ma and rapidly moved towards the Eurasian Plate. The Himalaya started to rise when the

24

Geological framework of the tropical lands

Indian plate collided with the Eurasian plate, starting in 65 Ma. The suture zone is marked by Cretaceous basic and ultrabasic rocks from the ocean floor. The contact slowed down the northward movement of India and the welding of the two landmasses was completed by 55–50 Ma (Patzelt et al., 1996). It is generally accepted that after the collision, the lower part of the crust under India slid below Asia, probably raising the elevation of the current Tibetan Plateau, and the upper part detached itself ending in an upwarp. Valdiya (2002) described two results of this collision. First, the land immediately to the north of the collision zone was elevated to form a highland (the Kailas–Mansarovar region of southwestern Tibet), from which large rivers drained in different directions. These rivers, the Indus, Satluj, Karnali and Tsangpo (the Brahmaputra in Tibet), persisted in deep gorges through subsequent stages of Himalayan orogenesis in the next 35–45 Ma that created high mountain chains across their flow. Second, crustal sagging created an elongated depression south of the rising mountain chain, which temporarily became a shallow marine basin in which rivers from the north and south deposited sediment. The rise of the mountains also probably intensified the monsoon system (Valdiya, 2002). As Valdiya (2002) described it, the new river systems established themselves on the rising mountain, and the eroded sediment started to travel generally southwards. By 20 Ma ago, the Himalaya formed an orographic barrier and a water divide. Tectonic disturbances, especially on the Lesser Himalaya, continued at intervals. Particularly strong movements at the beginning of the Quaternary generated huge slope failures and a number of debris flows that carried mud and gravel to form a thick boulder conglomerate southwards. The Siwalik Hills were created towards the Late Pleistocene by intense compression of such sediment. A narrow northern part of the pre-existing basin south of the Himalaya was compressed into the rising Siwalik Hills, and its wider southern part subsided to form an elongated basin that was subsequently filled up with sediment to form the Indo-Gangetic Plains (Valdiya, 2002). The Himalaya continues to be a very mobile mountain range with a number of active faults. Levelling across the mountain has indicated that the ranges are rising at an annual rate of several millimetres per year, with the highest value for the Great Himalaya. The denudation rate is also high. Burbank et al. (1996) measured an annual erosion rate of 2–12 mm for the Indus River near Nanga Parbat, a rate that has speeded up over time as determined by cosmogenic isotope dating. Four huge earthquakes with magnitudes over 8.0 occurred in 1897, 1904, 1934 and 1950 in different parts of the mountains. Smaller earthquakes are frequent. The Himalayan Ranges are very large mountains with high relief, steeper southern slopes, varying lithology, shattered rocks and clearly demarcated zones of structural weakness. Furthermore, the mountains are frequently disturbed tectonically. The Himalaya, because of its elevation, was extensively glaciated during the Pleistocene. All the ranges were glaciated in the past and ice still covers the higher elevations. Mass movements are common, especially during rainfall in the wet monsoon. A huge volume of sediment is released that is temporarily stored in valley bottoms and ultimately transported out of the mountains to the Indo-Gangetic Plains (Fig. 2.6). Almost the entire discharge of water and sediment occur during the southwestern monsoon in the summer.

25



Fig. 2.6

Major landforms across the tropics

The Himalaya Mountains and Ganga Plains:€a tectonically active high-relief environment. From NASA/MODIS. See also colour plate section

The Indo-Gangetic Plain lies between the Himalaya and the ancient craton of the Indian Peninsula. As mentioned, it is a structural foredeep filled by the alluvium of three major rivers, the Brahmaputra, the Ganga and the Indus, and their tributaries. The relief is subdued except for ridges of ancient rocks emerging in a few places. A number of alluvial fans soften the contact between the Himalaya and the plains. The Indian Peninsula, which consists of ancient rocks, lies south of the Indo-Gangetic Plain, and is separated from it by a series of sandstone and lava plateaus. A line of hills and plateaus runs parallel to the west coast of the peninsula, marking the drainage divide. This is the Western Ghats, the northern part of which consists of flood basalt (the Deccan Traps) and the southern part of which is made of Archaean gneisses, schists and charnockites (a type of metamorphosed igneous rock). The 1600 km long Western Ghats acts as an orographic barrier to the southwestern monsoon and short steep streams flow down its western slopes to the Arabian Sea. In contrast, longer rivers such as the Godavari, Krishna and Kaveri flow eastward to the Bay of Bengal through a range of disconnected hills of ancient rocks, known as the Eastern Ghats. The only major westflowing streams, the Narmada and the Tapi, flow through rifted valleys towards the north of the peninsula. Further to the east, Southeast Asia is a corner of Asia which starts from a mountainous region and ends in an assemblage of peninsulas, archipelagos and partially enclosed seas. From the mountainous region of the northwest that includes the eastern Tibetan Plateau and the eastern Himalaya Mountains, a number of large elongated river basins

26



Fig. 2.7

Geological framework of the tropical lands

Northern Australia, an old arid landscape. From NASA/MODIS. See also colour plate section

extend north–south or northwest–southeast. These are the basins of the Irrawaddy (or Ayeyarwady), Salween, Chao Phraya, Mekong and Sông Hóng (also known as the Red River). An east–west traverse of the mainland of Southeast Asia therefore crosses a series of alluvial valleys of large rivers separated by mountain chains or plateaus. Tapponier et al. (1986) have explained this pattern as related to an eastward lateral extrusion of Southeast Asia due to the Himalayan orogenesis. Large-scale strike-slip faulting opened up the basins in Southeast Asia. This also created the present drainage systems of the region. It has been suggested that a large pre-orogeny drainage net of the Sông Hóng (Red River) disintegrated (Brookfield, 1998; Clark et al., 2004) due to progressive river capture following the uplift of the Tibetan Plateau and the present separate large river systems (Chang Jiang, Mekong, Salween, Irrawaddy) coming into existence. The Chang Jiang is also known as the Yangtze. The hills and the valleys end in coastal plains, rocky peninsulas or deltas. Beyond are the arcuate islands of Indonesia and the Philippines with steep volcanic slopes, intermontane basins and flat coastal plains, primarily controlled by subduction-related convergent plate margin activities. Australia is a low, arid continent (Fig. 2.7) with a mean elevation of 330 m and about 40 per cent of the land area below 200 m (Goudie, 2002). The relief is low over most of the continent with flat plains dominating the landscape (Twidale and Campbell, 1993). Australian rocks, and even the surface, tend to be aged, a large fraction being derived from the Gondwana. Even certain sand dunes have been reported to be Tertiary in age (Benbow, 1990). In general, a low surface with weathered crusts has developed on the ancient rocks with plateaus and mesas rising from it and inland basins occurring at lower elevations. The

27

Interrelationships

uplands are in bedrock with an association between their forms and rock type. Domes, tors and pinnacles have formed on granite; escarpments, ridges and cuestas on quartzites and sandstones; and dissected uplands on metamorphic rocks. The heart of the continent is mostly stony or sandy deserts, but narrow coastal plains provide a contrast. Aridity, low relief and hard ancient rocks combine to restrict the volume of sediment carried by the rivers of Australia into the sea. This is a very brief summary of the major landforms of the tropics, primarily controlled by plate tectonics. In Part II we examine the various types of geomorphic process that operate in the tropics on this template.

2.4╇ Interrelationships:€plate tectonics, landforms, erosion and sediment production The worldwide delivery of sediment from the land to the ocean via rivers has been reviewed several times (Milliman and Meade, 1983; Walling and Webb, 1983; Milliman and Syvitski, 1992; Meade, 1996; Hovius, 1998). In general, basins with high relief correlate with high sediment production, and section 2.3 helps us to understand the sources and sinks of sediment in the tropics. Milliman and Syvitski (1992) indicated that elevation and relief is possibly a surrogate variable for tectonism. Relief, as measured from elevation data, probably reflects, as they describe it, ‘the entire tectonic milieu of fractured and brecciated rocks, oversteepened slopes, seismic and volcanic activity’ (Milliman and Syvitski, 1992:€540) that occurs in high mountains. Active orogenic belts, positioned by plate tectonics, are eroded heavily, becoming primary sediment sources. This is true for the sediment travelling along large rivers such as the Amazon, Ganga, Brahmaputra and Mekong. The Amazon is over 6000 km long, but more than 90 per cent of its sediment is derived from the Andes Mountains and its foothills and carried across this long distance to the South Atlantic (Meade, 2007). In tropical South America, four major rivers (Amazon, Magdalena, Orinoco and Paraná) collectively drain more than half the entire continent, each with their headwaters in the Andes. The sediment they deliver to the oceans has travelled for thousands of kilometres taking hundreds, perhaps thousands of years. In contrast, the rivers draining west from the Andes are short and steep. Sediment transfer is also affected by the generally dry conditions on the western slopes of the Andes. Significant sediment transfer in these rivers therefore takes place only when large flows occur sporadically, as during the El Niño Southern Oscillation (ENSO). Little sediment is sourced from the lowlands or uplands of ancient rocks such as the Guyana Shield or the Brazilian Shield. This is the common pattern unless the regional lithology is highly soluble (e.g. consists of carbonate rocks or evaporites) or can be eroded very easily (Stallard, 1985; Meade, 2007). The number is high for the east-draining Amazon (about 1500 tkm−2yr−1) but not for the rivers draining west from the Andes. According to Milliman and Meade, the annual

28

Geological framework of the tropical lands

sediment yield from Central America is a little higher than these west-draining rivers, being about 230 tkm−2. The pattern of mountain-derived sediment travelling along short steep rivers in episodic events such as tropical cyclones is found in many rivers of Central America and the Caribbean islands. There the mountains are located near the coast due to plate tectonics. The sediment deposited on a high-gradient coast forms a steep delta resembling a fan, called a fan-delta. A comparatively small amount of sediment seems to be deposited in the surrounding seas by the rivers of Africa (Milliman and Syvitski, 1992). Although the southeastern and eastern part of tropical Africa is high in elevation and bounded by steep escarpments, the rocks are ancient and metamorphosed and not as prone to erosion as the fractured rocks of young folded high mountains. In tropical West Africa, most of the basins of major rivers, such as those of the Senegal and Niger, are not mountainous. The same pattern frequently holds true for small river basins. The Congo drains a huge central basin and then goes over a steep escarpment to flow into the Atlantic Ocean. Although its headwaters drain the steep slopes of the Western Rift, little sediment reaches the sea. The current average annual suspended sediment discharge for the Congo is estimated as 43 million tonnes. Sediment discharged by the Orinoco, which drains a basin less than one-third of the area of the Congo, is about 150 million tonnes (Meade, 1996). In tropical Africa, rivers that carry high sediment discharges are those that drain the high relief associated with rifting in East Africa:€Nile, Zambezi, Limpopo and Rufiji. Even so, many rivers of the world surpass these in sediment discharge. Anthropogenic changes in the channel and river basins may have a greater influence on the amount of sediment load than the natural environment in Africa. The impounded Nile is a good example. In tropical Asia, rivers with high sediment discharge are large rivers that originate in the Himalaya Mountains:€Ganga, Brahmaputra, Irrawaddy, Salween, Mekong, Zhujiang and upper Changjiang, (also known as Yangtze). Short streams that drain the mountainous islands of Indonesia, Philippines and Papua New Guinea also contribute high volumes of sediment given their drainage area. High mountains and volcanoes on these islands are eroded heavily, giving rise to large sediment loads. Small streams draining volcanic islands in the Indian or Pacific Oceans may carry a very high sediment load for their size; the annual contribution from their basins being as high as 3000 tonnes for each square kilometre (Milliman and Syvitski, 1992). The world map of suspended sediment discharge prepared by Milliman and Meade (1983) indicates that the highest annual sediment yield of approximately 3000 tkm−2 is derived from the hilly islands of Southeast Asia and the Himalayan drainage to the Bay of Bengal. Rivers draining the Southeast Asian mainland annually contribute more than 900 tkm−2. In contrast, rivers draining the old landmass of South India were shown to release less than 300 tkm−2yr−1 to the Arabian Sea, a figure comparable to that from East Africa. The rest of Africa erodes at an even slower rate, the estimated volume of annual sediment yield barely reaching 100 tkm−2. Milliman and Syvitski (1992) recognised topography, i.e. tectonics and volcanism, as the prime explanation of high sediment production, being more crucial than precipitation

29

Questions

and runoff, although the latter are also important factors. Thousands of years of human occupation of land has, however, significantly changed the natural rate of sediment production in two ways. Erosion and sediment production have increased dramatically with deforestation and farming. Saunders and Young (1983) estimated that such increases range from 2–3 times (in what they have described as moderately intense land use) to about 10 times or more with intensive land use and associated soil erosion. On the other hand, the multiplicity of dams and reservoirs across the world results in the storing of a huge amount of sediment in the reservoirs, depriving downstream rivers of their sediment load. It is a safe generalisation that rivers with high sediment load originate in the mountains. Most of the large rivers originate in mountains associated with active plate margins (front edge of a moving plate) but discharge to the sea at passive margins (trailing back edge of a moving plate), building big deltas with their high sediment load. If the mountains (volcanic or otherwise) are located close to a coast, a number of small rivers drain their steeper shorter side. Similar streams drain mountainous islands. As many of these short rivers flow into an active plate margin, their sediment deposits may be subducted and deposition forms at their mouth could be limited in size, perhaps forming a small, steep fan delta. This pattern, however, may change if these slopes are affected by large storms, volcanic eruptions or earthquakes. Such large events may bring down a huge amount of sediment that needs a number of years to be reorganised. Erosion and sedimentation depend on a range of environmental aspects:€ tectonics, relief, climate, land cover, anthropogenic modification of the land and geomorphic processes. We look at some of these aspects briefly in the next two chapters before moving on to a detailed discussion of the operation of geomorphic processes in the tropics.

Questions 1. Which parts of the tropical world originate from Gondwana? Which parts do not? 2. What surface features should you expect on a craton? What would be a reasonable Â�volume of sediment eroded from a craton? Sediment is commonly measured in tkm−2yr−1. It is known as the sediment yield. 3. A large volume of sediment is usually derived from high steep mountains forming orographic barriers to moisture-laden winds. Name any five locations in the tropics where it happens. 4. List (A) the principal source areas of sediment in the tropics, (B) the large rivers (at least 1000 km in length) that transfer high volumes of sediment to the sea, and (C) where the sediment is finally deposited by these rivers. 5. Examine Figure 2.6 which is a MODIS (Moderate-Resolution Imaging Spectroradiometer) image taken from space. (A) Identify the expected areas of high erosion. (B). Which are the locations where large amount of sediment is expected to be stored?

30

Geological framework of the tropical lands

6. Compare the sediment-producing capacity of the land shown in Figure 2.7 with Figure 2.6. 7. Milliman and Syvitski (1992) identified tectonics and volcanism as the prime factors behind high sediment production. Do you agree? 8. Is it reasonable to expect the current rates of erosion and sedimentation to be natural rates?

3

Tropical hydrology

A set of ‘rainfall rules’ to apply in Africa has been suggested by some cynic, perhaps a tropical meteorologist but certainly a person with experience in the tropics. These rules are:€(a) the rainfall is seasonal in nature; (b) the amount increases as one approaches the equator; (c) do not put too much faith in (a) and (b). J. F. Griffiths

3.1╇ The tropical climate:€a brief review Climatic characteristics do affect local geomorphology in the tropics, although not as deterministically as was presented by many geomorphologists in the past. The warm and moisture-laden condition of the humid tropics certainly accelerates the process of weathering, and the high-intensity rainfall of the tropics gives rise to frequent slope failures and flooded rivers. In contrast, physical weathering and shorter rare floods from cloudbursts are common in the arid tropics. This chapter is not on the principles of climatology but reviews the basic nature of the tropical climate with special reference to the factors that influence tropical hydrology and geomorphology (for a detailed review of tropical climate please see McGregor and Nieuwolt, 1998).

3.2╇ Temperature Climate in the tropics varies according to location. In 1918, the Austrian meteorologist Köppen characterised the tropics as the region where the average temperature of the coldest month is above 18°C. This would be the A-type and part of the B-type in the Köppen– Geiger classification of world climates (Trewartha, 1954). But the tropics also include highlands and high mountains where the temperature certainly drops below this designated threshold, and seasonal differences in temperature with a clearly defined winter become perceptible at latitudes higher than about 10°. Near the equator a thermal uniformity prevails with year-round high temperature at low altitudes. Here the diurnal range in temperature is usually larger than the annual range. Cloudiness, wherever it occurs, tends to reduce the diurnal range. If rainfall is strongly seasonal, as it usually is at latitudes above 10°, temperature may drop during the wet season due to both cloudiness and rain. Köppen described this pattern as the Ganges-type temperature distribution, where the temperature 31

32



Fig. 3.1

Tropical hydrology

Distribution of average annual precipitation in the tropics rises to the highest level in early summer, then drops at the summer solstice which falls within the rainy season, and rises again to a post-rain secondary peak before the arrival of a cool mild winter. Both the annual and diurnal ranges of temperature are high in the arid tropics. The annual range of temperature is as much as 20–30°C, higher than elsewhere in the tropics. The annual range may even equate the diurnal range, although the latter is usually large.

3.3╇ Wind circulation From the lower latitudes, where there is surplus solar radiation, heat is dissipated to the higher parallels by wind. This is seen as a three-dimensional circulation known as the Hadley Cell, a generalised version of which is shown in Fig. 1.1. The surplus heat energy of the solar radiation is transferred to kinetic energy near the equator, and also evaporates moisture from the ocean. The heated air mass rises with high moisture content. As it rises, condensation forms clouds, and in some instances rain, releasing part of the latent heat, which takes part in further warming and lifting of the air mass. This rising mass of air is then carried to higher latitudes by upper air currents. The upper air currents slow down around 20–30° of latitude. This is followed by subsidence, building of regional highÂ�pressure zones (the subtropical highs) and air mass warming. As a result, drier atmospheric conditions prevail around 30° latitudes where the arid tropics are located. At low atmospheric levels, a large and steady movement of air takes place from the subtropical highs towards the equator thus closing the circulation (Fig. 1.1). As moving air tends to turn to the right north of the equator and to the left south of it, this low-level wind steadily blows over the surface of the tropics from the northeast or southeast (depending on the hemisphere), and is collectively known as the trade winds. The trade winds are especially steady over the ocean, bringing moisture westward to land surfaces. The northeast and southeast trade winds converge at a belt of low pressure round the globe near the equator, known as the Inter Tropical Convergence Zone (ITCZ). This completes the circulation of the Hadley Cell and contributes to the lifting of moisture-bearing air at the ITCZ to higher altitudes, leading to equatorial rainfall. The ITCZ shifts its position slightly between the seasons. The Hadley Cell does not operate quite as simplistically as described here but this account provides a basic understanding of the general movement of moisture

33

Wind circulation

in the tropics. Oceanic islands with high volcanoes are therefore wetter on their eastern side, which is the direction from which the moisture-bearing trade winds arrive, and much drier on the west. The Hawaiian Islands are good examples. Over the ocean, the trade winds usually operate throughout the year with minor shifts in latitude and small variations in intensity. In contrast, there is a seasonal rhythm over the continents and coastal waters, which, being warmer in summer and colder in winter, influence a seasonal variation in wind direction, particularly in the higher tropics. This is a gigantic manifestation of land and sea breezes. In winter, subtropical high-pressure cells develop over the cooler continents, which in summer are replaced by centres of low pressure over heated landmasses. This gives rise to a seasonal reversal of wind systems between land and sea, leading to alternate dry and wet seasons, and winds characterised by this direction reversal are known as monsoons. There are always two monsoons coming from different directions at different times of the year. Monsoon systems of reversal are seen over northern Australia, tropical Asia (South, Southeast and parts of East Asia), East and West Africa, and parts of southwestern North America. The Asian monsoon is influenced by two factors:€(1) the seasonal thermal contrast over land and (2) the east–west alignment of the high Himalaya Mountains. In the winter season, winds near the surface flow round the Himalaya and the Tibetan Plateau, moving eastward towards the coast of East Asia where they meet the moisture-carrying trade wind. The combination turns southeast to bring rain over the eastern coasts of Indochina and Malay peninsulas. This is the northeastern monsoon. When the wind systems reach Indonesia, further south, and the equator, they are well saturated, warm and unstable, and continue to produce rain. On crossing the equator, these winds change direction and blowing from the northwest bring seasonal rainfall to Northern Australia. In contrast, winds over Northern Australia come from the southeasterly direction during the southern hemispheric winter. Rainfall over Northern Australia is seasonal and individual disturbances with embedded cells of high rainfall arrive in the wet season, causing bursts of heavy rain. This is a common characteristic of the wet monsoon, the monsoon that brings rain. Over Asia, the summer monsoon arrives at different times in different places. It is a deep circulation, reaching 6000 m over India and even higher over Myanmar (Nieuwolt, 1977). Most of the rain that falls over South Asia, western coasts of Southeast Asia and East Asia arrives during the summer monsoon. The Indian Ocean and the Western Pacific are the main sources of moisture. For example, moisture-bearing winds from the East China Sea move inland up the Chang Jiang (Yangtze) Basin, bringing monsoon rain inland progressively later. Rainfall from the moisture-bearing air is triggered by a number of factors:€orographic lifting, convergence and atmospheric disturbances. The disturbances travel embedded in the monsoon wind system. The passage of a depression is followed by a slackening of rain for several days until the next depression arrives. This type of pulsating nature is a characteristic of the monsoon system of rainfall. It is also reflected in river flows. Rain-bearing depressions arrive embedded in the monsoon system of winds during the wet season. For example, the east coast of the Indian Peninsula is subjected to the passage of a series of depressions between June and September. Das (1968) mentioned that on average, eight cyclonic depressions pass from the Bay of Bengal to the eastern coast of India between June and September. Widespread rain occurs in the southwestern quadrant

34

Tropical hydrology

of these depressions (Rao, 1981). The rain is periodically concentrated by higher-intensity precipitation pulses. Sikka estimates that 8–10 storms arrive in an average season, with a lifetime of 2–5 days. Monsoon storms of up to 9 days' duration have been recorded. Such storms tend to be 1000–1300 km across, with a 7–9 km deep cyclonic circulation with an anticyclonic outflow on top (Sikka, 1977). Arriving in the middle of the wet monsoon, when the soil is already saturated and rivers high, these high-rainfall events often create floods and failures on hillslopes. Tropical disturbances arrive in a similar fashion over the Indochina Peninsula, usually late in the summer monsoon when the ground is wet and the rivers are high. The African monsoon is not as deep as the Asian version. In the northern hemispheric winter the ITCZ moves close to the equator and dry stable northeasterly winds arrive over West Africa, frequently carrying dust particles from the Sahara Desert. In contrast, during the northern hemispheric summer, high temperature and a thermal low build up over West Africa, and the ITCZ moves northwards to about 15°N. South of it, warm moist southwesterly winds from the Atlantic Ocean bring rain. In East Africa both monsoons arrive from the east, and whether they are locally northeasterly or southeasterly depends mostly on the seasonal position of the ITCZ. The winds, however, are relatively dry and either convergence or orographic uplift is required for precipitation. Thus, the rainfall over Africa is not uniformly distributed. For example, the eastern mountain slopes of Madagascar, first reached by southeasterlies from the Indian Ocean in the southern hemispheric summer, receive higher rainfall than the continent of Africa (Nieuwolt, 1977). Orographic conditions cause a general increase in rainfall. Seasonal wind reversals and wet and dry differentiations also occur near the coast of the southwestern United States and western Mexico, known locally as the monsoon. The arrival of the wet monsoon and the amount of precipitation both vary each year. Various explanations, such as control by upper-air jet streams or pressure oscillations in El Niño or La Niña years, etc, have been offered for such variations. Geomorphologically such variations lead to a high number of slope failures and floods in the wetter years and less effective erosion and sediment transfer in drier ones.

3.4╇ Precipitation Average annual total precipitation over the tropics is shown in Fig. 3.1. The amount of rainfall decreases with distance from the equator unless local factors, such as windward location on orographic barriers, override this general rule. The ITCZ remains close to the equator throughout the year producing high annual rainfall with two maxima separated by two relatively less wet periods (Fig. 3.2). From 10–15º of latitude, the pattern changes progressively to one rainy season and a long dry period, although the annual rainfall remains high. At higher latitudes in the tropics, the rainy season becomes shorter, the amount of rainfall drops, and at around 25–30° the influence of the subtropical high-pressure cells becomes dominant. Semi-arid and arid tropics appear at these locations. Rainfall intensities are commonly high in the arid tropics, although the duration of rainfall is usually

35



Fig. 3.2

Precipitation

Examples of variations in annual temperature and rainfall in the tropics. Entebbe (Equatorial highland), Jodhpur (arid subtropics), Mumbai (monsoon) and Townsville (seasonal climate, southern hemisphere). From McGregor and Nieuwolt, 1998. By permission of Wiley brief, lasting only for tens of minutes. Over the oceans, the latitudinal pattern of rainfall is clearer than over the continents, where convection and orographic lifting add complexity to the general pattern. Even over the oceans, heavy rainfall from disturbances such as tropical cyclones and easterly waves (discussed below) produces deviations from the general model. Rain falls over the tropics with high intensity, causing slopes to fail, floods to occur and land to erode unless it is covered with vegetation. Even on a daily basis, the mean rainfall intensity may reach figures in tens of millimetres. The short-term intensity is much higher (Table 3.1). A high percentage of this rain comes in short-duration tropical disturbances which cause continuous and intense rain while they last. Tropical disturbances that may last for several days bring in even higher amounts, always with strong intensity. To illustrate, rainfall over southern Mexico in late September 2010 exceeded 700 mm in two days,

Tropical hydrology

36

Table 3.1╇ Examples of recorded very heavy one-day rainfall in the humid tropics Date (if known)

Location

Amount (mm)

15–16 Mar. 1952

Cilaos, Reunion

1870

10–11 Sept. 1963

Pai Shi, Taiwan

1248

Typhoon

14–15 Jul. 1911

Baguio, Philippines Bowden Pen, Jamaica Cherrapunji, India Funkiko, Taiwan Bowden Pen, Jamaica Silver Hill, Jamaica Cinchona, Jamaica Rochambeau, Fr. Guinea Nkhotakota, Malawi Mumbai, India Jaffna, Sri Lanka Zomba, Malawi Diego-Suarez, Madagascar

1168

Typhoon

1109

Frontal

1036

Monsoon depression

23 Jan. 1960 14 Jun. 1876 31 Aug. 1911 22 Jan. 1960 6 Nov. 1909 25 May 1889

Santo Domingo, Dominican Rep. Darjeeling, India Dehradun, India Bhuj, India Akyab, Burma Tana Tave, Madagascar

Cause (if known)

1034

Source Landsberg, footnote in Flores & Balagot (1969) Landsberg, footnote in Flores & Balagot (1969) Flores & Balagot (1969) Vickers (1967) Das (1968) Jennings (1950)

977

Frontal

Vickers (1967)

775

Incipient hurricane Mixed

Vickers (1967)

718

Vickers (1967)

596

Snow (1976)

570

Torrance (1972)

548 520

Rao (1981) Rao (1981)

509 508

Torrance (1972) Griffiths & Ranaivoson (1972) Portig (1976)

508

493

Monsoon depression

Rao (1981)

490 468 465

Dhar et al. (1984) Rao (1981) Nieuwolt (1981)

442

Griffiths & Ranaivoson (1972)

Tropical disturbances

37

Table 3.1 (cont.) Amount (mm)

Date (if known)

Location

10 Apr. 1955

Rungwe, Tanzania Vango, Kenya

Oct. 1966

Cause (if known)

425 404

Source Temple and Rapp (1972) Griffiths (1972)

Note:€This table includes only rainfall over 400 mm in a day. Examples of lower figures are too common to be listed. Figures are rounded to the nearest millimetre. Some figures are one-day rainfall, while others are rainfall recorded in any 24-hour period. In most cases rain probably continued for more than a day. Generalised from Gupta, 1988

with about 400 mm falling in a single day. Rainfall from the super-typhoon Megi (October 2010) rose to almost 1200 mm in 48 hours in Taiwan. Such rain is an effective erosive agent, especially without ground protection. Deforestation removes the protection given by dense vegetation to the slopes of the humid tropics from such heavy and intense rainfall.

3.5╇ Tropical disturbances Temporary low-pressure systems of varying size, duration and intensity are superimposed on the general atmospheric circulation system in the tropics. These disturbances cause intense and high rainfall which could be associated with surface runoff, floods and slope failures. Locally, cells of convection may give rise to thunderstorms, which are the smallest of the tropical disturbances. Thunderstorms also develop when two air masses converge, forcing the convergent air to rise and condense. A number of separate thunderstorms may occur along a strong zone of convergence, forming a linear system of squalls often recognised as a line of discrete cumulonimbus clouds rising into the air. Such linear systems may form due to the convergence of sea breezes over a narrow neck of land or orographic lifting along mountain crests or coastlines. These squall lines could cover hundreds of kilometres in length and 10–30 km in width, including zones of both strong and weak thundershowers. In West Africa, linear thunderstorms may produce a significant amount of rain at the two extremities of the wet monsoon season. They often arrive in the pre-monsoon summer in northeastern India and Bangladesh. The seasonal sumatras of southwestern Malaysia and Singapore are a regionally important linear system of storms. They originate over the eastern lowlands of Sumatra but tend to reach the southwestern coast of the Malay Peninsula in early morning bringing copious rainfall (Nieuwolt, 1977). They comprise a 200–300 km long band of cumulus and cumulonimbus clouds (McGregor and Nieuwolt, 1998). Over parts of the tropics, thunderstorms may occur frequently at any time of the year, producing a large proportion of local rainfall. Thundershowers that instead of travelling become stationary produce even heavier rainfall. For example, 390 mm of rain fell in three days in the Virgin Islands in May 1960 resulting in floods (Portig, 1976). Henry

Tropical hydrology

38

Table 3.2╇ Some very heavy multi-day rainfall in the humid tropics Date

Location

Amount (mm)

11–19 Mar. 1952

Cilaos, Reunion

4130

9–16 Jun. 1876

Cherrapunji, India Cherrapunji, India Bowden Pen, Jamaica Cinchona, Jamaica Baguio, Philippines Funkiko, Taiwan Tacajo, Cuba Silver Hill, Jamaica

3388

24–30 Jun. 1931 22–25 Jan. 1960 4–11 Nov. 1909 14–20 Jul. 1911 18–20 Jul. 1913 3–8 Oct. 1963 5–7 Oct. 1963

Cause

Monsoon depression

3213

Source Landsberg, in Flores & Balagot (1969) Jennings (1950) Jennings (1950)

2789

Frontal

Vickers (1967)

2287

Hurricanes

2210

Tropical storm

Jamaica Weather Reports (1909) Jennings (1950)

2071 2025 1524

Tropical storm Hurricane Flora Hurricane Flora

Jennings (1950) Vickers (1967) Vickers (1967)

Note:€Generalised from Gupta, 1988

(1974) examined tropical rainstorms (defined as precipitating at least 10 mm of rain) over parts of Southeast Asia and tropical America. He expected a ‘typical tropical rainstorm’ to cover at least 30 km in diameter. But within it, rain falls with variable intensity in time and space. Thus rainfall distribution within the area covered by these storms can be patchy, with inter-storm rain gauges receiving little or no rain. The rainfall at the core of a typical storm exceeds 25 mm, and more than 60 per cent of the daily rainfall is derived from one storm. Other studies confirm this picture of a short heavy downpour, in which most of the rain falls within the first 30 minutes or so, although the storm may tail off slowly. More than 200 thundershowers a year can be expected over much of Africa (Griffiths, 1972). Such storms are important in geomorphology because they can cause intense soil erosion. According to Hudson (1971), the erosive power of tropical rain is about 126 times more than from rainfall in the temperate areas. The largest tropical disturbances are tropical cyclones (Box 3.1). Very high-intensity rain (Table 3.2) has continued for a day or more in the tropical cyclones described below. These are impressive high-magnitude events that have caused some of the world’s heaviest rainfall, especially when tropical cyclones have stalled or confronted orographic barriers. Such heavy rainfall usually causes numerous slope failures and high-magnitude flooding. Typhoon Morakot in early August 2009 produced rainfall in excess of 600 mm over nearly the entire southern half of Taiwan. Certain areas received more than 1000 mm of rain on the western slopes of the central mountain range, giving rise to massive flooding, devastating mudslides, and huge losses of life and property. The geomorphic effects of tropical storms including the cyclones are discussed in Chapter 7 and also their common locations.

39



Box 3.1

Tropical disturbances

Tropical cyclones A tropical cyclone is a very large low-pressure system that occurs in the tropics and subtropics and has a revolving wind velocity of at least 119 kph (33ms−1). Clouds that develop with this system resemble a funnel with a clear, circular and relatively calm centre, known as the eye. The diameter of the eye ranges between 5 and 60 km. Very low pressure is found in the eye, the lowest value recorded falls in the range 860–890 mb (Simpson and Riehl, 1981), which contrasts remarkably with normal surface atmospheric pressure of 1013 mb. Generally, the pressure in the eye of a tropical cyclone is around the 950 mb level. Rapidly circulating winds occur round the eye with the highest velocity about 1–2 km from its boundary. In most tropical cyclones, a vertical wall of cloud develops round the eye, which rises into the top of the funnel about 5–10 km above the surface and then curves outwards, identified as a vortex in satellite images (Fig. 3.3). This huge vortex revolves counterclockwise in the northern hemisphere and clockwise in the southern. The extent of a tropical cyclone could be about 1000 km. Generally, two arms of clouds extend out further as long spiral bands. The outflow at the top is balanced by a strong inflow of air at the lowest levels. Tropical cyclones develop from persistent but travelling rain disturbances over warm oceans which have temperatures of 26.5–27°C. Hence they develop only in summer, particularly in late summer when the sea surface temperature (SST) is especially high. Two factors are essential for their formation:€Coriolis force and latent heat. As the Earth’s Coriolis force is zero at the equator, tropical cyclones are not formed within 5° north and south of the equator. Once the system is seeded and pulling in evaporated moisture from the ocean surface, it is driven by the energy released by the latent heat of condensation. The growth of tropical cyclones is also helped by (a) the presence of an anticyclonic circulation in the upper troposphere that facilitates outflow at the top of the rotating system and (b) absence of a vertical wind shear allowing development of the tall cumulonimbus convection (Fig. 3.4). The small disturbances or lows that may develop into a tropical cyclone usually come from small vortices near the ITCZ or easterly waves (McGregor and Nieuwolt, 1998). These summer storms form over tropical oceans except the South Atlantic where the cooler temperature, northerly position of the ITCZ and vertical wind shear restrict their growth. Table 3.3 summarises their estimated annual frequency over various oceans. Tropical cyclones are known under different names in different regions:€hurricanes in the tropical Atlantic,

Fig. 3.3â•… Tropical cyclone Katrina over the Gulf of Mexico, 2005. From NASA/MODIS

Tropical hydrology

40



Box 3.1

Continued

Fig. 3.4â•… Diagram of a tropical cyclone. From Simpson and Riehl, 1981. By Permission of Wiley/Blackwell

Table 3.3╇ Estimated annual frequency of tropical cyclones Ocean of origin

From Simpson and Riehl (1981)

From Reynolds (1985)

Northwest Pacific Northeast Pacific Australian seas Northwest Atlantic Southern Indian Ocean Northern Indian Ocean South Pacific

22 15 5 8 5 8 5

26 13 10 9 8 6 6

Note: Generalised from Gupta, 1988

Caribbean, and Northeast and South Pacific; typhoons in the Northwest Pacific; and tropical cyclones in the Indian Ocean and Australian seas. These huge revolving storms tend to travel from their area of origin, embedded in the trade winds. They may either reach a landmass where they tend to dissipate after destructive activities on the coastal areas or follow a near-parabolic path circumventing the subtropical anticyclones over the oceans to reach the westerly winds of temperate latitude and curve back eastward as extra-tropical storms. The wind strength of a tropical cyclone varies across the system, possibly as a result of the vortex and the steering air current such as the trades. As a result the front right quadrant of the tropical cyclone, where the directions of both coincide, has the strongest wind. The clouds are seen as long streets spiralling inward from the outer edges. A tropical cyclone is capable of producing huge amounts of rainfall (Tables 3.1 and 3.2). This rain mostly comes from the part of the cyclone where wind velocity is high and thus is difficult to measure with conventional rain gauges. It is better determined by numerical computations or from satellite radar images. In general, precipitation from a tropical cyclone shows a logarithmic distribution when the amount is plotted against the distance from the eye, with the maximum concentration of the rain just outside the core. However, the absolute volume of rainfall is very high even on the periphery because of the huge size of the storm. Many

Tropical disturbances

41

Table 3.4╇ The Saffir–Simpson damage potential scale Category of tropical cyclone

Central pressure (mb)

Maximum sustained winds (kph)

Storm surge (m above normal)

1

≥ 980

119–153

~ 1.5

2

965–979

154–178

~ 2–2.5

3

945–964

179–209

~ 2.6–3.9

4

920–944

211–249

~ 4–5.5

5 (called supertyphoons in W. Pacific)

>920

>249

>5.5

Damage Damage to shrubbery, trees and unanchored mobile homes. Low-lying coastal roads inundated Considerable damage to trees, some damage to building roofs, coastal roads flooded, possible damage to moored shipping in unprotected sites Large trees down, small buildings damaged, serious flooding at coasts and low areas, structural damage near coast Trees down, extensive damage to buildings, structures near coasts damaged, coasts and low areas flooded, beaches eroded Trees down, extensive damage to buildings, roof failures, small buildings overturned, extensive flooding

Generalised from Simpson and Riehl, 1981

tropical cyclones therefore generate rainfall over a substantive part of a river basin and create huge floods. If the tropical cyclone slows down over a mountainous area, a huge amount of cumulative rainfall occurs over several days, as happened from Hurricane Flora in 1963 in Cuba and Typhoon Morakot in 2009 in Taiwan. In geomorphology, tropical cyclones are considered high-magnitude low-frequency events that are capable of causing large landform alterations. Such changes occur due to the strong wind, high rainfall and storm surges associated with them. They are classified by their properties and ability to damage natural and anthropogenic features according to the Saffir–Simpson Damage Potential Scale (Table 3.4). As tropical cyclones make landfall, wind drives huge waves inland and the super-elevation of such waves during high tide levels are known as storm surges. These are extremely destructive on beaches, barrier islands and dunes. The level of the storm surge rises in bays and estuaries as these become narrower inland causing even greater damage. The wind round the eye, and especially in the front right quadrant of the storm, is very high and

42



Box 3.1

Tropical hydrology

Continued capable of destroying forests, individual trees, buildings and other structures. Apart from the sustained wind speed of the storm (average wind speed over time, usually measured in minutes), gusts (high instantaneous wind speed) also cause damage. When hurricanes or typhoons confront slopes, the destruction of vegetation by wind is followed by huge rainfall, a combination that leads to extensive slope failures, gullies and river floods (Chapter€7). Tropical cyclones lose their power after crossing the coastline, as they are no longer able to evaporate water from the ocean surface, a principal source of their energy. They also decay from friction as they cross mountain barriers across their track. These, however, are times when huge amounts of precipitation may occur, leading to large-scale flooding. Tropical cyclones thus are effective geomorphologic agents in parts of the humid tropics. The Atlantic tropical cyclones were given women’s names from 1953. The system has been modified. At present men’s and women’s names alternate, in alphabetical order. There are six standing lists of names which are used in rotation. For example, the list for 2011 hurricanes runs Arlene, Bret, Cindy, Don, Emily, etc. This list will be used again in 2017. If a hurricane is particularly destructive, its name is retired and replaced by another one starting with the same letter. Tropical cyclones for every ocean have their own special set of names. A complete list is available from the National Hurricane Center (USA) website:€http://www.nhc.noaa.gov.

They are not only powerful but also expected. Over a period of 70 years between 1884 and 1953, 22 storms arrived on average each year in an area bounded by 5°–30° N latitudes and 105°–150°E longitudes, an area off the coasts of South China, the Philippines and Vietnam (Nieuwolt, 1981). In the 70 years between 1891 and 1960, a total of 360 tropical cyclones arrived over the Indian subcontinent, averaging 6 per year. About 70 per cent of these storms were between June and November, when the ground was already saturated and the rivers high. These tropical cyclones repeatedly cause slopes to fail and rivers to flood, thereby strongly affecting local landforms and geomorphic processes. Perturbations in the trade winds may show a wave in the pattern of isobars forming a low-pressure trough at right angles to the direction of the trade wind. These are known as easterly waves as these waveforms move with the easterly trades (Fig. 3.5). The waveform leads to convergence followed by rain, which may reach a large amount if the convergence takes place near an orographic feature. Heavy rainfall arrives from time to time from easterly waves in the Caribbean and eastern Philippines along orographic barriers. Tropical cyclones may originate from easterly waves as they move across the Atlantic, from the African shores towards the Caribbean. Expectations of high-rainfall events have been computed for certain regions. For example, the 5-year 24-hour rainfall over an area of 280 km2 in eastern Jamaica was estimated to be between 350 and 600 mm (Lirios, 1969). Vickers (1967) estimated nearly 450 mm for a maximum 24-hour rainfall for a small watershed in southeastern Jamaica. If these figures are representative, then episodic rainfall of very high magnitude recurs in the Caribbean and adjoining areas at intervals of several years. Other areas of seasonal humid tropics may carry similar figures. High-magnitude storm rainfall may be expected

43



Fig. 3.5

Miscellaneous factors

Generalised diagram of an easterly wave. Rain falls from the shaded area at short intervals within the geographical location of 10 and 30 degrees of latitude, where tropical cyclones are common in summer. Within such a latitudinal belt, if orographic barriers run across the common paths of these depressions, as reported from the Philippines, Taiwan, eastern Jamaica, Puerto Rico and Madagascar, rainfall can be especially high and intense.

3.6╇ Miscellaneous factors The general character of climate at a particular location in the tropics may alter for a brief period of time due to the impact of factors such as the release of aerosols in the atmosphere by volcanic eruptions or the El Niño Southern Oscillation. Large eruptions may release a high amount of volcanic material into the atmosphere. The coarser fragments return to Earth within a short time, but the finer parts stay much longer in the atmosphere as aerosols. Their presence influences various climatic properties during this period. If the eruption is very large, the temperature of the next year may be reduced and the nature of the rainfall changed. These may have some minor influence on geomorphological processes, especially as the volcanic ash travels a long distance worldwide. The June 1991 eruption of the Pinatubo volcano in Luzon, the Philippines, created a stratospheric aerosol cloud that had travelled round the Earth by mid-July. The specific eruption that took place on 15 June 1991 was estimated to have released about 20–22 km3 of ash. The temperature of even the temperate countries fell below average along with increased storminess in the year after the huge 1815 eruption of Tambora in Sumbawa, eastern Indonesia. The El Niño Southern Oscillation (ENSO) is a disturbance that occurs periodically, altering the rainfall and storminess across the tropics. This is related to a change in pattern, known as the Southern Oscillation, between sea-level barometric pressures over the eastern tropical Pacific and the western tropical Pacific and Indian Oceans. In a majority of years, winds blow from the high-pressure area off the west coast of South America across the Pacific Ocean towards the low-pressure area off East Asia. Blowing across the ocean, this wind generates surface

44

Tropical hydrology

currents, which travel in the same direction as the wind. Surface water off the Peruvian coast thus is pushed westward and deeper colder water comes up from below to fill the gap. The process is known as upwelling. The upwelling water is not only cold but also rich in plankton which attracts fish. The coast of Peru therefore has a cool and dry climate, while warm water driven by wind across the equatorial Pacific piles up on the Asian side of the ocean. The atmospheric pressure difference between the East and West Pacific disappears in certain years. Then the east-to-west wind across the ocean weakens, and the surface water that was piling up against the coast of East Asia tends to flow back east across the Pacific. If this condition is fully developed, warm water reaches the coast of Peru, altering the regional climatic condition (making it warmer and drier) and also reducing the fish catch. Such a condition is locally known as El Niño. The phenomenon has been named after infant Jesus by the fishers of Peru, as it tends to develop towards the end of December. The years of El Niño strongly correlate with years of drought in Australia, parts of Southeast and South Asia and Africa. In contrast, El Niño years could be times of heavy rains and floods on the west coasts of the Americas. Such related environmental behaviours across vast distances are known as teleconnections. El Niños tend to recur after several years (the average time difference between two El Niños being 3–7 years), and these conditions last for 1–2 years. El Niños vary in strength. The El Niños of 1982–83 and 1997–98 were particularly strong, bringing floods, landslides, forest fires and death and destruction across many parts of the tropics. The 1972–73 El Niño was remembered for severe droughts in Africa, especially in the Sahel area around the southern boundary of the Sahara Desert. It has been noticed recently that the atmospheric change that promotes conditions for El Niños to form may occur in certain years at the Central Pacific rather than off the coast of Peru. Such El Niños are known as the Central Pacific El Niños. In certain years, the departure from the norm happens the opposite way and cold water flows east across the Pacific. This is the year of La Niña. During La Niña, rains and floods may threaten parts of Australia, Asia and Africa and the coastal waters off Peru could be especially cold. All these cause parts of the tropics to undergo accelerated dry and wet conditions periodically, which not only affect the economic and social conditions but also the geomorphic processes by changing the hydrological norms. Aalto et al. (2003) have shown that major depositional events related to the floodplains of the Beni and Marmore Rivers in northern Bolivia, tributaries to the Madeira, have a recurrence interval of about 8 years which corresponds to the La Niña years. Periodic slope failures and flooding in parts of the tropics therefore occur, interrupting the geomorphological processes which are associated with the expected climate of a location. A strong linkage exists between episodic meteorological phenomena and geomorphology.

3.7╇ Water balance Water balance, a term first used by the climatologist C. W. Thornthwaite in 1944, determines an accounting relationship between the incoming and outgoing water for a soil profile or a drainage basin. The incoming water could be from various forms of precipitation

Water balance

45



Fig. 3.6

Pattern of water moving through soil

and snowmelt. The outgoing water is the sum of evapotranspiration and overland flow (in case of a soil profile) or streamflow (in case of a drainage basin). The positive difference in volume between the incoming and outgoing water is usually stored in the soil; it raises groundwater level or enhances streamflow (Fig. 3.6). The negative difference leads to dry soil and low streamflow. It can be written as

P + Sn = AET + SF (or Q) + ∆SM + ∆GWS + GWR

(3.1)

where P = Precipitation Sn = Snowmelt (on high mountains in the tropics) AET = Actual evapotranspiration SF = Surface flow Q = Discharge ΔSM = Change in soil moisture ΔGWS = Change in groundwater storage GWR = Groundwater runoff Water balance (also called water budget) is based on the two important criteria of precipitation and actual evapotranspiration. If precipitation is greater than actual evapotranspiration, the extra water first fills up the available void space between the solid grains in the soil, then moves down the soil profile to recharge groundwater. If the incoming volume of water continues to be more than actual evapotranspiration, as often happens during the rainy season, the extra water flows out of the soil profile or drainage basin via streams. Streams receive the excess water either as groundwater runoff or as overland flow running on the surface. Computing water balance is an easy way of determining the behaviour of soils and rivers. Instructions for computing water budgets are available in Thornthwaite and Mather (1957) and Dunne and Leopold (1978).

Tropical hydrology

46

(a)

(b)

(c)



Fig. 3.7

Water balance. (a) Moisture-surplus equatorial location, Januarete, Brazil (0.5°N, 69°W); (b) Moisture-deficit arid location, Hyderabad, Pakistan (25.5°N, 68.5°E); (c) Seasonal climate, Kericho, Kenya (0.2°S, 35°E). Data for (a) and (b) from Strahler, 1975, from C. W. Thornthwaite Associates, 1962–65. (c) from Dunne. and Leopold, 1978. With permission from T. Dunne Three basic patterns of water balance can be recognised in the tropics. Near the equator where rainfall is high and a dry season is not conspicuous, precipitation is higher than actual evapotranspiration (Fig. 3.7a). Here the actual and potential evapotranspirations are the same because moisture is always available from high rainfall to meet the demand for evapotranspiration. As a result, streams are seldom dry, the groundwater table is near the surface except in special cases and the soil is moist. All this is reflected in the ambient dense natural vegetation. Arid areas near the subtropical high-pressure systems show the other extreme. In these areas precipitation is low, but under a hot and cloudless sky potential evapotranspiration is high. Actual evapotranspiration is less than the potential evapotranspiration due to the

47

Climate and geomorphology in the tropics

limited rainfall. There is not enough precipitation to meet the evaporation demand, and the water balance is negative throughout the year. The soil is therefore dry, groundwater is deep and streams are usually low or dry (Fig. 3.7b). The annual pattern of water balance is more complex in places with strongly seasonal rainfall. Dunne and Leopold (1978) provided the example of Kericho, Kenya (Fig. 3.7c), which remains relatively dry during December to February but receives high rainfall in April and May. Kericho is located at low latitude, but its climate is seasonal due to the location in the interior and on the western slope of a mountain range. As Figure 3.7c illustrates, the end of the dry season has a moisture deficit as water is lost by evapotranspiration and rainfall is absent or low. With a rise in precipitation, soil moisture is recharged from March, becoming saturated by April. Beyond April, the excess water flows to the river either as groundwater (a process known as baseflow) or overland (also called surface runoff). Rivers are high so long as surplus moisture is present, i.e. precipitation is higher than potential evapotranspiration. Potential and actual evapotranspirations are the same at this phase. As precipitation drops and temperature rises, the volume of precipitation falls below potential evapotranspiration, the difference being met in the early stage of moisture shortage by the evaporation of soil moisture. Potential and actual evapotranspiration continue to be equal. Some time into the dry season, water is no longer available from the soil, potential evapotranspiration is no longer reached, and the period of moisture deficit completes the annual cycle. Rivers in the seasonal tropics thus fluctuate considerably between wet and dry seasons. During the wet season the rivers are high. In many parts of the tropics, large storms (such as tropical cyclones) arrive at this time to cause flooding. Rivers of the monsoon tropics are not only seasonal in flow but they are also floodprone. Tropical slopes also tend to fail seasonally, commonly when they carry too much moisture.

3.8╇ Climate and geomorphology in the tropics Broadly speaking, geomorphological processes in the tropics are influenced by three kinds of climatic variation. Near the equator, it is nearly always warm and humid with a positive water balance and, unless anthropogenically altered, the land is under well-developed vegetation. As discussed in Chapter 5, this leads to advanced weathering and development of deep soils in most places (Dudal, 2005; FAO, 1998). High-magnitude storms are not common and most of the rain is from ITCZ convergence and thundershowers. Away from the equator, the water balance carries a seasonal component with the associated variation in streamflow. The rivers are high during the wet season when most of their work is done. Large tropical storms also are characteristic of seasonal tropics. Weathering will be advanced but not as much as the equatorial tropics in most places. The arid tropics tend to have a negative water balance and weathering and soil formation are restricted. Streams tend to carry water only for a very short period after a rare rainstorm. Such climatic variations give rise to different types of geomorphic processes and associated landforms. Episodic changes such as tropical cyclones or floods and droughts associated with ENSO are superimposed on the general pattern in certain parts of the tropics,

48

Tropical hydrology

periodically influencing the landforms in such regions from time to time. Even if we consider only the climate, we should expect a good deal of variation in the tropical landscape. The reality is even more complex, as the climatic variations operate in conjunction with different types of tectonism, lithology and relief as introduced in Chapter 2. The tropics are a mosaic of different landforms and a range of geomorphic processes, depending on the regional climatic characteristics that operate on this wide range of landforms. The surface of the humid tropics is protected by vegetation, but the vegetation cover varies, depending on local climate and relief. In the next chapter we review tropical vegetation.

Questions 1. You are visiting an island with a high volcanic mountain in the tropical Pacific. What would be the pattern of distribution of annual rainfall on this island? 2. Describe the nature of rainfall during the wet monsoon season. Does it rain uniformly throughout the wet season? 3. You have two rivers flowing in opposite directions. In one case, moisture-laden winds move upstream and in the other the winds travel downstream. If a large rainstorm is embedded in the wind system, what would be its effect on these rivers? 4. Tropical cyclones have been described as geomorphic events in parts of the humid tropics. Why? In which areas are tropical cyclones expected each summer? 5. Explain how climatic variations related to ENSO events influence geomorphic processes in the tropics. Can you find good evidence of such effects in literature or on the web? 6. Describe a typical tropical rainstorm. 7. Figure 3.7c also shows runoff. Explain how this correlates with seasonal variations in temperature and precipitation. 8. Construct a water budget (including surface runoff) for a station for which you have access to the necessary data. See Thornthwaite and Mather (1957) and Dunne and Leopold (1978) for the methodology required for computing water budgets. 9. Explain the quote from J. F. Griffiths at the beginning of the chapter.

4

Erosion and land cover in the tropics

We passed through extensive forests, along paths often up to our knees in mud, and were much annoyed by the leeches for which this district is famous. Alfred Russel Wallace

4.1╇ Erosion from tropical rainfall Tropical rainfall can be strikingly erosive (Fig. 4.1) because it falls with high intensity and often in large amounts. Erosion results from a combination of erosivity and erodibility. Erosivity is the potential ability of a geomorphic agent to erode. Erosivity of a particular rainfall event depends on the physical characteristics of the rain, such as intensity, duration, amount and drop size. Erodibility is the measure of the level of vulnerability of the surface material to erosion. It depends on the properties of the surface material, such as the texture of the surface material, the slope of the land and the protection offered by the local vegetation cover. Intense rainfall, as from convectional thundershowers, erodes, especially when falling on loosely structured soil. The kinetic energy or momentum of a rainfall event determines its erosivity. The intensity of rain is measured as the rate of rainfall, the amount of rain arriving over a fixed period. It is generally measured in mm hr−1, but other time periods are also used, such as millimetres per 30 minutes. Hudson (1981) mentioned that rainfall intensity seldom exceeds 75 mm hr−1 in temperate countries, and that only in summer thunderstorms. In comparison, intensities reaching double that amount are often experienced in many parts of the tropics. Hudson recorded a maximum of 340 mm in an hour in Africa. Low-intensity rain that lasts for long periods usually falls as small raindrops, whereas high-intensity rain carries a range of sizes including big raindrops. The kinetic energy of the rain, on which its erosivity depends, is determined by rainfall intensity and drop size. Hence, a tropical thundershower on unprotected soil can be extremely erosive. When raindrops with high kinetic energy impact the surface, they break it up and splash particles in the air, a process known as splash erosion. This is the beginning of erosion. Most of the splashed-up material comes to rest in the downslope direction. With time and more rain, water starts to run downslope over the surface, removing the detached particles. Runoff begins as a very thin sheet of water moving downhill, concentrated selectively in rills after travelling for some distance. More erosion and removal of surface material happens progressively as rills get bigger and become gullies (Fig.€4.2). Based on empirical evidence, Hudson (1981) identified the intensity of 25 mm hr−1 as 49

50



Fig. 4.1

Erosion and land cover in the tropics

Gully erosion at Baringo, Kenya. From Gupta and Asher, 1998. By permission of Wiley. Photograph: A. Gupta

a threshold that separates erosive and non-erosive rain, with an excellent correlation between rain exceeding this threshold and soil loss. Erosion of the surface material therefore may depend on the proportion of the rain in a rainfall event that falls at intensities greater than 25 mm hr−1 or a comparable value. Research in Nigeria indicated another index, AIm, where A is the amount of rain and Im the maximum intensity over a 7.5 minute period (Lal, 1976). A large volume of intense rainfall, as from convective thundershowers, is highly erosive. Hudson (1981) noted that most of the annual erosion happens in large storms with high-intensity rain such as convective thunderstorms, and not from gentle low-intensity rain. He attributed this as a probable cause for intense soil erosion in tropical and subtropical regions. Drops between 1.2 and 2.5 mm in size and with a terminal velocity of 4.6–7.4 ms−1 have been measured in monsoon rainstorms in India with hourly rainfall intensities of 8.5–55 mm (Pisahorty, 1991). Tropical rainfall often falls with even higher intensity. Tropical rainfall thus has the potential to erode surficial material much more effectively than temperate precipitation, where gentle rain may commonly fall over a long period. The amount of annual rainfall and the extent of natural vegetation cover are directly related. As a result, in spite of the prevalent intense tropical convective thunderstorms, well-vegetated areas such as the tropical rain forest are usually lightly eroded. Areas most vulnerable to erosion are those where intense rainfall falls on land with only a partial coverage of vegetation. In a study that has become classic, Langbein and Schumm (1958) plotted annual sediment yield against annual precipitation adjusted for temperature, i.e. corrected for evapotranspiration. Their data were from small river basins in the United States. The graph showed a single peak, the region with maximum erosion being under short grass with about 300 mm of adjusted precipitation in temperate United States. This peak in annual sediment yield was moved to about 450–500 mm by Dendy and Bolton

51



Fig. 4.2

Erosion from tropical rainfall

Rainfall giving rise to runoff and erosion on a slope

(1976), who analysed data from 800 drainage basins in the United States. The erosion peak was still in grassland. In comparison, a protective cover of vegetation reduces sediment yield in temperate continental areas with more rainfall. In drier areas, lack of running water on the surface due to low rainfall limits erosion and sediment movement. Wilson (1973) found double peaks in the global pattern, one for the semi-arid grassland with about 380 mm of rain and the other for seasonal climates including the Mediterranean and seasonal humid tropics. The data for the tropical wet–dry climate indicated sediment yield peaking at a mean annual precipitation of 1780–1905 mm. With higher rainfall, the natural ground surface is well protected by tropical rain forest and sediment yield drops (Douglas, 1967). Fournier (1960) located the highest rate of erosion also in the seasonal tropics and a double peak in the global pattern. Sediment yield, however, does not depend solely on annual rainfall. It depends also on a multiplicity of other factors that determine erodibility, such as ambient geology, local slopes, properties of rainfall and vegetation cover. Natural vegetation has been destroyed across large parts of the tropics, especially since the mid-twentieth century, for various reasons:€ demand for tropical timber; expansion of agriculture; migrant settlements; and urbanisation. Heavy and intense rain therefore falls, in many places, on bare ground or a cultivated cover. Where the forest cover has been destroyed, the rate of sediment loss increases tremendously as the land surface becomes exposed to high and intense rainfall. For example, an annual sediment yield of 30 t km−2 from the Javanese rain forest rose to 1590 after logging (Anderson and Spencer, 1991). It is therefore necessary to examine both the distribution of natural vegetation in the tropics and its destruction in order to understand the prevalent pattern of erosion and sediment production.

52



Fig. 4.3

Erosion and land cover in the tropics

Map of tropical vegetation, pre-destruction stage

4.2╇ Distribution of natural vegetation in the tropics Temperature and precipitation jointly determine the amount of moisture available to plants. Rain forests occur in the tropics in areas with high rainfall and without a pronounced dry season. Warm areas need higher precipitation to maintain a forest, as a large amount of moisture is lost to the atmosphere by evaporation. The vegetation types in the tropics vary on a gradient determined by moisture availability, seasonality of climate and elevation. With progressively less rainfall we find dry-season deciduous forest (also known as tropical dry forest), grasslands and thorn scrubs, and deserts (Fig. 4.3). Altitudinal zoning of vegetation occurs on high mountains, where with elevation and changing rainfall and temperature, the rain forest of lowlands is altitudinally replaced by lower montane forest, upper montane forest, subalpine forest and, on mountains that rise beyond 3500 m, a treeless community towards the top. On wet mountains such as in New Guinea the treeline may rise to nearly 4000 m, provided the forest has not been disturbed. Ice covers the surface on very high mountains such as the Andes and Himalaya and tops of high volcanic peaks such as Mount Kenya or Kilimanjaro. The vegetation is denser on the wetter sides of the mountains and the different altitudinal zones of forests and tree vegetation reach higher elevations. Specialised forest types occur in specific environments. The major examples are mangroves on tidal coasts, freshwater swamp forests on floodplains and deltas, heath forests on relatively infertile lands, and peat swamp forests on deep ombrotrophic peat rising in a dome above the water table, as on the island of Borneo.

4.3╇ Tropical rain forests Tropical rain forests occur in three large blocks within the tropics. The most extensive area is found in South and Central America covering 4 million km2 in area. Most of this block is located in the basins of the Amazon and Orinoco rivers. A much smaller and discontinuous stretch lies on the Pacific side of the Andes in Ecuador and Colombia, extending northwards through Central America to the southern border of Mexico. A third patch, now almost entirely deforested, used to occur as a narrow strip on the Atlantic coast of Brazil.

53



Fig. 4.4

Tropical rain forests

Shorea curtisii, a tall tree of the tropical rain forest in Southeast Asia. Photograph: A. Gupa The second largest area of the tropical rain forest is an Australasian block, extending discontinuously from the Queensland coast of Australia and Papua New Guinea and other Pacific islands through the islands and peninsulas of Southeast Asia to the wet slopes of the Western Ghats of south India overlooking the Arabian Sea. The smallest extent of the rain forest is in Africa, centred on the basin of the Congo, with small outliers in East Africa. Small patches are found on the wet east coast of Madagascar where rain forest used to be continuous until recently. A strip detached from the Congo forest occurs along the coast in West Africa. The nature of the forest varies depending on elevation, rainfall, soil characteristics, etc., but the well-developed variety found in the lowlands is mostly evergreen and high. The general canopy is at 30–40 m above ground with scattered tall trees up to 50 m high, known as the emergents, rising above this level. Smaller trees occur underneath the canopy. The trees tend to have trunks with a wide girth and branching only near the top (Fig. 4.4). Trees of the montane rain forest are smaller but still grow 15–33 m tall on the lower slopes. Ground vegetation is usually sparse. A litter layer covers the forest floor. It varies in thickness from 1–2 cm of recently fallen litter to about 15 cm of material on infertile soil (Kellman and Tackaberry, 1997). The most extensive tropical forest type in the past was probably semi-evergreen rain forest with a largely deciduous canopy and an evergreen understorey, which still occurs in Africa. With repeated burning, this type tends to change towards deciduous forest or savanna (R. Corlett, personal communication).

54



Fig. 4.5

Erosion and land cover in the tropics

Downward passage of rainwater through a tropical rain forest

4.3.1╇ Passage of rainwater through tropical rain forest The structure of the forest affects the falling of raindrops. Only a fraction impacts directly on the ground. In a tropical rain forest, or indeed in any high forest with well-developed canopy, rainwater, after reaching canopy level, moves as (Fig. 4.5) • throughfall (raindrops pass through gaps in the canopy to reach the ground) • drips from the canopy (rain intercepted at canopy level may drop later from leaves and twigs) • stemflow (water accumulated in leaves and twigs at the canopy level, slowly running down straight, long and wide trunks to reach the ground). Both throughfall and drips from canopy may cause splash erosion, throughfall because of the high intensity of the tropical rain and canopy drips as the canopy in a rain forest is high enough for the falling drops to acquire some velocity. In general, however, the potential for erosion by rainfall in a tropical rain forest is low (Brandt, 1988). Vegetative litter on the ground may also reduce rain splash. After a period of intense rain, water on the ground may briefly flow on leaf-littered surfaces or flow through the subsurface via soil pores and pipes at shallow depth as described by Sidle et al. (2000). A spell of intense convectional showers is often followed by water emerging and cascading down from exposed ends of soil pipes midway up a vertical bank

Tropical deciduous forests, grasslands and deserts

55

or a road cut. A number of case studies have been carried out on rain forest slopes, for example in eastern Puerto Rico, Southeast Asia and northern Australia. Bruijnzeel (1989) generalised that 80–85 per cent of the rain reaching the ground surface arrives via throughfall and 1–2 per cent as stemflow. The rest is retained by the canopy and evaporated back to the atmosphere. Comparable figures came from a study in the Danum Valley of Sabah, East Malaysia where the annual rainfall is around 3600 mm: interception throughfall stemflow

17.4% 80.7% 1.9%.

According to this study, (Sinun et al., 1992), about 2.0–2.5 per cent of the rain on reaching the ground travels as overland flow. The rest infiltrates and probably flows mostly through soil pipes. Surface runoff and erosion therefore is limited in a tropical rain forest and widespread overland flow is common only for short periods following heavy and intense rain. The water, however, may react chemically with organisms and plant litter to turn acidic and act on underlying soils and rocks as described in Chapter 5.

4.4╇ Tropical deciduous forests, grasslands and deserts Trees of the rain forest shed their leaves individually and never in a synchronised fashion species-wise, but the tropical deciduous forest displays a seasonal pattern, with trees becoming bare in early summer. The leaves grow back with the arrival of the wet monsoon, but the ground surface is prone to erosion in the first few weeks of the wet monsoon before new leaves grow to provide some protection from the momentum of raindrops. Compared to the rain forest, the canopy is less dense and at a lower height. Shrubby undergrowth partially covers the ground surface and bamboos are common although not ubiquitous. Erosion by sheetflow and gullies happens on the slopes under tropical deciduous forest. This type of forest is found skirting the drier edges of the rain forest in East, Southeast and South Asia, with the best example in South Asia. It occurs in patches in Africa, the Pacific slopes of the Andes in tropical South America, Brazil, Central America, Mexico and the Caribbean islands. Tropical grasslands or savannas are plant communities dominated by grasses or grasslike plants (graminoids), although not entirely without trees. Trees occur where the moisture supply is better, e.g. in areas of locally high rainfall or along stream banks, a phenomenon known as gallery forest (Fig. 4.6). The canopy is never closed. Savannas therefore run the whole gamut from open forests with a graminoid understorey to grasslands without trees (Fig. 4.7). Tropical grasslands and thorny scrubs cover enormous areas in tropical Africa. Kellman and Tackaberry (1997) associated African savanna with subhumid areas receiving between 500 and 1500 mm of annual rainfall as an intermediate region between the tropical rain forest of Central Africa and the arid subtropical deserts. Savannas, however, also occur where the soil is unsuitable for tree growth or in areas affected by anthropogenic alteration, and thus are found outside the subhumid belt. For example, the savanna known as llanos in South America is associated with 800–1500 mm of annual rainfall. These grasslands,

56

Erosion and land cover in the tropics



Fig. 4.6

Grassland with gallery forest, Masai Mara, Kenya. Photograph: A. Gupta



Fig. 4.7

Short grassland, Masai Mara, Kenya. Photograph: A. Gupta however, have been reported to occur also in areas receiving much higher rainfall but with a very long dry season. Tropical grasslands extend over huge areas south of the Amazon Basin rain forest, and along the Andes foreland in the western Orinoco Basin. Scattered areas of various types of grassland occur throughout tropical South America, Central America and the Caribbean islands, not only where the annual rainfall is low (subhumid) but if the geology or landform is also unsuitable for plant growth. Grasslands thus occur on old crystalline rocks of the Brazilian Shield and Guinea Shield and the seasonally flooded Pantanal swampland in southwestern Brazil. In Asia-Pacific, tropical grasslands are extensive in northern Australia, in the interior of Papua New Guinea and the dry interior of Myanmar and the Indian Peninsula.



57

Anthropogenic alteration of the tropical vegetation

Fig. 4.8

Vegetation clearance for oil palm plantation in eastern Sumatra, Southeast Asia. Note lines of young oil palms towards the upper left corner, canals put in for drainage and burning of vegetation in the centre. IKONOS satellite image © Center for Remote Imaging, Sensing and Processing, National University of Singapore (2006), reproduced with permission. See also colour plate section Tropical grasslands cover about 60 per cent of the uncultivated tropical land area (Kellman and Tackaberry, 1997) and, together with thorny scrublands, indicate that a seasonal climate with limited annual rainfall covers most of the tropics. The extensive occurrence of grasslands, thorny scrubs and subtropical deserts in the tropics (Fig. 4.3) is a reminder that geomorphic processes vary across these latitudes, and a perhumid climate or vegetation should not be expected everywhere.

4.5╇ Anthropogenic alteration of the tropical vegetation The destruction of the natural vegetation started with the beginning of agriculture at the beginning of the Holocene. Within a few thousand years, forests were being depleted not only for agricultural expansion but also for construction material and urbanisation. Fertile river valleys and coastal plains such as the Ganga Plains or volcanic slopes as in Java were cleared thousands of years ago. From the mid-twentieth century the demands for timber and for rubber and oil palm accelerated the destruction of the tropical rain forest, especially in South and Southeast Asia (Fig. 4.8). Rain forests were also destroyed in large-scale projects for resettling people from densely populated areas, as in Indonesia and Malaysia, and by spreading of pasture in Central America. Tropical seasonal forests were depleted

58

Erosion and land cover in the tropics

for agricultural land and firewood. Details of deforestation have been reported by many authors (Douglas et al., 1993; Gupta and Asher, 1998; Corlett, 2009). Vast areas of lowland forest have disappeared since the 1970s in Malaysia and western Indonesia, in fact from everywhere. Such widespread deforestation implies that: 1. the present forests are a fraction of their past extent 2. with the disappearance of the forest canopy and other natural protective devices, soil erosion has increased and the present rates of erosion on slopes and deposition of sediment in rivers and coastal waters are not necessarily natural rates 3. a significant part of this destruction followed by changes in the physical environment has happened very recently, in the last few decades, and change is continuing. The anthropogenic component of these changes affecting the physical environment and geomorphic processes in the tropics is examined in detail in Chapter 17. It is, however, important to emphasise even at this stage that current landforms and operating processes in the tropics are not entirely natural and include a strong anthropogenic component. This, of course, is a global trend, but the tropical countries have recently experienced a particularly rapid rate of population growth and developmental changes which are still continuing. As a result, many tropical slopes are now exposed to wind and rain damage following the removal of the protection offered by a vegetation cover. These slopes fail repeatedly with a greater volume of sediment loading the streams and ultimately reaching the coastal waters. The present landforms and geomorphic processes have been significantly altered from the natural appearance and rates. Tropical geomorphology therefore requires an understanding of both the distribution of natural vegetation that still survives and its progressive anthropogenic alteration. This, given the nature of tropical rainfall, probably had a greater impact in the humid and subhumid tropics than in other parts of the world. A natural analogue of this anthropogenic destruction is the nature of post-hurricane (or typhoon) slopes when the destruction of the vegetation in high winds is followed by torrential rain. Slopes fail, rivers flood and coastal waters are polluted with sediment.

Questions 1. Why is tropical rainfall highly erosive? 2. What amount of annual rainfall has been associated with high sediment yields? Can you explain such figures? 3. Why is erosion limited in an undisturbed tropical rain forest? 4. Is it correct to state that a seasonal climate, with limited annual rainfall and grassland with scattered trees, best represents the tropics? 5. Distribution of the surviving natural vegetation and its progressive anthropogenic alteration are both important factors in tropical geomorphology. Do you agree?

Part II

Process geomorphology in the tropics

5

Weathering in the tropics

Hover through the fog and filthy air William Shakespeare

5.1╇ Introduction The process of weathering alters rocks at or near the surface of the Earth, transforming hard rocks into soft material that can be removed easily (Fig. 5.1). Rocks, which are formed at various depths inside the Earth’s crust under conditions of high pressure and temperature, undergo a series of physical and chemical changes when they are exposed to the atmosphere at or near the surface of the Earth where much lower temperate and pressure prevail. The term ‘weathering’ encompasses the full range of such changes. A formal definition can be constructed as Weathering is a process, at or near the interface between the crust and atmosphere, which alters the physical and chemical nature of rocks in situ.

The process of weathering comprises a set of sub-processes that can be grouped broadly into two classes:€ physical and chemical weathering. Physical weathering includes the sub-processes that lead to the breaking up of rocks into smaller units without significant chemical alterations. In chemical weathering, original minerals in the hard rock react with ambient atmospheric conditions and with any water flowing across the rock surface to change into new and softer minerals. Usually both types of weathering occur together, although one may predominate. In the arid tropics, where atmospheric moisture is very low, rocks tend to break up into smaller components without much chemical change. In the humid tropics, chemical weathering is accelerated in the presence of high temperature and moisture. The pattern and products of weathering are therefore different in various parts of the tropics (Fig. 5.2). The nature of the final weathered product depends on • mineral composition of the original rock • prevailing climate • ambient vegetation • time. A rock, when well weathered, becomes softer. It is the first step to erosion. The upper part of the soft weathered rock is further transformed by soil-forming processes into soils. The subsurface material therefore shows a vertical zonation:€ soils, weathered rock and 61

Weathering in the tropics

62



Fig. 5.1 Schematic diagram indicating the sequence of changes over time from hard rock to soft weathered rock and soil

(a)

(b)



Fig. 5.2

(a) Physical weathering in southern Israel:€note clasts of broken rock on the surface; (b) Chemical weathering in Mauritius:€red fine-grained soft material developed over basalt; growth of sugarcane indicates a well-developed soil. Photographs: A. Gupta



Fig. 5.3 The vertical zonation of soil, weathered rock and unweathered bedrock. Duricrusts (hard layers of precipitated

material) may occur in soil and weathered rock. Regolith describes the total material above unweathered rock. Saprolite usually refers to weathered rock that still carries the rock structure inherited from the unweathered rock

63

Sub-processes of weathering:€a brief review

unweathered rock (Fig. 5.3). The structure of the original rock may still be recognised in parts of the weathered rock layer below the soil. The total thickness of soil and weathered rock increases with time and rate of weathering. This rate is extremely slow in the arid tropics, giving rise to a thin layer of physically broken material on hard rock. In the humid tropics, where accelerated chemical weathering changes minerals at a rapid pace, a thick layer of soil and weathered material forms and is found over hard rocks. Chemical weathering is therefore an important process in the humid tropics where slopes are underlain by thick layers of soft erodible material, except where such material cannot accumulate as it is quickly removed by erosion. Slopes affected by tectonics and storms therefore do not generally carry a thick weathered layer. The eroded material from the soft soil and weathered rock ends up in rivers. Sediment carried by rivers is thus weathering-dependent.

5.2╇ Sub-processes of weathering:€a brief review Weathering combines several sub-processes, which may act together. These, however, are easier to comprehend when discussed individually. This section is a brief review of the sub-processes.

5.2.1╇ Physical weathering Physical weathering leads to mechanical disintegration of a rock into smaller particles; the process therefore is also known as mechanical weathering. Apart from breaking down a large piece of rock into small individual particles, each of which is easy to remove individually, such disintegration also increases the specific surface of the rock. The specific surface of a solid is defined as the sum of its surface areas. A cube with the volume of 1€cm3 has six sides, each with an area of 1 cm2. The specific surface of the cube therefore is 6 cm2cm−3. On breaking this cube into 10 units of 0.1 cm3 each, the total specific surface becomes 60 cm2cm−3. When broken to units of 0.001 cm−3, the size of a colloid, the total specific surface becomes 60 000 cm2cm−3 (Ruhe, 1975). As chemical reactions take place when water flows across the surface of minerals, an increase in surface area accelerates chemical weathering. Furthermore, chemical bonds in the structure of minerals break down when particles disintegrate to a very small size, leaving unbalanced charges at various points on the surface of the mineral, which leads to chemical reactions. Physical weathering thus accelerates chemical weathering. Physical weathering happens when external stresses set up by ambient conditions are stronger than the internal resistant forces of the rock, and the rock disintegrates mechanically. These stresses are set up by pressure release, crystal growth, alternate wetting and drying, and, in case of weak material, by biological agents. The pressure on the surface of a rock such as granite is reduced when its overburden (the material lying on top of the granite which was originally formed inside the Earth’s crust) is removed by erosion. This may lead to a slight expansion of the rock and the development of shear planes at right angles to the direction of stress release. The upper part of the

64



Weathering in the tropics

Fig. 5.4 Exfoliation in andesite, Lombok, Indonesia. Photograph: A. Gupta

rock therefore develops joints that are curved and run parallel to the surface of the rock. Concentric slabs then come off the rock as various processes of weathering act on it. The process of curved slab removal is called exfoliation (Fig. 5.4). Joints may also develop in a rock as it cools from a hot magma or is subjected to tectonic forces. Such joints usually occur as parallel or right-angled sets. Where joints intersect, rocks may break up into blocks as the first step to disintegration. Joints also allow water to enter a rock and promote chemical action. Growth of ice crystals from water that enters openings in rock and freezes also generates stress. The change of state from water to ice leads to a 9 per cent expansion in volume and with ice filling the cracks the stress on the rock can be high, the maximum recorded being over 2000 kgcm−2. This far exceeds the tensile strengths of rocks which split up into small particles. Crystallisation of minerals such as halite, gypsum and calcite from salt solutions filling openings in rock may lead to rock disintegration in a similar fashion. Alternate wetting and drying leads to physical disintegration of soils and the breakdown of certain rock minerals, such as phyllosilicates, that undergo volumetric changes by absorption or loss of water. Direct heating of rock faces to high temperature, as in bushfires, also leads to rock disintegration. Tree roots entering cracks in a rock may expedite its disintegration, but possibly with simultaneous chemical weathering by accompanying acid solutions as discussed below. Tree roots by themselves, however, are only capable of disturbing relatively softer material such as soil or weathered rock. Ants and termites also turn up huge volumes of soft unconsolidated material that lie over hard rock (Fig. 5.5). It follows that physical weathering in the tropics is significant in (a) relatively dry areas where crystal growths and episodic bushfires occur and (b) in mountains and deserts where the temperature may oscillate across the freezing point of water, falling below freezing at night or in winter. But physical weathering does not happen in isolation even in deserts. Weathering in the central Namib Desert is associated with salt weathering, and lichens and

65

Sub-processes of weathering:€a brief review



Fig. 5.5 Termite mounds, Kakadu, Northern Australia. Photographs: A. Gupta



Box 5.1

Phyllosilicates and clay minerals Phyllosilicates are a group of minerals which are physically made of alternate layers of silicon tetrahedra (SiO4) stacked vertically with either layers of minerals gibbsite (Al(OH)3) or brucite (Mg(OH)2). As a result, these minerals tend to cleave easily in thin sheets. In some minerals, water can be stored in between the sheets or stored water can be lost from there. As a result, these minerals can swell or shrink, a property which makes them unstable and prone to disintegration. Minerals such as kaolinite, talc, montmorillonite, smectite and various types of mica such as muscovite or biotite, fall in this group. A number of fine-grained minerals that belong to this group are known as clay minerals. Kaolinite, for example, is a clay mineral. Clay minerals are earthy masses, which are very soft and may hold some water. They are products of weathering and are easily eroded. The word 'clay' denotes grains smaller than 2 μm in diameter. Clay minerals are also fine-grained but they have extra characteristics of being a group of hydrous silicates with a layered structure. Although a number of clay minerals have the texture of clay, the two terms do not mean exactly the same. The clay fraction, being a size term, may also include quartz, feldspar, oxides of iron and carbonates. In comparison, certain clay minerals may have a texture coarser than clay.

Weathering in the tropics

66

the moisture in fogs dissolving carbonates in the ground. Minor solution features on granite and marble are quite common (Viles, 2005). Elsewhere in the tropics chemical weathering dominates, although disintegration by physical weathering, if it happens, would be a subprocess in support.

5.2.2╇ Chemical weathering with special reference to the humid tropics The process and rate of chemical weathering depend on the combined effect of a number of factors: 1. temperature 2. rainfall 3. vegetation 4. physical weathering 5. lithology 6. relief. The first three factors are common in the humid tropics; others may be of local importance. Chemical weathering is therefore extremely effective in the humid tropics. Chemical weathering happens when a weak acidic solution passes across the rock face. This acidic solution is derived in several ways. Rainwater may contain a limited quantity of carbonic (H2CO3) and sulphuric (H2SO4) acids. Carbonic acid is also produced by microbes in the soil by oxidising organic matter to CO2. Sulphuric acid is produced by bacteria from sulphide minerals, and a low concentration of this acid may also occur in rainwater. Water in the subsurface (also called soil water) may carry organic acids (humic and fulvic acids) in the tropics, released by fast partial decomposition of vegetative matter on the forest floor. Schlesinger (1991) provided a list of mean residence time for litter in different kinds of forest. The mean residence time for forest litter is 353 years in boreal forest, 4 years in deciduous forest, 3.8 years in Mediterranean scrubland, but only 3 months in the tropical rain forest. The prevailing high temperature in the tropics increases the decomposition rate of soil organic matter (Berner and Berner, 1996). The importance of biomass storage has also been demonstrated by Taylor and Velbel (1991), working in the southern Appalachians, which are outside the tropics but their evidence can be treated as an indicator. According to their calculations, stored biomass increases the weathering rate of minerals by a factor of four. Biomass may be defined as the dead organic matter resting on the surface of the ground and in the soil. The chemical constituents between the rainwater and the soil water change because of the selective storing of elements (such as phosphorus and nitrogen) by vegetation. As only pure water is lost by transpiration, vegetation increases the concentration of dissolved ions such as Cl− in soil water (Berner and Berner, 1996). Acids thus produced and carried in soil solution expedite rock weathering. The hydrogen ions in acids replace the cations in minerals as the acid solution flows across the rock face, a process known as dissolution, which happens at a fast rate in the humid tropics. In addition, organic acids colour soil solutions brown or red (a colour common in the tropics), disintegrate minerals, and remove iron and aluminium by chelation (formation of a ring structure round the

Sub-processes of weathering:€a brief review

67

Table 5.1╇Common primary minerals and weathering Mineral

General composition

Type of reaction

Olivine

(Mg, Fe)2SiO4

Pyroxenes

Ca(Mg, Fe)Si2O6/(Mg, Fe)SiO3

Amphiboles

Ca2(Mg, Fe)5Si8O22(OH)2, Na and Al may be present Solid solution between albite (NaAlSi3O8) and anorthite (CaAl2Si2O8) KAlSi3O8 K(Mg,Fe)3(AlSi3O10)(OH)2

Formation of iron oxides, acid dissolution © Formation of iron oxides, acid dissolution © Formation of iron oxides, acid dissolution © Acid dissolution (i)

Plagioclase feldspar

K-feldspar Biotite Muscovite Volcanic glass Quartz Calcite Dolomite Gypsum Anhydrite Halite Pyrite

KAl2(AlSi3O10)(OH)2 SiO2 CaCO3 Ca,Mg(CO3)2 CaSO4.2H2O CaSO4 NaCl FeS2

Acid dissolution (i) Formation of iron oxides, acid dissolution (i) Acid dissolution (i) Water and acid dissolution (i) Difficult to dissolve Acid dissolution © Acid dissolution © Dissolved by water © Dissolved by water © Dissolved by water © Formation of oxides of iron and sulphur

Note:€© congruent or simple dissolution; (i) incongruent or dissolution with reprecipitation Source:€Berner and Berner, 1996

metallic ion). These acts are followed by a downward transfer to the subsurface, where iron and aluminium are reprecipitated, usually at a depth of tens of centimetres. Dissolution is one of a series of chemical reactions, probably the most important, that together constitute chemical weathering. In dissolution, the primary mineral (Table€5.1) is either dissolved completely by acidic solutions passing the rock face (congruent dissolution) or a fraction of the original mineral is reprecipitated to form one or more secondary minerals (incongruent dissolution). Oxygen dissolved in water also reacts with reduced iron and sulphur to form new minerals. Minerals such as halite, calcite and gypsum easily dissolve in subsurface or rainwater, both seasonally abundant in the humid tropics. Acidic solutions pass downwards and laterally through the soil and already weathered rock to reach primary minerals in unweathered material. Berner and Berner (1996) described a selective beginning of dissolution on the face of the mineral crystals, where solutions first attack rows of atoms slightly out of place in the crystal lattice. Reactions with the flowing solution start in these locations, leading sequentially to selective etching, formation of etch pits on the mineral surface, and coalescence of such pits, ending finally in the disappearance of the original mineral.

Weathering in the tropics

68

Silicate minerals commonly dissolve incongruously. Congruent dissolution is limited to only olivine, amphiboles and pyroxenes with low iron content, as the following chemical reactions from Berner and Berner (1996) illustrate. Mg2SiO4 (forsterite, an olivine with very little or no iron) + 4H2CO3 (carbonic acid) → 2Mg2+ + 4HCO– + H4SiO4 (dissolved silica)

CaMgSi2O6 (diopside, a pyroxene with very little or no iron) + 2H2O + 4H2CO3 → Ca2+ + Mg2+ + 4HCO– + 2H4SiO4

The iron-free olivine and pyroxene, therefore, can be completely dissolved and removed in solution. This is a simple case, but most silicate minerals weather by incongruent dissolution, which is complex. The following equation summarises the decomposition of a feldspar (Berner and Berner, 1996). 2NaAlSi3O8 (albite, a plagioclase feldspar) + 4H2C2O4 (oxalic acid) + 2O2 + 7H2O → Al2Si2O5(OH)4 (kaolinite) + 2Na+ + 2HCO3– + 4 H4SiO4 + 6CO2 After incongruent dissolution of albite, kaolinite will be left at the place of weathering, CO2 will dissipate as a gas and other products will be removed in solution. Feldspars occur primarily in igneous rocks along with quartz (SiO2), and quartz remains in situ with kaolinite from feldspar as an undissolved residue. Granite is commonly weathered in this fashion, the hard rock being replaced by sand from quartz and clay from weathered feldspar and other minor constituent minerals. Loughnan (1969) showed the importance of acidity in water for dissolving minerals and rock component. Silica is not dissolved unless the water is highly alkaline. Oxides of iron or aluminium are dissolved only in highly acidic or alkaline water. Microorganisms, so commonly present in the soils of the humid tropics, produce carbonic and sulphuric acids that expedite dissolution. Oxalic acid is exuded by fungi in the forest litter and the upper parts of the soil. As a result, iron and aluminium in the humid tropics are efficiently removed from the minerals undergoing weathering and deposited elsewhere. In the forests of the humid tropics, therefore, organic acids colour soil solutions (subsurface water circulating through soil) red or brown, decompose minerals, and chelate and solubilise iron and aluminium and transfer them to lower soil horizons. Other chemical reactions may happen separately or simultaneously with dissolution. Absorption of water without chemical change in a mineral is known as hydration. Anhydrite (CaSO4) changes into gypsum (CaSO4.2H2O) by absorbing water and swelling up. Hematite (Fe2O3) changes into limonite or goethite (HFeO2). Iron is transferred from ferrous state to ferric by oxidation with the atmosphere. For example, an iron-bearing mineral may oxidise to iron oxides. Such oxides may then change to goethite (FeO.OH). On losing water, goethite changes back to hematite. When metallic cations on the surface of a mineral are replaced by free hydrogen in solution, the process is called hydrolysis. For example, 4KAlSi3O8 (potassium feldspar) + 4H+ + 2H2O → 4K+ + 2Al2Si2O5(OH)4 + 8SiO2

Products of weathering

69

Table 5.2╇Common secondary minerals formed by weathering Mineral

General composition

Hematite Goethite Gibbsite Kaolinite Smectite Vermiculite Calcite Gypsum

Fe2O3 FeO. OH Al(OH)3 Al2Si2O5(OH)4 (Ca,Na)Al3MgSi8O20 (OH)4.nH2O CaCO3 CaSO4.2H2O

Potassium feldspar thus changes into clay minerals and silica. Carbonate minerals weather by simple congruent dissolution, when attacked by carbonic acid. CaCO3 (calcite) + H2CO3 → Ca2+ + 2HCO3− CaMg(CO3)2 (dolomite) + 2H2CO3 → Ca2+ + Mg2+ + 4HCO3− Limestones undergo selective dissolution along existing bedding planes, joints and other openings. The openings are enlarged by removal of Ca2+ and HCO3− in solution and, given enough time, caves are produced. The air in caves, especially those connected by an air passage to the atmosphere, has a low partial pressure of CO2. When water flows into such caves along subterraneous passages carrying high levels of CO2, the gas is lost to the cave atmosphere. As a result the water becomes saturated with CaCO3, which is reprecipitated as calcite in the caves. Even in dolomite caves, only calcite is deposited, as dolomite is difficult to precipitate. Calcite is also deposited in soils from solutions due to similar degassing. The more arid the climate, the shallower is the precipitated layer of calcite, an extensive layer of which is known as caliche.

5.3╇ Products of weathering When a primary mineral is attacked by a reaction agent such as an acidic solution, it undergoes chemical changes. Such changes produce one or more secondary minerals (Table 5.2), free cations and anions that travel with the solution, and remnants of the original mineral. The final product is not always the same. It depends on the rate of flow and the pH-value of the attacking solution. In an earlier example, plagioclase feldspar changed into kaolinite, but two other alternatives are common:€smectite ((Ca,Na)Al3MgSi8O20(OH)4.nH2O) and gibbsite (Al(OH)3). The formation of the clay mineral depends on the rate of flow of water. A slow passage of water across the face of the plagioclase increases the time of contact of a unit volume of water with the mineral, and the solution becomes progressively more saturated

70



Fig. 5.6

Weathering in the tropics

Formation of different clay minerals depending on rainfall, following Sherman’s (1952) example from Hawaii. Montmorillonite is a type of smectite with the dissolved material. The proportion of both silica and cations builds up in the solution and resultant chemical change forms smectite instead of kaolinite. In contrast, when water travels fast, concentration remains low, and cations and silica are flushed out of the system. This gives rise to gibbsite being formed. Kaolinite is formed when slow flushing takes place but the water is not stagnant enough to allow a pronounced build-up of cations and silica. Therefore gibbsite is found where high rainfall and good drainage prevail and smectite where water stagnates, i.e. at a break in a slope or a relatively dry area. An excellent and often used illustration comes from Sherman’s work in Hawaii (Sherman, 1952), where the volcanic highlands create a wind and lee effect against the moisture-bearing easterly trade winds and formation of smectite, kaolinite and bauxite (gibbsite) from basalt is directly related to annual rainfall on slopes (Fig. 5.6). A second example is the case of fast-flowing water down the steep sides of andesitic volcanoes in Indonesia that slows down near the break in the slope at the base. Kaolinite is found on the upper slopes and smectite in poorly drained or swampy depressions at the base (Mohr and van Buren, 1954). In South India, weathering in humid conditions gives rise to the formation of kaolinite and gibbsite, but where rainfall is low, smectite prevails (Bronger and Bruhn, 1989). The presence of extra cations such as Na+ in the primary mineral also favours the formation of smectite. Variable residence times of water in different layers of soil may give rise to a vertical zonation of clay minerals with gibbsite towards the top and gibbsite and kaolinite below. Entrapment of water in slightly weathered buried rock produces smectite (Berner and Berner, 1996). The characteristic clay mineral over most of the humid tropics, however, is kaolinite. This discussion is based on plagioclase feldspar as the primary mineral undergoing weathering. Other silicate minerals likewise weather to different clay minerals, the end product depending on their chemical composition and rate of weathering. Generalisations can be made regarding weathering of different classes of rock.

71



Fig. 5.7

Products of weathering

Weathering in volcaniclastic rock, Puerto Rico, producing both boulders and red clay. Photograph: A. Gupta

5.3.1╇ Weathering of common extrusive igneous rocks Common extrusive igneous rocks are basalt and andesite. Basalt is fine-grained and consists of olivine, pyroxene and calcium-rich plagioclase feldspar. Andesites are generally made of plagioclase feldspar, amphibole and pyroxene. A small fraction of olivine may be present. Both these rocks weather rapidly in the humid tropics, and smectite, kaolinite or gibbsite is produced depending on the hydrologic conditions. Unless volcanic clasts or quartz veins are present in the original rock, basalt and andesite would, over time, weather entirely to clay. If clasts or quartz veins are present, coarse gravel is found embedded within the clay (Fig. 5.7). The soils on extrusive rocks are generally clay-rich, dark and contain cations such as Ca, Mg or Fe. This pattern of weathering is seen in the humid tropics in mid-ocean islands, over flood basalts on continents and in subduction-related volcanic mountains. Stable slopes on extrusive rocks and volcanic mountains in the humid tropics therefore tend to have a thick weathered layer near the surface (Fig. 5.2b). Rivers draining such areas carry fine-grained material. If volcanic clasts or vein quartz is present, boulders are found with the clay in stream channels. Over time the total thickness of soil and weathered rock may extend to several tens of metres.

5.3.2╇ Weathering of common intrusive igneous rocks Granite and granodiorite are common intrusive rocks. Granite consists of quartz, potassium feldspar, plagioclase feldspar and small amounts of muscovite, biotite and amphibole. As expected, amphibole, biotite and plagioclase feldspar weather early in a hot and humid environment, followed by potassium feldspar and muscovite. All these minerals change into clay minerals and quartz is left as individual grains of sand. Granite under advanced

72



Weathering in the tropics

Fig. 5.8 Schematic zones in weathered granite after Ruxton and Berry, 1957

chemical weathering therefore transforms to a mass of sand and clay, usually producing more clay by volume than sand. Granodiorite with comparable mineralogical composition changes to similar end products. Like igneous extrusive rocks, granite weathers deep in a hot and humid climate. More than 100 m has been recorded. Ruxton and Berry (1957) proposed a general vertical profile (Fig. 5.8) for weathered granite based on their work in Hong Kong. The profile indicates fresh bedrock at depths with joints and various openings in it and various states of weathered rocks above it. Solutions penetrate along the openings in granite and weather the rock both vertically and laterally, the depth of weathered rock increasing with time. For most of the profile, isolated bits of granite (known as corestones) are separated from each other by sandy clay. The final end product is found at the top of the profile, a mixture of clay with sand. The profiles tend to develop better on gentle slopes and ridge tops and less well on steeper slopes. Weathered profiles up to 85 m have been reported from Hong Kong, although most profiles tend to be less than 20 m deep (Bennett, 1984). Intensity of weathering increases with distance from the centre of a corestone. For example, the centre of a corestone in zone 3 of the Ruxton and Berry model is likely to be fresh, unweathered rock. Its margins, however, may be reddened indicating the beginning of chemical weathering. Beyond the margins of this corestone, feldspars and biotite of the original rock have undergone chemical changes leading to sand and clay minerals replacing weathered rock away from the corestone. Further away clay will be present, reddened by iron oxides. This is zone 2 with only a few corestones. With time the entire corestone will disappear, as in zone 1, replaced by sand and clay minerals. A granitic hill is commonly dome-shaped, reflecting its formation inside the crust where the granitic magma commonly fills a large dome-shaped cavity and also forces a bulge to form. Such hills are often seen standing in isolation in the middle of a plain which appears

Products of weathering

73



Fig. 5.9 Corestones, Southeast Asia

to have formed by erosion. Such hills are called inselbergs (island mountains) or bornhardts (after an early German geologist). With time, the granite is weathered into the standard bimodal type of granite regolith:€finer material and corestones. The fine material is commonly moved by slope processes and running water over time, leaving the slopes of very old granitic hills covered by corestones (Fig. 5.9). Bare granitic domes are also seen protected by a thick weathering layer on the surface (called a rind), cut with deep grooves.

5.3.3╇ Weathering of sedimentary rocks Weathering of sedimentary rocks depends on their composition. Sandstone primarily consists of sand-size grains, mostly quartz with some feldspar and fragments of earlier rocks, held together either by cements such as calcite or silica, or by a clay matrix. Only feldspars and rock fragments weather chemically and so does the cement. Weathering is usually slow and limited, but over time a thin layer of loosened sand grains tends to overlie the unweathered sandstone. The decomposition of the cement and the feldspar loosens the quartz grains. Individual grains drop off, the rock becomes fragmented and sand grains are removed over time. If water is present in the subsurface, then piping and further removal of the rock mass occur. Water penetrates sandstones along joints and other openings, breaks down the rock by dissolving its cement or loosening its matrix and then removes the individual grains (Fig. 5.10).

74

Weathering in the tropics



Fig. 5.10 Subsurface piping in sandstone, Northern Territories, Australia



Fig. 5.11

Weathering in conglomerate, Kuta Tjuta, Australia Water circulating inside sandstone along pipes is more effective than water running over the sandstone surface (Young, et al., 2009). Water running on top may even protect the surface by precipitating crusts of silica or iron oxides. This happens best in alternating wet and dry conditions in a seasonal climate. Young et al. (2009) explain the decomposition of the massive rock of Uluru in Australia by drawing attention to the presence of 80° dips in

75

Weathering and vertical zonation

rock which allows penetration of water throughout its mass. The surface of Uluru displays a number of pipes emerging from inside the rock mass. The surface also shows a series of accelerated spalling (breaking of the rock in thin layers starting from the surface) in arkosic sandstone, which has been attributed to both tensional stresses set up during valley excavations and hydration. The role of a protecting caprock is crucial, as the absence of it allows water to penetrate the sandstone thereby weakening it (Young et al., 2009). In the case of coarser conglomeratic sandstones, as the cement or matrix decomposes the rock breaks down to pebbles, cobbles or boulders as in the case of the conglomeratic hills of Kota Tujah in Australia (Fig. 5.11). Usually the weathering and erosion rates are slow in sandstone, but if the sandstone is part of a craton, the rock has time to gradually disintegrate and be removed. The solubility of silica has been directly related to temperature, which suggests that tropical streams may carry a relatively higher silica load (Meybeck, 1987). This has not always been the case. Very low silica concentration, ≤ 1 mg/l, has been usually recorded for waters draining quartz-rich areas in Venezuela (Chalcraft and Pye, 1984). Weathering and subsequent erosion along joints may give rise to characteristic topography in sandstone. These include cliffs with rounded tops, sandstone towers separated by jointcontrolled corridors, deep shafts, caves, and surface features ranging from small pits, hollows, boxwork joint-driven patterns to larger hollows in rock called alveolis and tafonis. Unlike sandstone, limestone undergoes rapid chemical change in the humid tropics. The calcium carbonate in the rock goes into solution and then is transported out of the source area. The calcium may be reprecipitated as the water circulates because of temperature and pressure change in underground caves. The non-carbonates in the limestone are released, frequently as precipitates, as the rock is dissolved. If iron is present, it colours the weathered material and soil. Soils on limestone in the humid tropics are usually red. Even a small amount of iron in the soil water will accomplish this. In the arid tropics where moisture is limited, physical weathering tends to break up limestone first in blocks and then into individual grains. The solubility of limestone and the resulting topography in the tropics is discussed in Chapter 15.

5.3.4╇ Weathering of metamorphic rocks Our knowledge about weathering of metamorphic rocks in the tropics is limited. Chemical weathering of such rocks depends on mineralogical content and grain size.

5.4.╇ Weathering and vertical zonation Weathering transforms a hard rock into a vertical sequence of layers. From the surface downwards these are soil, weathered rock and unweathered rock (Fig. 5.3). These layers differ from each other in texture, composition and erodibility (the capacity of a material to be eroded). In general, weathered material erodes from the top. Other terms have been used to describe the different components of the column shown in Figure 5.3. The term saprolite is used, usually in North American literature, to describe a deep section of soft weathered material consisting of clay, iron oxides and sand, some of which may still retain

Weathering in the tropics

76



Fig. 5.12

Diagrammatic sketch of soil horizons

the structure of the original rock, e.g. joint patterns. The term regolith is used to refer to the part of the column consisting of soil and the weathered rock from which all marks of rock structure have been obliterated. Regolith is sometime used loosely to indicate any soft residual material. Each of these layers can be subdivided. Soil is divided into horizons (Fig. 5.12). Weathered rock can be subdivided into different layers, for example the generalisation of Ruxton and Berry (1957) showing progressive penetration by weathering both downwards and from the outside of a rock towards its centre. The cross-country vertical profile, however, differs in total thickness or thickness of individual layers. It is possible to trace the complete profile of weathering from a hilltop to the neighbouring valley bottom but the details of the profile, especially the thickness of different layers in the subsurface, would vary even if the underlying lithology remains the same. This is primarily due to variations in the flow rate of water through the subsurface material. Rock types, frequency of slope failures, climate and vegetation are other determining factors. Given suitable conditions, e.g. gentle slopes, igneous rocks, a location in the humid tropics and an absence of human interference, weathered material may extend to 100 m or more in the subsurface. This is an extreme figure, but depth and rates of weathering in the humid tropics are strikingly more advanced than in cooler and drier areas. Engineers denote the strength of the subsurface material and changes in permeability on a scale of weathering divided into six grades (Table 5.3). It reflects the vertical sequence of changes in the weathered material.

Pans and crusts

77

Table 5.3╇Mass weathered grades as used by engineers Grade

Class

Summary description

VI V

Residual soil Completely weathered

IV

Highly weathered

III

Moderately weathered

II

Slightly weathered

I

Unweathered fresh rock

Soil with developed horizons Mostly soil; fragments of original rock or rock structure may be present. Such fragments, if present, are discoloured Discoloured rock with openings. Not more than 50 per cent of the rock mass is decomposed and can be extracted with a geological hammer. Scattered corestones may be present Mostly discoloured rock but less than half decomposed; penetration of weathering along openings, interlocked corestones Discolouration along openings, rest of the rock resemble fresh rock. Original rock, no weathering

Source:€Dearman, 1974, 1976

5.5╇ Pans and crusts Soil and soft weathered rock are eroded more easily than the hard fresh rock underneath. This is the general rule, but hard, usually thin, layers of reprecipitated material present within the soil or weathered rock may complicate the pattern. Such layers are known as pans or crusts and they may occur at any depth between the surface and fresh rock. The thickness of the hard layer varies but a measure of 10 m and more has been reported from various locations (Goudie, 1973). Dissolved material such as calcium carbonate or silica is carried by solution to the subsurface and, under suitable conditions, reprecipitated as hard layers called crusts. These crusts have three important geomorphological functions. First, being difficult to erode they protect the underlying soft soil or weathered rock. Second, they hinder soil formation, which requires vertical movement of material. Third, such crusts prevent water from penetrating to depths below. If the crust is near the surface and the rainfall is high, water tends to saturate the material on top of the crust making the upper part of the subsurface unstable. The effects are discussed in detail in Chapter 6. The general term for such pans and crusts is duricrust, but self-explanatory and specific terms are also used:€silcrete, calcrete, ferricrete, etc. As encrustment requires both dissolution and reprecipitation, seasonality in climate helps, with crusts being very common in subhumid and semi-arid tropics. Calcrete shows this well. For example, isolated nodules of calcium carbonate (locally known as kankar) are common in India in a strong seasonal climate with about 1000 mm annual rainfall. In more arid parts of India, such as Rajasthan, precipitation of calcium carbonate becomes extensive, forming bigger nodules and ultimately a continuous subsurface layer close to the surface in the driest areas. Such layers are

78



Fig. 5.13

Weathering in the tropics

A stone line in deeply weathered igneous material, Kenya called caliche. In very dry tropics, caliche may occur on the surface itself. Ferricrete is common in more humid areas. A particular type of ferricrete is laterite. Laterite is an old term (Buchanan, 1807), and associated with a characteristic profile which is not shared with all kinds of ferricrete. Laterite has a hard surface layer of iron oxide that could be nodular, pisolitic or spongy. A mottled zone lies below this surface layer, a pallid zone underneath the mottled zone and a ferrigenous bedrock underneath. Unlike common ferricrete, laterite is soft unless exposed to the atmosphere, hence the hard layer at the surface. In places, a thin layer of single or multiple clasts have been noticed within a deep weathered layer. Such layers are known as stone lines (Fig. 5.13). Stone lines have been associated with various origins. A stone line has been interpreted as lag gravel, remnant of the pavement of an old surface, a climate change indicator and the result of bioturbation. Whatever their origin, stone lines are not ubiquitous.

5.6╇ Effects of weathering There are two major results of weathering. First, the land surface is mantled with soft material that can be easily eroded and removed. Second, the soft weathered material can be transformed into soils, the base for vegetation. To a large extent, the vertical zonation of soil, weathered material and the underlying hard rock determine the nature of hillslopes and the type of downhill processes that operate on it. The weathered material is eroded from the slopes and ends up in streams as dissolved or solid load. Weathering is the first step towards eroding a land surface. Chemical weathering slows down when fresh rock is not exposed either to the atmosphere or to the circulating subsurface water. A thick cover of weathered material and soil prevents the land surface from being eroded and the hard rock underneath remains in a

79

Effects of weathering

nearly unchanged state. If the cover is thin, then the processes of erosion directly impinge on the hard rock, slowing down denudation. In between lies the optimal situation where chemical weathering is active at the hard rock surface creating new weathered material, but the regolith does not thicken significantly, as its upper part is regularly or episodically removed by erosion. Carson and Kirkby (1972) identified slopes with a thick cover of regolith as transport-limited slopes. On these slopes, weathering produces material much faster than it is removed. Transport-limited slopes are expected to be covered with a thick regolith that buries underlying surface irregularities in the hard rock, promotes vegetative growth and produces a gentle slope on the surface. On the other hand, material is transported out fast from weathering-limited slopes, where the rate of erosion is higher than the rate of weathering. Such slopes have a thin regolith, even fresh rock surfaces may be exposed, the rock exposures tend to steepen the slopes, and vegetation is usually sparse due to thin soil and active erosion. Streams draining transport-limited slopes often correlate with a high dissolved load in the local rivers whereas weathering-limited slopes tend to produce clastic material as suspended and bed load in the streams. Advanced chemical weathering should contribute solution loads to the streams of the humid tropics, but as the tropical slopes are often unstable and prone to failures (as discussed in Chapter 6), tropical rivers carry considerable amounts of solid load as well. As most of the sediment load in a river is frequently contributed by the headwaters in tectonic mountains of variable lithology, the relationship between weathering and river sediment correlates well for individual rivers but an overarching rule is difficult to construct. Gibbs (1967, 1972) studied the geochemistry of the Amazon River system in the mid1960s. Both the Amazon and the Orinoco system were studied in detail by Stallard (1985) and Stallard and Edmond (1983, 1987). Even for the 6300 km long Amazon, the sediment load is derived primarily from the headwaters in the Andes. Gibbs (1967) indicated that, in the Andes, the continuous availability of unweathered rock following removal of the weathered material by active erosion leads to progressive chemical weathering. In contrast, the very thick cover of secondary minerals on fresh rock isolates the rock surface from incoming rainwater and not much material gets to the river as suspended or bed load. Stallard (1985) and Stallard and Edmond (1983, 1987) related the chemistry of the water of the Amazon and Orinoco to lithology and relief. The amount and nature of the sediment load in a river therefore depend significantly on weathering of the bedrock of the river basin and, as most of the load comes from the headwaters, ultimately on mountain lithology and relief. High weathering rates and transport-limited slopes are often correlated with high dissolved loads of regional rivers. In contrast, rivers draining weathering-limited areas usually carry suspended load and bed load. Stallard (1985) showed this for the basins of the Amazon and the Orinoco. He divided the Amazon Basin into five major erosional regions:€the Andes, the foothills, the elevated shields, the lowlands and the lowland shields. The highest erosion rates and the biggest sediment concentration in river waters came from the tectonically active areas, not the geologically ancient low-relief areas or the lowland built by deposition. This pattern holds for a number of large river basins. Very little dissolved or solid load is found in the rivers flowing across only the lowland and the shields,

80

Weathering in the tropics

lithology permitting. Most solids are in the resistant mineral phase. Weathering reflects both bedrock and erosional regime. The nature of the sediment, however, changes along a long river due to episodic transfer and long-term storage on floodplains (Chapters 7–9). Even the feldspar-rich sediment that arrives in the Orinoco from the granites of the Guinea Shield is finally transformed into the quartz sand of the Orinoco, given the long-distance transportation and long-term storage (Johnsson et al., 1988). Weathering continues on detritus grains as they move from the headwaters to the sea. The transformation of the weathered product to the river load in the tropics is examined in detail in Chapters 7–9. Rapid rates of erosion are indicated by the presence of chemically unstable and cationrich minerals in the suspended load and bed load. Geomorphologists now have techniques based on isotopes to identify the main source areas of sediment even for very large rivers. For example, the isotopic composition of the sediment of the Brahmaputra River in Assam, India has revealed that 45 per cent of the river’s sediment comes from 4 per cent of the drainage area, the Eastern Syntaxis, where the Himalaya mountain bends and the Brahmaputra flows through a 5 km deep gorge past the Namche Barwa Peak (Singh, 2007). We need more case studies relating weathering to stream load, but a general picture has started to emerge.

5.7╇ Tropics and weathering A number of generalisations can be made regarding weathering in the tropics. 1. Given the ambient conditions of high temperature and humidity, weathering, especially chemical weathering, is accelerated in the humid tropics, unless the underlying rock is chemically inert. 2. The slopes of the humid tropics are usually underlain by regolith, consisting of soil and weathered rock, which is several metres thick. 3. The weathered rock is normally clay-rich and kaolinite is the common clay mineral. 4. The weathered rock is not necessarily homogeneous and duricrust layers are common. 5. The impediment caused by such crusts to the downward flow of water leads to slope failures, as discussed in Chapter 6. 6. Tropical streams may carry a significant amount of dissolved load, but these streams mainly carry clasts of varying size as suspended and bed load. Most of the suspended and bed load is derived from the mountains at the headwaters of the basin. 7. The river sediment tends to be episodically transferred with periodic deposition on the floodplains, where the grains are weathered further. 8. The accelerated chemical weathering of the humid tropics influences subsequent slope and channel processes. 9. Weathering in the arid tropics has a different pattern, physical disintegration being the principal weathering process.

81

Questions

Questions 1. Name a mineral that exists both as primary and secondary types. How is it formed? 2. State any reasonable annual rainfall figures associated with the weathering of a plagioclase feldspar (albite) to a) gibbsite:€mm b) smectite mm c) kaolinite mm. Now name three locations in the tropics to go with your rainfall figures. Be specific with your choice of location. Broad answers such as 'Africa' are unacceptable. 3. Imagine you are in Java (Indonesia) walking down a steep volcanic slope. What kind of volcanic rock would you expect there? The area gets about 1800 mm of rain a year. Do you expect to walk on the same clay mineral all the way? Justify your answer. 4. On which side of the Hawaiian Islands should you expect bauxite? Why? 5. Why should an increase in the specific surface of a rock fragment raise its potential for chemical weathering? 6. Suppose you have hillslopes in the humid tropics underlain by the following rock types. What size material should you expect on the slopes after advanced weathering? Why? What type of sediment load should you expect in rivers draining such areas? a) basalt b) granite c) sandstone d) limestone. 7. You are in eastern Puerto Rico. The slopes are in volcanic rocks weathered to boulders set in red clay. What would be the sediment load of a local stream if a) a landslide occurs? b) a large flood comes down the river? c) headwaters of the river receive augmented drainage? As a result, the average discharge increases significantly but not flood peaks. 8. Should there be any difference in appearance between a weathering-limited slope and a transport-limited slope? If so, how would they differ? Provide brief descriptions and two diagrammatic sketches. 9. How do pans and crusts form in the subsurface? Do you anticipate any significant role played by these features on erosion of slopes? 10.╇ What happens after weathering?

6

Slopes:€forms and processes

I have climbed and ruminated upon too many great bornhardts, in company with the leopard and the baboon, to believe that these most powerful of landforms, glorious in the sun and rain alike, ever originated foetally within the dark body of the earth. The leopard and the baboon don’t believe so either. L. C. King

6.1╇ Properties of a slope The word slope indicates both (1) an inclined unit on the surface of the Earth and (2) the measured level of such inclination. The surface of the Earth is a combination of slopes of various kinds. This assemblage of slopes mostly consists of hillcrests, valleyside slopes and slopes along streams orthogonal to the termination of valleyside slopes. As Dunne and Leopold (1978) have described it, hillslopes cover virtually the entire landscape. A slope is probably visualised best as a profile drawn from the hillcrest to the valley bottom. Both water and surficial sediment are transported under gravity downslope. Sediment commonly takes several moves to traverse the entire length of a hillslope, and between such moves it rests at progressively downslope locations. The surface of a slope therefore reflects this pattern of storage and transfer of material. The geometry of a slope is expressed by three measures:€gradient, length and width. The gradient of a slope is measured in degrees, ratios or percentages, and its length and width in linear measures such as metres. A complete description of the slope requires details of its geometry and a hillcrest to valley bottom cross-section showing a vertical sequence of soils, weathered rock and bedrock along the profile. Accumulation of soft regolith towards the top of the slope and active erosion by water towards its base tend to give a slope profile a convexo-concave appearance, but not always. For example, a weathering-limited slope, as described in Chapter 5, tends to have a steep, almost vertical segment below the crest of the slope. Material falling off this steep segment comes to rest below it, giving the next downslope section a straight profile. In 1942, Wood proposed a four-fold division of a slope profile. In the downward direction these are waxing slope (convex crest), free face (vertical), constant slope (straight) and waning slope (concave). King proposed a similar four-fold division in 1962 but with different terms:€crest, scarp, debris slope and pediment. These units differ not only in form (appearance) but also in subsurface material and geomorphic processes that operate on the surface (Fig. 6.1). In a transport-limited slope, accumulation of weathered material covers up the free face (scarp) and the profile of the slope is modified to a convexo-concave appearance. 82



83

Properties of a slope

Fig. 6.1

The sub-units of Wood and King. Note that with time the straight debris slope (also called a talus slope or constant slope) increases in area due to the accumulation of material arriving from the top. It then proceeds to bury the scarp face and change the original appearance of the slope. The size of the talus or debris accumulation is reduced when its constituent material is removed episodically by a slide or flow. Its appearance depends not only on the nature of its constituent material, but also on the time elapsed since the last removal or significant addition of material from the top Any given point on a slope displays a balance between two forces over an interval of time:€(1) the erosive force that removes material from the surface and moves it downslope, and (2) the resisting force that opposes its removal. The erosive force usually depends on two gravity-driven processes:€movement of material en masse and running water. The first of these is collectively termed mass movement or mass wasting. Both erosive forces increase if the slope is high and water is present. The upper parts of high mountains, such as the Andes or the Himalaya, are also eroded by ice, a third operating process on selective slopes. The high mountains are discussed later as a special case in Chapter 13. The resisting

84

Slopes:€forms and processes

force essentially depends on the slope material. Hard rock with little structural weakness close to the surface is difficult to erode. Erosive processes in general are also limited where an abundance of vegetation binds the regolith together. In contrast, loose regolith, tectonically shattered rocks and bare slopes tend to accelerate slope erosion. In sum, slopes are failure-prone when the following properties are high: • gradient • weathered or soft rock in the subsurface • geological structures susceptible to erosion • presence of water on the surface or in the subsurface • absence of vegetation • tectonic disturbance • undercutting of the base of the slope by a river or the sea. For example, tropical cyclones passing over hilly areas will lead to multiple slope failures. The high winds will destroy the vegetative cover and then intense high rainfall will lead to slope failures. Similarly, steep slopes in an area of weak rocks and tectonic disturbances, e.g. in the vicinity of an active plate boundary, are hazardous. Anthropogenic activities may also increase the potential for slope failures. Destruction of vegetation, local steepening of slopes or construction of houses on the mid-slope increasing the weight on the slope may increase slope failures.

6.2╇ Mass movement on hillslopes Mass movement is the general term for movement of material downslope under gravity, usually in a relatively dry state. Three variables characterise such movements. First, the nature of movement is an important characteristic. The material may fall down a cliff, slide down a surface, move as a hyperconcentrated flow when lubricated by water, overcome physical barriers by riding a cushion of air mixed with fine dust particles, etc. Second, the nature of material that moves is a crucial property. The moving mass may range in texture from boulders to clay and it usually occurs in an unsorted mixed state. The third variable is the speed of movement. The moving material may creep downslope very slowly. In contrast, certain types of mass movement are near-instantaneous. Various classifications have been proposed to organise mass movements into classes based on these three variables. Of these classifications, the better-known ones have been proposed by Sharpe (1938), Varnes (1958; 1978), Hutchinson (1968), Carson and Kirkby (1972) and Nemcok et al. (1972). These attempts vary in the criteria used for classification and in their complexity. Some also require more careful field observations than others. We will use a simple classification system, adequate for our purpose. This classification (Table 6.1) is based on Dunne and Leopold (1978) and has been used before in Kale and Gupta (2001). It is based primarily on the nature of movement and each type is further subdivided according to (1) the property of the moving material, (2) the presence or absence of moisture in the material and, in certain cases, (3) the speed of movement.

Mass movement on hillslopes

85

Table 6.1╇ Types of mass movement Types

Subtypes

Fall

Rock fall Debris fall Topples Planar slide (often further categorised as rock slide, debris slide, mudslide, etc., depending on the material) Rotational slide, also known as slump (always in fine material) Debris avalanche Debris flow Earth flow Mudflow Lahar Solifluction Soil creep

Slide

Flow

Creep

Modified from Dunne and Leopold, 1978

6.2.1╇ Falls This term is used when a mass of material breaks off from a near-vertical or vertical slope and falls through the air to the lower slopes (Fig. 6.2). If this falling material consists of a single chunk or several pieces of rock, it is a rock fall. The falling mass may also be a mixture that includes material from boulders to fine particles, in which case it is known as a debris fall (Table 6.1). The movement is very fast, almost instantaneous. The coarser components of a fallen mass (particularly the broken bits of rock) may roll or bound further downslope after landing. Rock falls from a vertical cliff of hard rocks could be large in size or may fall from a considerable height. Soft rocks or unindurated material can only build a steep slope of limited height and therefore falls in such material are smaller in both size and height. Potential conditions for rock falls exist when the rock is well jointed and weakened over time by several factors such as alternate freezing and thawing of the slope material, intrusion of tree roots or the development of stress-release joints as explained below. A trigger action for starting a fall usually involves a well-jointed rock cliff being vigorously shaken by earthquakes or undercut by a river or the sea. The undercutting of the lower part of the slope makes it unstable. Falls commonly occur in seismic high mountains such as the Himalaya or Andes. The potential for falls in these high mountains is also higher because the present glaciers used to extend further down the valleys during ice ages. The retreat of glaciers up-valley in the warmer Holocene removed ice masses from the lower valleys, removing the lateral and downward pressure of ice against the rock of valleysides and valley bottoms. This release of pressure fractured sidewalls and uparched valley floors, and led to the development of new joints, splitting rocks into slabs and raising the potential of rock falls. Walls of mountain valleys therefore also fail as a delayed effect of past glaciation.

86



Fig. 6.2

Slopes:€forms and processes

Different types of slope failure. After Dunne and Leopold, 1978

Topples is a special case. If joints are near-parallel to the rock face or bedding planes are in a similar position due to the folding of the strata, rock slabs often fail along the cross-joints and topple forward. Igneous rocks such as basalts or dolerites with columnar joints and folded bedded sedimentary rocks or metamorphic slates or schists tend to topple (Selby, 1993). A mountain slope affected by rock or debris fall therefore would have a near-vertical section with screes of fallen material below. This is the free face or scarp and constant or debris slope mentioned earlier. The material on the scree slope (also known as talus) tends

87

Mass movement on hillslopes

to be unstable and fails again, to provide streams at the bottom of the valley with coarse bed load. At times, boulders released by rock falls may also bound downslope to reach valley bottom streams.

6.2.2╇ Slides Slides are failures along a surface of separation, over which the failed mass slides downhill. Translational (planar) slides indicate movement en masse over a surface separating two different types of material such as regolith and unweathered rock or shallow surficial material and a subsurface encrustation (Fig. 6.2). Translational slides occur in a range of material and hence it could be qualified by the property of the slid material. For example, a planar slide can be identified as a rockslide (contains broken bits of rock), debris slide (a mixture of rock and regolith) or mudslide (clay-size material, usually water is present). In contrast to translational slides, rotational slides occur in deeply weathered, near-homogeneous fine material (Fig. 6.2). The type and level of weathering of the slope material therefore would be a determining factor, and the slid material, when it comes to rest, could be at any state between nearly undeformed and well deformed. Slides vary widely in size, distance travelled and rate of movement. Dunne and Leopold (1978) mentioned that the slid material has been reported from cubic metres to cubic kilometres in volume and the distance travelled between metres and kilometres. The rate of movement has been found to range from several centimetres a year (very slow) to several centimetres per second (fast). The slower rockslides are also called rock glides. The rate of movement depends on three factors: • gradient • amount of moisture in the slid material • presence of lubrication. Antecedent rainfall and trigger rainfall (Box 6.1) in combination determine the last two factors and lead to the crossing of the erosional threshold. An erosional threshold is the condition, level or stage at which a striking change is produced, such as a slide, which significantly modifies the landscape (Schumm, 1977). There are three types of erosional threshold:€extrinsic, intrinsic and geomorphic. An extrinsic threshold is crossed when, for example, an increase in velocity and depth of a river allows it to pick up sediment previously resting on the channel bed. An intrinsic threshold is crossed with ongoing changes in the property of material. For example, progressive weathering of slope materials may eventually cause the slope to fail. Geomorphic thresholds are crossed when a landform changes over time to a state of incipient instability. For example, undercutting of a rock cliff may result in the collapse of part of the cliff face (Schumm, 1977). A combination of antecedent and trigger rainfalls results in crossing of both intrinsic and extrinsic thresholds. In translational slides the failed mass moves along a subsurface plane of failure that tends to coincide with (1) a structural surface such as a bedding or foliation plane, (2) the contact between regolith and unweathered rock with little cohesion between the two, or (3) a subsurface pan such as an iron or silica encrustment layer. Cracks and openings in the material above the plane of failure allow rapid penetration of water which is interrupted or slowed

88



Box 6.1

Slopes:€forms and processes

Antecedent and trigger rainfall Slope material fails when it is saturated with water and the increased weight due to the stored moisture is more than the resisting force at the plane of failure, a weak zone to begin with. Moreover, the presence of moisture lubricates the plane of failure. The moisture accumulates over time. Antecedent rainfall refers to the total amount of rain that has fallen over a selected number of days prior to the failure. The number of days selected depends primarily on location, and usually is between 5 and 30. Moisture from antecedent rainfall saturates the soil to some extent, following which water from a large rainfall event (the trigger rainfall), supersaturates both soil and weathered rock, and a failure occurs. The trigger rainfall has to be large if the antecedent water in the soil is low, but if the soil is near saturation, a small amount of rain may act as the trigger. This explains why slopes fail in intense rain from a tropical cyclone (measured as 102 mm in a day) without much antecedent moisture. On the other hand, slopes often tend to fail late in the rainy season when the antecedent moisture is high and only a small amount of trigger rainfall is required. Numerous landslides occur when both are high. Pitts (1992) has shown that slopes fail in a number of places in western Singapore when rainfall thresholds are exceeded, but not earlier. Gabet et al. (2004) studied the Himalaya Mountains in Nepal, which are prone to slope failures because of steep slopes, weathered bedrock and intense monsoonal rainfall occurring in combination. Based on three years of fieldwork in a small drainage basin, they have shown that a number of small landslides start on these slopes from the beginning of the wet monsoon season and larger ones usually after the monsoon has been in place for some time. They observed slope failures in the basin of the 136 km2 drainage basin of the Khudi Khola on the southern flank of the Annapurna Himalaya in Nepal. It is a rugged tectonic catchment with a mean elevation of 2565 m. Heavy seasonal rainfall (3000–5000 mm yr−1) occurs over the basin due to monsoondriven moisture impinging on the mountains. The average soil depth is only 50 cm but the bedrock is deeply weathered and permeable. Thus the regolith can store water up to a certain amount (Fig. 6.3). Gabet et al. (2004) determined that landslides are triggered after about 860 mm of rain has fallen from the beginning of the monsoon to fill the regolith. This is the antecedent rainfall. As the regolith is gradually filled,

Fig. 6.3â•… Role of antecedent and trigger rainfall in landslides in the Nepal Himalayas. From Gabet et al., 2004. With permission from Elsevier

89



Box 6.1

Mass movement on hillslopes

Continued the slope begins to fail with trigger rainfalls, the threshold amount of which drops with time until the required amount stabilises at 11 mm per day (Fig. 6.3). Thus two distinct rainfall thresholds, the seasonal accumulation (antecedent rain) and a daily total (trigger rain), are required to initiate slope failures. Larsen and Simon (1993) reached a similar conclusion regarding slope failures in the hills of Puerto Rico. There, landslides tended to cluster towards the end of the hurricane season when intense rain falls on saturated ground. Gabet et al. (2004) mentioned suspended sediment waves travelling at the rate of 2–3 ms−1 in a local river. In this kind of hilly or mountainous environment with large storm rainfalls, the material coming off slopes is simultaneously transferred out of the small drainage basins to major rivers or, as in the case of islands like Puerto Rico, to the sea. Only the very coarse material may remain in the valley (Chapter 7).

down at the structural surface, contact surface or encrustment layer, leading to accumulation of water in the mass overlying the potential plane of failure. The slope then fails due to both the extra weight of water and lubrication. As gullies advance, slides in alluvial or colluvial material may occur due to pressure release along the channel walls (Fig. 6.4) Rotational slides, also known as slumps, tend to occur in deep homogeneous material such as a thick layer of clayey silt formed by advanced chemical weathering. Slumps, therefore, have been associated with the humid tropics, but the weathered layer in the humid tropics is not always homogeneous as it frequently includes impermeable layers such as iron pans or silcrete. Shallow planar slides are therefore also common in the humid tropics. In a slump, the failed material is deformed as the mass above the ruptured surface is rotated backwards giving rise to tilting, formation of small faulted blocks at the back of the slide, a hummocky surface, ponded drainage and a curving rampart at the toe (front end) of the slide (Fig. 6.2). If the water ponded between the blocks at the back of the slide seeps through the weakened failed material, subsequent failures may occur. Slumps tend to happen when water from heavy rainfall enters the subsurface, usually through an opening into a thick, fine homogeneous material, or when a river undercuts its bank in thick beds of fine material. In steep mountains, very large slides may come down a mountain slope, cross the valley bottom and ride up the opposite hillslope for a limited distance. The ground between the source and the final location of the slid material surprisingly does not always show much disruption along the track of the slide. Even nearly undisturbed snow, fallen prior to the failure, has been recorded. Shreve (1965) studied the huge Blackhawk landslide in southern California, and suggested that as the huge slide came downhill, it began to ride a cushion of compressed air in the manner of a hovercraft, which allowed the material to travel a large distance. The cushioning effect has also been explained by the production of steam from water vaporised by the heat of friction at the base of the sliding material (Habib, 1975). The concept of sturtzstrom was proposed by Hsü (1975). According to this concept, the innumerable solid units inside the moving mass of a cloud of dust exchange kinetic energy by collision. This energy keeps the particles separate and thus the mass moves like a fluid (fluidisation) with very little internal friction. An earlier version was provided by Heim in 1882 after studying a large failure at Elm, Switzerland. There the failed material travelled for about 3 km at a velocity exceeding 160 km per hour. For this to occur, the flow needs to be cushioned.

90



Fig. 6.4

Slopes:€forms and processes

Slides along headwalls of a gully cut into the alluvial fill of an intermontane basin, Upper Araguaia Basin, eastern Brazil. Photograph: A. Gupta

6.2.3╇ Flows Unlike slides where the deformation of the failed mass is limited, the entire mass is deformed internally in flows and the moving mass behaves like a viscous fluid. Flows also are accompanied by higher volumes of moisture and fine-grained material. Certain classifications recognise very large slides such as sturtzstorms as dry flows and identify them as rock avalanches if the primary constituent is broken rock. When the constituents are mixed, such failures are called debris avalanches. The term avalanche recognises the very high downslope velocity of these failures. Debris flows consist of fast-moving solid units of varying texture including fine material and relatively more moisture and air mixed with the solids in the moving mass. Debris flows are common in the mountains of the humid and arid tropics, where high-intensity rain usually causes a series of debris flows to come down the valleys of small steep streams, especially if sediment is already stored in the valley bottoms. For example, the hillslopes of the Caribbean islands are marked by numerous debris flows after high rainfall or the passage of a hurricane (Fig. 6.5). Rapid snowmelt in the high mountains, collapses of glacial ice or a moraine with a lake behind it, and overflows from or collapse of volcanic crater lakes all have the potential for creating debris flows. A debris flow commonly starts as a slide or a fall of debris at the head of a small valley draining hilly areas. The addition of more moisture downslope, and probably dilatancy, transforms the slide into a flow. Debris flows usually come down channels of small streams, but, as Costa (1984) states, debris flows are quite capable of eroding their own channels across sloping surfaces. They act as an erosive process, especially over steep slopes, and may erode up to several metres in bedrock and more in loose material, breaking tree trunks at ground level. The destruction along the track is related to the velocity of the flow, but such destruction is seldom seen where the debris flow terminates. Morphologically, debris flows

91



Fig. 6.5

Mass movement on hillslopes

Debris flow after a heavy rainstorm in the Blue Mountains, Mahogany Vale, eastern Jamaica. Photograph: A. Gupta

terminate in a curved rampart of large boulders; flows behind the rampart are bounded by levees of boulders, displaying hummocky tops. Individual flows measure between several metres to about a hundred metres wide with a maximum reported depth of 12 m. Usually the flows are shallow and less than 5 m deep (Selby, 1993; Costa, 1984). Moving debris flows have been videotaped, displaying a mass like wet concrete flowing over the ground surface. Pebbles, cobbles and boulders collide with each other with banging noises as they are tossed round within the mass and carried downslope. Debris flows move in a series of surges. The surges arrive enriched with boulders and watery turbulent slurries rich in suspended sediment, but only a few boulders arrive between the surges (Costa, 1984). Observed velocities range from 0.5 to 20 ms−1 and they are fast over steep slopes. Even over low gradients, debris flows are capable of travelling long distances, especially those with a high clay content. Debris flows not only erode their tracks but also bring down a huge amount of material, which often reaches valley flats at the base of the slope, augmenting the sediment load of streams. They also build parts of alluvial fans at the mouth of valleys emerging from mountains. The solid content of a debris flow varies between 35 and 90 per cent and its water content between 10 and >30 per cent by weight. Unless the flow originates from a clay-rich source, the clay content of debris flows is quite low, below 5 per cent. The increased bulk density supports the transfer of boulders. Excellent general accounts of debris flows have been provided by Costa (1984, 1988). Slow, dry flows of clay-sized material are earthflows. These usually travel less than a metre annually. Earthflows are derived from clay-rich rocks or clayey regolith. They usually start by a slope failure that releases clay which then spreads over a low gradient as a toe-shaped feature. In contrast, mudflows are formed when a fine-grained failed material incorporates a large volume of water. Mudflows may occur in the arid tropics when rapid rainwash from cloudbursts sends a large volume of mud into river gorges. They can carry



92

Slopes:€forms and processes

Fig. 6.6

Lahar filling a stream channel on the southern slope of the Merapi Volcano, Java, Indonesia. The parapet of a bridge which used to take the traffic over the river indicates the volume of the fill. The material is used in construction, explaining the vehicle in the channel. Note the texture of the material. From Gupta and Ashar, 1988. By permission of Wiley. Photograph: A. Gupta material as coarse as boulders because of their thickness and viscosity. On reaching flatter ground beyond the mountains, mudflows tend to spread out and terminate. Lahars are mudflows of volcanic origin (see Chapter 14). The name is adopted from Bahasa Indonesia, lahars being extremely common around Indonesian volcanoes. Lahars start by the addition of a large volume of water to an accumulation of volcanic material. Slopes of explosive volcanoes, as in the Andes Mountains and on the islands of Indonesia and the Philippines, are usually covered with old pyroclastic material ranging in texture from fine-grained ash to volcanic boulders. Thicker layers accumulate in river channels radiating out of a central volcanic cone. Large and intense falls of rain, as expected in tropical latitudes, create lahars from the stored pyroclastic sediment which then move downslope, especially along the river valleys. Pyroclastic flows travelling down a valley and mixing with the water in the channel may also start lahars. In high mountains, such as the Andes, lahars may be caused by snowmelt. Lahars travel fast and bring down huge volumes of volcanic material to bury the pre-existing topography and settlements (Fig. 6.6). They are extremely dangerous in terms of loss of life and property. The 1991 and 1992 eruptions of Mount Pinatubo in Luzon, the Philippines, deposited huge amounts of pyroclastic deposits and lahars (Newhall and Punongbayan, 1996; Nossin, 2005).

6.2.4╇ Creep Creep is the slow downslope movement of the upper parts of the surficial material. It is frequently described as soil creep because it is the soil part of a weathered profile that tends to move imperceptively downslope. The shallow upper part of a weathered profile may move due to three reasons:€slow plastic deformation of the moving mass under its weight;

Running water on hillslopes

93



Fig. 6.7

Running water on hillslopes alternate wetting and drying; and alternate freezing and thawing. The last characteristic is limited in the tropics to the high mountains. It is a very slow movement, usually unnoticed, and continues for a very long time. Soil creep is common under the tropical rain forest, especially if a thick weathered layer is present and, although imperceptible, it may move a considerable volume of material over time to a stream at the bottom of the valley. Water, by reducing the friction between grains, tends to speed up soil creep but even then it is a very slow and imperceptible process. Its morphological manifestation breaks up a hillslope into a series of terracettes, and piles up material against the base of an upstanding object such as a tree trunk or post. Over time, such upstanding objects may also bend downslope in an area of continuing soil creep. Thus the presence of a soil creep is identified. Soil creep is not a hazard, except where the creeping material is unstable, e.g. a clay-rich plastic soil.

6.3╇ Running water on hillslopes Splash erosion, as discussed in Chapter 4, is caused by intense rain with large raindrops splashing on bare or near bare ground. Splash erosion also depends on the nature of the surface material. Sand or silt is disturbed easily, but a soil that contains a high amount of any binding material such as organics, clay and CaCO3 reduces splashing. Running water, however, erodes slope material more effectively. Water accumulates on the surface when rain intensity is higher than infiltration (the steady rate at which water enters the subsurface). A thin film of water thus accumulates on the surface before flowing downhill. This is the Horton overland flow. If, however, all the empty space in the soil (i.e. pores between soil grains) is already full of water, rainwater cannot enter the subsurface. Instead it emerges to flow on the surface. This is the saturation overland flow (Fig. 6.7). Horton overland flow tends to occur over most of the area covered by a hillslope, the saturation overland flow is common towards the base of the slope and near streams that flow at the bottom of valleys. Exceptions may occur to this general rule with unusual breaks in gradient or particular types of soil. Erosion, i.e. removal of grains of material on slopes by running water, requires a minimum velocity which is dependent on the depth of flowing water. Thus a belt of no erosion

94

Slopes:€forms and processes

occurs near the crest below which the running water reaches sufficient depth and velocity to remove surface material. The length of the belt of no erosion depends on the gradient and soil properties. The running water downslope is known as sheetwash. Sheetwash is about a centimetre or so in depth but not a uniform spread, as small threads of water with deeper and faster flow move across the sheet. These threads are effective in moving surface material, especially those loosened up by splash erosion. If the sheet of water is deeper (up to 1 m) and faster because of high and intensive rainfall, it is called sheetflood. Small channels (rills) are eroded where, over time, flow is concentrated and some of the small channels deepened. Sheetwash is thus replaced by rills downslope. Rills tend to enlarge into gullies, which are bigger rills. Dunne and Leopold (1978) have defined gullies to be at least 30 cm (one foot) deep and wide. Bare slopes or slopes with limited vegetation tend to develop a network of rills and gullies over time (Fig. 4.1). Dunne and Leopold (1978), after reviewing a number of case studies, concluded that in many cases flowing sheets of water remove more sediment than rills or gullies. The nature of a specific location (rainfall intensity and amount, gradient, soil properties, vegetation) determines which type of erosion would be more effective at a particular location. Material eroded and transported downslope ends up at the base where the gradient lessens and the velocity of the running water drops. A fan-shaped accumulation of sediment tends to occur at the mouth of gullies, which is removed by the valley bottom stream at intervals.

6.4╇ Storage and transfer of surficial material on tropical slopes A range of slope processes occur in the tropics, of which some are more common. The efficacy and frequency of the processes depend on the location. Rocks, relief and rainfall are crucial background factors. Slope processes vary between the humid and arid tropics. In the humid tropics, slopes are commonly under a thick layer of soil and weathered rock due to effective chemical weathering. The slope processes that operate are commonly moisture-driven as illustrated by Gabet et al. (2004) and slopes are expected to fail periodically, especially in the wet season. In contrast, weathering is limited on the slopes of the arid tropics which are usually under a thin layer of clasts of broken rock and sand. Slope processes are rare and episodic, and tend to occur following cloudbursts when the intense precipitation is heavy enough to cause debris flows, etc. without any antecedent rainfall. A number of case studies concerning the storage and transfer of surficial material on tropical slopes have recently been published, allowing generalisations on slope processes in certain environments. It is not possible to cover all types of slope, but we will discuss two cases for which a reasonable amount of information is available: • slopes of the tropical rain forests • steep slopes of hills and low mountains. Slopes of the arid tropics, volcanic landforms and anthropogenically disturbed areas are discussed later, in Chapters 12, 14 and 17 respectively.

95

Storage and transfer of surficial material

6.4.1╇ Tropical rain forest slopes Tropical rain forests are best developed on lower slopes of mountains and near the contact between hills and plains. In general this is an environment of rolling hills with a thick soil and a deeply weathered layer. The thickness and nature of surficial material, however, depend on local lithology and steepness. A number of case studies have been carried out on slopes of the rain forest, for example in eastern Puerto Rico, Southeast Asia and northern Australia. In a forest, rainwater, after reaching the canopy level where part of it is intercepted and evaporated back to the atmosphere, moves as (Fig. 4.4) • throughfall • fall from the canopy • stemflow • flow on leaf-littered surface, for a brief period after intense rain • flow through the subsurface via soil pores and shallow pipes. A description of such movement of water is in Sidle et al. (2000). In the rain forests of the Danum Valley in Sabah, East Malaysia where the annual rainfall approaches 3600 mm, rainwater ends up as interception (17.5 per cent of the total rainfall) throughfall (80.7 per cent) stemflow (1.9 per cent). According to this study (Sinun et al., 1992), between 2.0 and 2.5 per cent of the rain travels as overland flow on reaching the ground. The rest infiltrates and probably flows through soil pipes. Both throughfall and fall from the canopy may cause splash erosion, loosening the upper soil and preparing it for removal in surface runoff. Given the brief and intense rainfall, widespread overland flow lasting for short periods is common in the rain forests. The brief duration of flow on a litter of leaves overlying humus-rich clays is not particularly erosive. The movement of slope material under gravity is also limited, and commonly occurs by surface creep. Apart from slow but continuous creep, sediment moves downslope for brief periods in surface runoff (sheetflow) and through subsurface pipes emerging at the face of vertical banks cut in the regolith. Under well-developed forests, the amount of sediment that is eroded and removed is limited, usually not exceeding 102 tkm−2yr−1. Such a rate is unlikely to either form erosional scars or build extensive depositional features. The sediment commonly consists of sand, clay and oxides and hydroxides of iron and aluminium, particularly on igneous rocks. Gravels (up to boulder size) may occur on slopes as lag or exposed corestones. Rills and gullies, if formed, are few in number until the vegetation cover is destroyed. Such is the nature of the slopes of the forested hills of the tropics. With the destruction of vegetation, however, either in storms such as tropical cyclones or by anthropogenic clearing, sediment loss can be extreme. Such cases are discussed later (Chapter 17), but just to illustrate, Anderson and Spencer (1991) in a review of tropical rain forest conditions refer to a case study in Java that reported an increase in sediment yield from 30 to 1590 tkm−2yr−1, an increase of 53 times. Such a figure is comparable to erosion in natural disasters, such as tropical cyclones, where destruction of the vegetation cover in

96

Slopes:€forms and processes

high winds is followed by heavy rainfall. In all these cases, the slopes become dissected by landslide scars and gullies. Elsewhere, the rain forest is a quiet environment, except in tectonically unstable areas. Soil creep is the dominant mode of transportation down slope and, unless lithology intervenes, a convex–concave slope tends to develop over a thick layer of soil and weathered rock.

6.4.2╇ Steep slopes of hills and low mountains Traditionally, given the thick layer of soil and weathered rock in the humid tropics, rotational slides (slumps) were considered as the common type of failure on hillslopes. The failures, however, are just as often translational or planar. The plane of failure frequently turns out to be a relict structural surface or an encrustment layer that impedes the downward passage of subsurface water, as explained in section 6.2.2. Usually, when slopes fail in this fashion, a number of small and shallow translational slides are seen on a hillslope, often at a comparable elevation. Antecedent and trigger rainfalls (Gabet et al., 2004) are important for both translational and rotational slides (Box 6.1). Simon et al. (1990) reviewed failures and developments of hillslopes in the low mountains of eastern Puerto Rico. They came to certain conclusions which may be extended to cover most of the hillslopes of the humid tropics. Shallow translational failures are probably the primary mechanism of hillslope denudation. This happens because of (1) deeply weathered rocks and thick soils, (2) occurrence of discontinuous zones of contrasting density and permeability throughout such layers, and (3) presence of relict features and tensile stresses. Both pathways and impedances for the movement of surface water therefore occur in the regolith. Deep failures do happen, but shallow failures are common. Simon et al. (1990) measured the depth of failure in eastern Puerto Rico as between 3 and 7 m. Slopes around the heads of channels (cove slopes) tend to fail when the saturation level is high and subsurface water converges towards the channel. Although elevations in eastern Puerto Rico rise to beyond 1000 m, slopes usually fail at between 600 and 800 m, probably due to water accumulated from the upper slopes crossing the threshold of failure at this elevation band. Such elevation-selective failures may happen on slopes in other areas. These events combine to determine the characteristic form of the local hills. Large-scale failures also occur and these are normally associated with tectonically active areas or steep slopes within the tropical cyclone belt. The worst scenario is a combination, such as a tropical cyclone in action over steep and tectonically unstable hills. Simonett (1967), working in Papua New Guinea, estimated that the denudation rate of the earthquakeprone Toricelli Mountains where landslides are common is about 100 cm in 1000 years. In comparison, the denudation rate in the neighbouring Betwani Mountains, less affected by earthquakes, is 22 cm per 1000 years. Steep tectonic mountains are prone to slope failures that contribute a large proportion of the sediment that travels down major rivers whose headwaters originate in mountains associated with active plate margins. Common examples of failures from steep slopes in the literature are from the Himalaya, Andes, mountainous Papua New Guinea, Hong Kong, the escarpment zone of eastern Brazil, the Caribbean and Taiwan. Undoubtedly steep slopes fail almost everywhere in the tropics, but these areas have been better studied.



97

Storage and transfer of surficial material

Fig. 6.8

Landslide in the Blue Mountains, eastern Jamaica overlooking the Yallahs River at Mavis Bank. A number of factors control the slopes of the Blue Mountains. A regional fault runs along the hillside; the slopes are in alternate beds of greywacke and shale; and the area is periodically affected by storm rainfall, some of which is caused by hurricanes. Note also that, as a result, large quantities of very coarse material fall into the Yallahs, which can remove and redistribute most of it downstream if it is in flood from hurricane rain simultaneously with slope failures. Photograph: A. Gupta

Slope failures in the tropical mountains (Fig. 6.8) tend to contribute most of the river sediment, irrespective of the scale of the basin. Larsen and Torres Sánchez (1992) calculated that 81 per cent of the total sediment transported out of the Mameyes River basin (17.8 km2) in eastern Puerto Rico comes via mass movements. In this paper they also report the effect of a 1989 hurricane (Hurricane Hugo) which had a sustained wind speed exceeding 46 ms−1 and a 24-hour rainfall total ranging between 100 and 339€mm. Average rainfall intensity varied between 34 and 39 mmhr−1. Up to 97 per cent of the trees were defoliated in certain river basins; a very high proportion of these were destroyed. This was followed by hurricane rain falling on exposed ground. More than 400 landslides occurred on the slopes facing the approaching hurricane. Two hundred and eighty-five landslides were mapped within 64 km2 from aerial photographs. Field mapping after similar storms has indicated even higher figures in Jamaica, approaching 60 mass movements per km2, more than half of which being reactivated slope failures (Maharaj, 1993). The size of such slope failures varies. The biggest one mapped in Puerto Rico was a debris slide in which about 30 000 m3 of material was moved into the Rio Cubuy, exposing unweathered bedrock along the erosional scar. It happened three days after the hurricane. Most of the landslides occurred within the area of heavy rain, bounded by 200 mm isohyets in this case (Larsen and Torres Sánchez, 1992). Seventy-five per cent of landslides occurred on forested hillslopes, which indicates that vegetation cover may not be a restricting factor in large storms. Shallow soil slips and debris flows were identified as the primary types of failure. The failure plane of soil slips tends to be at the contact between soil and

98



Fig. 6.9

Slopes:€forms and processes

Debris flows and coarse channel sediment following the December 1999 flood in northern Venezuela. Photograph: M. C. Larsen, US Geological Survey weathered rock. Larsen and Torres Sánchez calculated an approximate denudation rate of 164 mm per 1000 years, taking into account the estimated volume eroded by landslides and the 10-year recurrence interval rainfall in the hurricane. High-magnitude formative events like this are frequent (once in 5–10 years) in certain humid tropical regions with steep slopes, indicating the effectiveness of climate in river basin geomorphology of these areas. Short-duration, high-intensity rainfalls regularly cause shallow slips and debris flows in humid tropical mountains. Bigger events such as tropical storms add larger and deeper slides to it at intervals of 5–10-years. These high-magnitude events erode huge parts of the hillslopes, contributing a large and coarse sediment load to the streams and, for islands or coastal mountains, to the coast and offshore regions. The effect of the excessive water and sediment loading on streams is discussed in Chapter 7. Larsen et al. (2001a, 2001b) described a sequence of rainfall-triggered debris flows and flash floods in northern Venezuela. The region received over 1200 mm of rain in the first 16€days of December 1999, including 911 mm between 14 and 16 December when the slopes were already saturated. This was a catastrophe with lives lost (estimated between 15 000 and 30 000) and property damaged, but the event also illustrated the role these episodic events play in modifying mountain slopes, stream channels, coastal features and river-mouth fans (Fig. 6.9). Creeps, planar slides and debris flows seem to be the dominant processes on tropical slopes, although all the types discussed earlier do occur in appropriate environments. For example, the type of mudflows known as lahars erodes and reforms volcanic hillslopes. Ice, steep slopes and tectonically shattered rocks cause a range of slope failures on high mountains such as the Himalaya or Andes, which is discussed in a separate chapter on tropical mountains (Chapter 13). It is, however, worthwhile to recall that the sediment load of rivers essentially owes its origin and characteristics to the tectonics, volcanism, weathering and mass movements on mountain slopes, often hundreds or thousands of kilometres away from the sea.

99

Questions

6.5╇ A general description of tropical slopes Slopes in the tropics, as everywhere else, are determined simultaneously by lithology, tectonics and a combination of geomorphic processes. The processes are climate-dependent to a large extent and vary across the tropics. Thus slopes on convergent plate margins tend to undergo rapid changes, especially if periodically visited by tropical storms. A large volume of sediment is derived from such mountains and volcanic highlands at a rapid rate and stored in wide valleys at low elevations. For example, the alluvium of the Ganga valley is mainly derived from sediment eroded by slope failures and running water from the Himalaya. Parts of the ancient Gondwana constitute stable areas of the tropics in every continent. Extensive low-relief features with scattered isolated hills or plateaux are common features of such a landscape. Slopes are either very low in the plains or near vertical as the sides of the plateaux often rise in steep escarpments. The rocks, however, are cratonic with a thin regolith on top. These areas, irrespective of erosional slope processes in operation, do not produce much sediment. Slopes, processes and sediment supply therefore vary across the tropics, and it is difficult to generalise about tropical slopes. It is necessary to qualify first the regional geological and climatic backgrounds. Many of the succeeding chapters provide examples of different types of tropical slope, and highlight the variability in tropical landforms due to geological history, climate and anthropogenic alterations. The quote at the beginning of this chapter highlights the controversy that once reigned regarding the origin of tropical landforms.

Questions 1. A number of geomorphologists believe that slump is the primary type of slope failure in the humid tropics because of the thick mantle of weathered material that covers the hard rocks. Do you agree? Justify your viewpoint. 2. How would you expect hillslopes to look in a) an area frequently visited by tropical cyclones? b) a granitic area in the humid tropics before and after deforestation? 3. What is your expectation of the sediment load of a river that (a) drains a steep, tectonically active mountain area? (b) drains a forested area of low granitic hills in the humid tropics? 4. Chapter 6 includes a list of slope properties that induce failures. List the properties which stabilise slopes. Give examples in support of your chosen properties. 5. Why is the amount of moisture retained by the surfical material so important for slope failures? 6. Fig. 6.3 is from Gabet et al. (2004) where the data are from a drainage basin in the Himalaya Mountains in Nepal. Does it demonstrate that landslides are more frequent towards the end of the wet monsoon?

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Slopes:€forms and processes

7. Pitts (1992) plotted a similar diagram for western Singapore. Antecedent rainfall was plotted against trigger rainfall. It was possible to identify the threshold beyond which landslides occur, but the data showed a wide scatter for smaller and fewer events that had occurred in urbanised western Singapore following a limited amount of rainfall. Why the scatter? 8. It has been said that most of the sediment of major streams is derived from headwater mountains. Is this solely related to failures in steep slopes? 9. How would you determine whether the surface material under a tropical rain forest is creeping downslope? Design a methodology for demonstrating this phenomenon and measuring the rate of creep. 10.╇Larsen and Torres Sánchez (1992) reported on slope failures following the arrival of Hurricane Hugo over eastern Puerto Rico in 1989. There were many failures, but the biggest one happened three days after the hurricane. Why the delay? 11.╇Do you agree with the summary statement that creep, planar slides and debris flows seem to be the dominant processes on tropical slopes? Justify your decision.

7

Rivers in the tropics

The moisture-laden wind blows at speed Awakening the mad river Rabindranath Tagore, translated from Bengali

7.1╇ Components of a river system Rivers are conduits for transferring water and sediment from the land to the ocean. Among the multitude of factors that control their dimensions, appearance and behaviour, four are of high importance:€water, sediment, sediment texture (grain size) and channel slope. If any of these changes, the river can be expected to accommodate the change with a proportional shift in its form and function. To illustrate, other factors being equal an increase in the supply of sediment would lead to deposition in the channel. However, if the increased sediment simultaneously becomes finer, the river may be able to transfer most of the fine sediment downstream. In contrast, an increase in stream flow due to rainfall from a large storm or tectonic tilting of the channel slope may lead to erosion. This relationship (Fig.€7.1) was expressed in a simple diagram by W. M. Borlaug based on an equation derived by E. W. Lane (1955). We need to explore these factors in order to understand rivers.

7.2╇ Water in river channels The basic measure of water in rivers is discharge. Discharge at a point on the river is defined as the volume of water passing that point in unit time. Discharge can therefore be computed by multiplying the cross-section of the channel (also known as channel area) with the velocity of the stream. It is written as

Q = A.v = w.d.v

(7.1)

where Q = discharge A = channel area v = velocity of the flowing water w = width of the surface of the channel d = mean depth of the channel.

101

Channel width and depth are both measured in m, velocity in ms−1, and therefore the Â�discharge is in m3s−1.

Rivers in the tropics

102



Fig. 7.1

The river balance. Generalised from the original diagram proposed by W. M. Borlaug and E. W. Lane. From Gupta, 2011. By permission of Royal Academy for Overseas Sciences, Brussels Velocity in a river is dependent directly on the component of gravity that pulls a volume of water downslope, and indirectly on the friction that has to be overcome before water can flow in the channel. The friction (also known as the resistance) is the summation of several characteristics of the channel:€unevenness of the surface of the bed and banks of the river, the bends in the channel, channel vegetation, nature of the sediment, water temperature, fish in the river, etc. It is a basket term that includes everything that retards the flow of water. It written as n (called Manning’s n) or ff (called friction factor). Various general equations of river velocity have been derived from these two forces (gravity and friction).

Manning’s equation:€v = (R2/3s1/2)/n = (d2/3s1/2)/ n Chézy equation:€v = C√(ds) d’Arcy–Weisbach equation:€v = √(8gds/ff)

(7.2) (7.3) (7.4)

where v, d and n have the same attributes as before and R = hydraulic radius of the channel defined as wetted perimeter/surface width, which can be approximated for wide channels by mean channel depth (this approximation works for most rivers) s = slope C = a constant that represents resistance to flow and can be derived as √(8g)/ff ff = friction factor used in the d’Arcy–Weisbach equation g = acceleration due to gravity. Both the Chézy and d’Arcy–Weisbach equations are dimensionless and theoretically preferable to Manning’s, which is not. But Manning’s equation remains popular, as it is easy to compute and physically perceive. It is so much in use that a manual exists for estimating Manning’s n of a river by comparing it with photographs of a number of streams with

Water in river channels

103

Table 7.1╇ Manning’s roughness coefficients (n) for various types of channels Channel type

Manning’s n

Canal with smooth concrete Canal with ordinary concrete Straight unlined earth canals Rivers in fair condition, some vegetation Winding natural streams Mountain streams with rocky beds and bank vegetation

0.012 0.013 0.020 0.025 0.035 0.040–0.050

Values of Manning’s n from Chow 1964

known values of n (Barnes, 1987). Tables are also available which provide a physical description of the river with a value of n (Table 7.1). We see from the velocity equations that stream velocity is related directly to channel depth and slope and indirectly to all types of friction summed together. The fastest velocity in a river occurs slightly below the surface due to friction with air which slightly slows down the flow at the surface. Velocity is low near the channel bottom as expected. The velocity that occurs at 0.6 depth of the channel is acceptable as the average vertical velocity but the mean of the velocities at 0.2 and 0.8 depths gives a better approximation. Cross-channel velocity is higher over deeper parts, and isovels (lines of equal velocity) in the channel tend to follow the channel shape (Fig. 7.2). The velocity of a river increases downstream. Its slope decreases downstream but its depth increases, and depth carries a higher power than slope in the Manning’s equation. Rivers tend to flow over sand, silt or clay in their lower courses and thus resistance to flow is also reduced. Flowing water exerts force on the bed and banks of a channel. This is the force that erodes. The force per unit area of the channel is called the boundary shear stress and can be written as

τ = γds

(7.5)

where τ = boundary shear stress, determined in Newtons per square m (Nm−2) γ = specific weight of water (9807 Nm−3 for clear water) d = depth (or the hydraulic radius) of the channel in m s = slope of the channel. The form of the equation indicates that rivers flowing with high velocity will exert greater stress on the bed and bank of the channel and hence erode. A river in flood is erosive. Erosion or deposition in a stream is also determined by its unit stream power or the power that a stream exerts over a unit area of the channel (Bagnold, 1966). This is measured in watts per square metre (Wm−2), and expressed as

ω = γQs/w = γwdvs/w = τv

where ω = unit stream power and the other symbols are as before.

(7.6)

Rivers in the tropics

104

(a)

(b)

(c)



Fig. 7.2

Idealised cross-channel distribution of velocity in a river:€(a) location of approximate mean velocity; (b) vertical velocity distribution; (c) cross-channel velocity distribution with shaded area indicating location of fast flow in channel

7.3╇ Sediment in river channels Sediment is commonly perceived as the load that the water of the river carries, which can be divided into two major types:€solution and solid (also known as particulate) load. The solid load is further differentiated by its mode of conveyance. If the immersed weight of a particle is carried by the water, it is classified as suspended load. If the weight is carried mostly by the solid bed of the river, it is bed load. Commonly, the solution load of a river is only a very small fraction of its total load. However, the solution load of the river may be significant if the upstream drainage basin includes a large area of rocks that are easily dissolved, such as evaporites or carbonates.

105

Sediment in river channels

The major chemical elements transported in solution are calcium, silica, iron and aluminium. Stallard (1985) has demonstrated that advanced chemical weathering tends to increase the solution load in parts of the Amazon and Orinoco Basins due to the presence of elements from highly weathered rocks in their solution load. The total volume of solution load increases with discharge, although the increase in discharge also tends to dilute the concentration at the same time unless a very large amount of soluble material is available to the river. Thus, a higher concentration is usually found in low flows. It is obvious from the equations discussed earlier that the power of a river and the stress it imposes on unit areas of the channel bed and banks are related to its velocity. As the velocity rises, the river tends to remove material from the channel perimeter and transport it. This is erosion by rivers. Very coarse material such as cobbles and boulders are always carried as bed load, except in extreme floods. Such catastrophic floods occurred in glacial lake bursts towards the end of the Pleistocene (Baker, 1981, 2007). Very large floods, such as those caused by tropical cyclones, have been known to suspend pebbles and cobbles for short distances. But this happens rarely, and pebbles and cobbles, just like boulders, are usually carried as bed load. Sand usually travels as bed load, but in high floods it is frequently carried suspended by eddies in a river. Silt and clay are transported as suspended load, unless they are in aggregates. The removal (erosion) from the channel of individual particles (clay to boulders) is a function of stream velocity. The critical erosion velocity of a particle is the velocity at which it starts to move. Grains of clay and fine silt, although of finer texture, tend to stick to each other in clumps and therefore require higher velocity for erosion than expected. Sand is removed at a relatively lower velocity than silt or clay, as each grain is independently subjected to stress. Once sediment grains are entrained, they can be transported with a velocity lower than the one for entrainment. A sediment grain being carried either on the bed or in water comes to rest (a phenomenon known as deposition) when the stream velocity falls below the level needed to carry it. This is the fall velocity, which is related solely to grain size. Hjulstrom mapped both velocities against grain size (Fig. 7.3) to explain the work of a river at a particular velocity. According to the diagram, sand is the easiest to erode from the perimeter of the channel. Anything coarser (pebbles, cobbles, boulders) requires a higher critical erosion velocity and, given their size and weight, they are only transported for a short time, as their fall velocity is also high. Sand is carried longer. The critical erosion velocity of silt and clay is higher than sand, but once suspended in water, they are transported for a long time as their fall velocity is very low. Thus river transport leads to a separation of material downstream, finer material travelling longer. This grouping of particles of different size over distance is called sorting. The channel of a river flowing from mountains to the sea therefore contains pebbles, cobbles and boulders near the mountains and silt and clay near the sea. Sand is ubiquitous, and is found just about everywhere although the modal size of the sand grains decreases downstream. Certain generalisations can be reached regarding the nature of the suspended load in a stream: • The suspended load increases directly with discharge. • The vertical distribution of the suspended load in a stream usually but not always parallels the vertical velocity distribution.



106

Rivers in the tropics

Fig. 7.3

Hjulstrom diagram of velocity required to remove and transport sediment load of a given texture. The velocity required to move a sediment grain varies and hence a band of erosion velocity has been shown. Sand (0.06–2 mm) is eroded at a relatively lower velocity than either clay or silt (< 0.0.6 mm) or coarser material (pebbles (2–64 mm), cobbles (64–256 mm) and boulders (> 256 mm). The line marked ‘fall velocity’ indicates the velocity required to keep a particular grain of sediment moving. Erosion and fall velocities together indicate that material coarser than sand is removed only at high velocity (floods) and deposited soon. Sand is picked up first and transported over a wide range of river velocities. Silt and clay are erodible at a velocity higher than sand but once in motion they are carried for a long distance. This leads to sorting of sediment in the downstream direction and also determines that rivers that flood will have a diagnostic appearance as discussed in the text. After Hjulström, 1939 • The peak concentration of the suspended load arrives before the peak in the hydrograph, unless the supply of sediment is abundant • The suspended load may exhibit hysteresis; for the same discharge, the rising limb of the hydrograph may carry more sediment than the falling limb. All these conclusions are diagrammatically illustrated in Figure 7.4. The amount of suspended load being carried can either be computed or directly measured from samples of river water. The numerical figure is called sediment discharge and is determined as milligrams per litre or tonnes per day. Bed load is difficult to measure and is often computed from theoretical bed load functions proposed by Shields, DuBoys, Meyer-Peter and Muller, Einstein and others. Data on the bed load of streams is therefore not as easily available as suspended load. However, the bed load generally constitutes only a small fraction of the total sediment load of most rivers and its amount may fall within the error term for suspended load measurement. It is therefore not always measured. Bed load, however, is important for certain rivers like the Brahmaputra, which are special cases. Material on the channel bed moves in three ways: • rounded or subrounded grains can be rolled • any shape of grain can be dragged by slipping and sliding along the river bed

Sediment in river channels

107



Fig. 7.4

(a)

(b)

(c)

(d)

Idealised pattern of suspended load in a stream channel; Q = discharge of water, Qs = discharge of suspended load

• grains may saltate, i.e. move in progressively decreasing hops after being hit by another moving grain. Two terms are commonly used to describe the transporting ability of a river. The competence of a stream at a given flow is the size of the largest grain that can be moved as bed load. The capacity of a stream at a given flow is the total sediment load. The ability of a stream to erode and transport sediment is related to its bed shear stress (τ) or unit stream power (ω). Bagnold (1966) has shown that stream power determines the capacity of a stream to transport sediment. The higher the values of shear stress and power, the more capable is a stream of carrying out erosion and sediment transport. Extremely high erosion is therefore carried out in floods or steep river gorges, especially where high values of depth and slope are periodically attained given that the drainage area upstream of the gorge is capable of producing large flood discharges (Baker and Costa, 1987). For example, Williams (1983), after examining a large database, stated that within his data set at least 1063 Wm−2 unit power, 225 Nm−2 shear stress and 2.52 ms−1 mean velocity were needed to move a 1.5 m boulder. Such figures are exceeded in many gorges and steep floodprone reaches. Coarse material is therefore transported through the gorges and steep reaches only to be deposited in flatter alluvial basins beyond. There are instances of floods with shear stresses in the region of 103 Nm−2, as happens in hilly areas during tropical cyclones, which can move boulders a few metres in diameter as bed load (Fig. 7.5).

108



Fig. 7.5.

Rivers in the tropics

Buff Bay River, eastern Jamaica, draining the northern slopes of the Blue Mountains. The coarse bed load material indicates the high competence of this river in flood. The very large rounded boulder is 5 m in diameter. From Gupta, 2000, with permission from IAHS. Photograph: A. Gupta

7.4╇ Channel geometry The dimensions of a stream channel are determined mainly by a number of background factors:€discharge, amount of sediment, texture of sediment and channel slope. These factors are very effective when the river is entirely in alluvium. If the river is cut into rock or flowing over a thin layer of alluvium on rock, the underlying lithology and geological structure also become a determining factor for the geometry of the stream channel. An alluvial river tends to be sensitive to changes in water and sediment and adjusts to changes easily. Richards (1982) has called them self-formed rivers. Alluvial rivers therefore tend to change their channel geometry frequently. On the other hand, a river flowing through bedrock requires a high amount of energy to erode the bed and adjust its morphology. Rivers on rock therefore only change their channel geometry under conditions of high energy, e.g. during a large flood or an increase in slope due to a regional tectonic activity. Such changes are separated by long time intervals. An extreme example is the Upper Indus, which has stayed pinned into a tectonics-driven suture zone. A river flowing through a rock canyon cannot change its course within the gorge without catastrophic events taking place. Temporary changes may, however, happen to the alluvial features deposited within the rock-bound channel (Chapter 8). It should be stressed that rivers in alluvium and rivers on rock tend to appear and behave differently and models or statistically generated parameters cannot be transformed directly from one type to another. The situation becomes complicated when rivers have alluvial banks but their bed is on rock or a thin layer of alluvium on rock. A long river usually flows over rock in its upper course in the mountains and in an alluvial valley downstream. Certain structure-controlled rivers, such as the Narmada, Irrawaddy and Mekong, flow

Channel geometry

109

alternatively in rock gorges and alluvial basins, thereby alternating their form along their course. Channel geometry for alluvial rivers is better known. In this section we review channel geometry for both types but with more details for alluvial rivers. The geometry of a river channel has four major characteristics: • channel slope • channel size • channel shape • channel pattern. These four are interrelated and the river is the sum of all four characteristics. It is, however, easier to examine each characteristic separately and then assemble the complete picture.

7.4.1╇ Slope of the river The slope (also known as the gradient) of a river is determined by plotting the elevation of its bed against distance from its source to its mouth (Fig. 7.6). The plotted line is called the longitudinal profile (long profile). As the headwaters of most rivers start in mountains or uplands, the upper part of a longitudinal profile is steep, giving rise to a concave-to-Â�Â�thesky appearance. The degree of concavity is important as the gradient of a river controls the velocity and power of a river to a large extent (see equations 7.1–7.6). When plotted on semi-logarithmic paper, long profiles become straight lines or a combination of straight lines. The standard form of the equation is

h = alb,

(7.7)

where h = elevation l = distance from the source of the stream and a and b are numerical coefficients for individual rivers. Rivers like the Ganga or Brahmaputra tend to erode and pick up sediment as they descend the steep southern slopes of the Himalaya Mountains. The sediment is deposited in the channel or overbank in the lower, flatter part of the long profile. Similarly, 90 per cent of the sediment of the Amazon is derived from the Andes (Meade, 2007). The gradient of a river tends to lessen progressively along its length. The term 'base level' is used to indicate the level of the water surface at the mouth of a river. The base level for a stream could therefore be the level of the sea or of a lake or of a bigger river at the confluence, and this level is recognised as the limit to which the river can deepen its channel. Mackin (1948) visualised an equilibrium condition in rivers where the gradient at each point is just right for passing a quantum of sediment from upstream to downstream without any significant erosion or deposition happening. This notional longitudinal profile is known as the profile of equilibrium or a graded profile. A graded profile is not necessarily smooth but can have sections that vary in slope in order to adjust to local rock types or changes in water and sediment delivered to the stream. In reality, the profile of a river at grade tends to fluctuate insignificantly on both sides of the expected profile. The ultimate profile is a function of a number of variables:

110



Fig. 7.6

Rivers in the tropics

Generalised long profile of the Nile from the White Nile headwaters to the sea. Note variations in slope. From Woodward et al., 2007. From Said, 1994, by permission of Wiley. • regional lithology • discharge • amount and texture of channel sediment • resistance to flow • stream velocity • width of the channel • depth of the channel • regional slope. The major channel properties (profile, size, shape and pattern) are mutually dependent. The design of alluvial canals in Egypt and India reflects this behaviour. After the initial scour and fill and finding of the appropriate width, depth and gradient, these canals become stable and are said to be in regime. The concept is comparable to that of graded rivers. A river maintains its channel dimension and form for a long time. It may undergo minor short-term fluctuations but usually returns to its long-term characteristics. A major disturbance in the drainage system is required to significantly change a graded river. It has been shown by Langbein and Leopold (1964) that channel forms adjust until the energy used per unit bed area for the river to flow is uniform along the channel. This indicates that power expenditure per unit bed area is constant (Richards, 1982). As the discharge increases downstream, it becomes necessary for the slope to decrease and the river develops the concave-to-the sky longitudinal profile. This implies that a river needs to adjust to the material lying on its bed. A river flowing across several rock types may display a composite profile formed of segments at varying slopes on different lithologies due to both differences in the erodibility of different rocks and the variation in size of the resultant bed load. Each segment will have a separate curve with different slopes and concavity. Coarser material will cause the gradient to steepen. A graded profile need not be a smooth profile but may have several protuberances, as demonstrated by Wolman (1955) for the Brandywine Creek, PA, USA.

111



Fig. 7.7

Channel geometry

The Narmada at Dardi, central India. Note the waterfall at the knickpoint in the background with a gorge with rock shoulders downstream of the falls. Imbricated boulders 2 m in size indicate the competence of the Narmada in flood. From Gupta et al., 1999. By permission of Wiley. Photograph: A. Gupta The concave profile is a generality that is not found everywhere. Rivers in arid environments may have a straighter or a convex profile, as discharges may not reach the mouth of the stream, being lost by evaporation and seepage through channel sediment. A concave profile is interrupted in places where rivers cross harder rocks or where uplift has taken place. Here the gradient is steep, so the river tends to flow over it with higher velocity and power. A break in the profile is called a knickpoint. Knickpoints frequently, though not always, give rise to waterfalls or cataracts (Fig. 7.7). The high energy of the river at a steep knickpoint leads to erosion and the knickpoint tends to retreat slowly and finally disappear. A retreating knickpoint may leave a gorge section downstream that starts from the original location of the knickpoint. Large rivers like the Mekong have deep scour pools in rock-bound sections, the base of which could be at elevations below the sea level, even a long distance from the sea (Fig. 7.8). More than 90 deep troughs were found eroded on the bed of the Changjiang (Yangtze), during the construction of the Three Gorges Dam, when the river was diverted for dam construction and its former bed exposed. The locations of such troughs were associated with tectonic zones of weakness and the deepest of these were about 80 m deep (Yang et al., 2001). The bed of a river can be either smooth or show significant relief. The longitudinal profile should not be expected to be smooth everywhere.

7.4.2╇ Size of the river channel Intuitively it seems that the cross-section of a river (also referred to as channel size or channel capacity) depends on the discharge it carries when full. This discharge is referred to as

112



Fig. 7.8

Rivers in the tropics

The Mekong eroding into quartzitic rocks along the border of Thailand with Lao PDR, upstream of Pakse, Lao, PDR. Scale indicated by trees at the bottom left. IKONOS satellite image © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2003), reproduced with permission. See also colour plate section the channel-forming discharge or dominant discharge. A channel-forming discharge should have the following properties. 1. It is capable of maintaining the channel at the existing cross-section without significant erosion or deposition. 2. It is not exceeded frequently enough to deposit sediment overbank to elevate the floodplain. Bankfull, discharge is the discharge that fills the channel when any addition of water will cause the river to flow overbank. As the two banks of a river are usually not at the same elevation, bankfull discharge is identified from the width–depth ratio of the channel. If this ratio at different heights above the bed is plotted on a graph, the lowest width–depth ratio occurs at the bankfull depth. It has been statistically shown that width–depth ratio occurs at an interval of approximately 1.5 years on the annual maximum series. The annual Â�maximum series is a listing of the highest discharge of the year for a number of years. Therefore a discharge of this recurrence interval may be taken to represent the bankfull discharge. Studies in the United States and Western Europe have shown that the bankfull discharge is statistically related to the channel dimensions. Measures of channel dimensions such as channel width, channel area and bend amplitude for various rivers show this relationship when plotted against their bankfull discharges (Leopold et al., 1964). This statistical relationship, however, does not hold without caveats for certain types of river. These are rivers deeply incised in rocks or alluvium, rivers periodically affected by high-magnitude floods or rivers in a strongly seasonal climate. Pickup and Rieger (1979)

113

Channel geometry

have shown that the size of many tropical rivers may depend on more than one discharge where floods and seasonality are common in both the arid and humid tropics. They attribute the task of channel forming to a set of discharges, termed by them as nested discharges, all of which to some extent control the channel size. This characteristic of tropical rivers is discussed in detail in section 7.6.

7.4.3╇ Shape of river channels Two rivers of the same cross-sectional area may have different shapes:€one deep and narrow, the other shallow and wide. Channel shape is calculated as the dimensionless width– depth ratio (also known as form ratio), which is the ratio between the water surface width and the corresponding mean depth of the river. This ratio changes as the water level rises or falls in the channel, reaching its minimum value at the top of the banks. The rate of change depends on the slope of the banks; a channel with near-vertical banks shows a different rate of change than the case where the banks (one or both) are sloping. The shape of a river channel at a given cross-section is a function of three variables: 1. bed and bank material 2. volume and texture of the sediment being transported by the river 3. river discharge and its variations. Channels with cohesive banks in silt and clay or fine sand tend to be steep. Rivers with a high proportion of coarse material in their banks tend to have sloping banks (Schumm, 1960). Finer material encourages vegetative growth in the channel and banks, which further stabilises them. Absence of vegetation on coarse bank material and lack of cohesiveness allow the active erosion of banks, widening the channel. Rivers that carry a high proportion of coarse bed load tend to have high form ratios but rivers with a high suspended load (usually fine sand, silt and clay) have relatively low ones (Lacey, 1930). Rivers with marked seasonality in discharge and those that flood at short intervals tend to have high form ratios (section 7.6). All these background factors are interrelated. A river that floods will bring in coarse material and a large bed load, thereby maintaining a wide channel. Leopold et al (1964) listed four generalisations regarding channel shape. 1. Most rivers have approximately parabolic cross-sections. 2. River channels become asymmetric at bends. 3. The width–depth ratio of river channels increases in the downstream direction. 4. Large rivers have high width–depth ratios. The shape of the channel is often referred to as the form of the channel. It is an important measure because it relates to two other properties of the river:€the channel pattern and the texture of channel sediment. The shape of each channel changes (1) as it is gradually filled with water or (2) in the downstream direction. Leopold and Maddock (1953) described a technique for measuring this change, called hydraulic geometry. It is determined by plotting width (w), mean depth (y) and velocity (v) of a channel against corresponding discharge values on logarithmic graph papers. The plotting can be done for two different sets of data. In the first set, all

Rivers in the tropics

114

b + f + m = 0.28 + 0.42 +0.30 = 1.00



Fig. 7.9

A classic example of at-a-station hydraulic geometry determination, Powder River at Locate, Montana, USA. From Leopold and Maddock, 1953. Courtesy of USGS measurements are from the same cross-section on the river. It is at-a-station hydraulic geometry (Fig. 7.9). In the second case, measurements are taken at various points along a river for the same type of flow such as half-bankfull. This is downstream hydraulic geometry. The relationships of width, mean depth and velocity against discharge can be expressed as



w = aQb y = cQf v = kQm,

(7.8) (7.9) (7.10)

where Q, w, y and v are the same as before and b, f and m and a, c, and k are numerical exponents. It can be shown that b + f + m = 1 and a.c.f. = 1. This allows a comparison between rivers as the relative values of b, f and m will vary although they would sum approximately to 1. A wide channel with banks in coarse material will have a high b and low f, as it gets relatively wider with increasing discharge. A narrow river with sloping banks will show the opposite. A

Channel geometry

115

Table 7.2╇ Values of at-a-station hydraulic geometry



Fig. 7.10

River and location

b

f

m

Source

Yallahs at Llandewey, Jamaica Godavari at Shagad, India

0.51 0.32

0.29 0.33

0.23 0.36

Godavari at Nanded, India

0.24

0.63

0.12

Tapi at Burhanpur, India Sina at Kolegaon, India

0.24 0.19

0.51 0.47

0.26 0.35

Bhima at Takali, India

0.13

0.43

0.44

Buff Bay at Balcarres, Jamaica Narmada at Rajghat, India

0.07 0.04

0.47 0.50

0.51 0.46

Gupta (1975) Deodhar and Kale (1999) Deodhar and Kale (1999) Kale et al. (1994) Deodhar and Kale (1999) Deodhar and Kale (1999) Gupta (1975) Kale et al. (1994)

Idealised sketch of four channel patterns selection of at-a-station hydraulic geometry values are shown in Table 7.2. Hydraulic geometry implies that channel dimensions are adjusted to flow conditions. This is comparable to the concept of stable alluvial channels designed by engineers in India mentioned earlier. It is an excellent non-subjective way of describing the shape of a river channel.

7.4.4╇ Channel pattern Rivers not only adjust their cross-sections: they also adapt a planform that is suited to their discharge, sediment and regional slope. The planform or channel pattern has been defined as the configuration of a single river as seen from the sky. Rivers are commonly classified as one of four types:€straight, meandering, braided and anastomosing or anabranching (Fig.€7.10). These classes can be subdivided, based on channel measurements. The majority of rivers are either meandering or braided. Many rivers display the same planform throughout their length, although exceptions do occur. For example, a braided river may

116

Rivers in the tropics

change to a meandering pattern with altered gradient and sediment load, or a seasonal river which braids during the dry season may take on a meandering appearance with the rising of the water level in the wet season. All natural rivers have an inherent tendency to follow a sinuous flow. The sinuosity of a river channel is calculated as a ratio called the Sinuosity Index. Sinuosity Index = Distance between two points along the channel/distance between the same two points along the valley.

Straight rivers are rivers whose channels run straight for several times the channel width or for which the computed sinuosity is very close to or exactly 1. A straight flow between two points obviously gives the channel the highest possible gradient, and hence these rivers flow on slopes steeper than that of other channel types in the region. The straightness in many cases is derived from geological controls such as faults, which determine the gradient of the channel and its ability to erode actively. For example, the Irrawaddy in upper Myanmar changes its planform from a meandering pattern through a wide alluvial valley to a deep straight gorge on reaching the Sagaing Fault (Fig. 7. 11). The banks of a straight channel may run straight but the thalweg (the line joining the deepest part of the channel) may not. It may follow a winding path between alternate bars on either side (Fig. 7.10). Pools (deep reaches with slow flow) alternate with riffles (shallow reaches with fast flow) along a straight channel. Fine sediment tends to accumulate in pools, coarse material in riffles. Measurements along straight channels have shown that pools and riffles tend to be spaced at regular intervals, about 5–7 times the mean width of the channel. A meandering river is expected to have a sinuosity index of 1.5 or above, but more importantly it should have a wandering channel with the thalweg (the deepest part of the channel) regularly winding from one bank to another, approaching different banks alternately (Fig. 7.12). Each bend is asymmetrical in cross-section with a deep pool on one side and a crescent-shaped bar (point bar) on the other. Halfway between two successive bends the channel is symmetrical with the thalweg in midchannel. A meandering channel tends to migrate at bends, eroding the meander cliff side but extending the point bar by deposition. Thus its original dimensions are maintained. A meandering stream thereby wanders across its valley floor. Brice (1983) recognised two different types of meandering channel. The first type carries approximately equal widths for each cross-section and is considered passive, i.e. does not move much across its valley floor. The second type is mobile and its cross-sections are wider at bends with well-developed point bars on the concave side. Schumm (1960) has shown that a significant proportion of clay and silt in the bed material is associated with high sinuosity of the channel. Streams with coarse bed load, however, also meander. Schumm (1968) suggested that the morphology of a meandering channel depends on the type of load − suspended, mixed or bed load − that it carries. The dimensions of a meandering river (Fig. 7.12)€– the wavelength of the meanders, channel width and the meander radius€– are statistically related to each other and to its bankfull discharge (Leopold and Wolman, 1957). Braided rivers display a number of sub-channels between the two banks of the entire river with mid-channel bars separating such channels (Fig. 7.10). These channels are characterised



117

Channel geometry

Fig. 7.11

Diagrammatic sketch of the Irrawaddy River changing its course from a free meandering river in an alluvial floodplain to a straight restricted gorge section as it enters the Sagaing Fault in north Myanmar. Sketched from a SPOT image on the web and checked with topographical maps. From Gupta, 2005b. By permission of Oxford University Press by their large width, high width–depth ratio and the frequently changing position of the smaller channels and mid-channel bars. Leopold and Wolman (1957) concluded from field data that for a given discharge, braided channels have steeper slopes than meandering rivers, an observation that has been supported by the experimental data of Schumm and Khan (1972). Braiding is associated with a high coarse sediment load that is transported and deposited repeatedly, following fluctuations of discharge. They are usually associated with steep slopes, abundant supply of coarse material and a fluctuating discharge. In such rivers, the pattern stays the same but the locations of channels and bars shift following individual high flows. It is therefore not surprising that braided channels have been reported from • semi-arid regions of low relief that receive drainage from mountainous areas • areas with seasonal rainfall • highland areas in a variety of climates

Rivers in the tropics

118



Fig. 7.12

The geometry of a meandering river • breaks of slope near the foot of a mountain • glacial outwash areas • periglacial areas, especially over permafrost. In all these areas an abundance of coarse bed load and a fluctuating flow regime are found. Anastomosing (also known as anabranching) rivers (Fig. 7.10) display a multichannel pattern where, unlike braided river channels, sub-channels and bars/islands are fixed in position. The stable islands of an anabranching river are usually in clay or silt, reflecting the load carried by an anabranching river. Such islands, being stable, are generally under

119

Channel network and nodes

vegetation. The individual channels of an anabranching system could be straight, meandering or braided (Schumm, 1985). The channels are not only stationary; they are also deep. Anabranching rivers are usually found in intermontane valleys, on low gradient fans, and in lacustrine and alluvial plains, where floods cause channels to avulse through erosionresistance banks. Nothing striking happens between floods. A similar pattern is sometimes seen in rock as described for the Sabie in South Africa by Niekerk et al. (1999) and the Mekong by Gupta and Liew (2007). The four characteristics of a river channel are interrelated and determined by the combination of channel properties such as discharge, sediment, bed and bank material. For example, braided rivers are associated with a high width–depth ratio, fluctuating discharge, coarse and abundant sediment, and steep slopes. Variations do occur from such generalisations but only as exceptions. Schumm (2005) has summed up the discussion by dividing rivers into two basic classes:€regime and non-regime. Rivers in regime are those described above which tend to be stable and demonstrate the expected characteristics. Non-regime rivers are either unstable or controlled by geological factors such as bedrock, the effect of which surpasses all other factors.

7.5╇ Channel network and nodes Individual channels join up to form a network that transfers water and sediment from an area of land surface. This area is the drainage basin which is named after the main channel in the network. The density and pattern of the network in the basin reflect various factors such as geology, relief, hydrology and land use in the basin. In a classical paper written almost 70 years ago, R. E. Horton explained certain rules behind the arrangement of this network (Horton, 1945). Horton proposed an ordering system for the network which has been modified several times since then. Figure 7.13 demonstrates Horton’s original system and two major proposed modifications. Horton also demonstrated that the meeting of the tributaries (nodes) and aspects of their dimensions follows certain relationships. Such relationships could be demonstrated using any of these systems. Horton called the fingertip tributaries the first-order streams. Two first-order streams come together to form a second-order stream, two second-order streams join to form a third-order river and so on (Fig. 7.13a). It is a wonderful system but it requires extending the ordering of the main stream all the way back to its source. This could be problematical as it is not always clear which one is the major stream upstream of a confluence, especially from a map. In the simplified system proposed by Strahler (1952a) both streams of the same order carry the same order up to the confluence (Fig. 7.13b). Shreve (1967) added the order of individual tributaries to the main stream, (Fig. 7.13c) so that the magnitude of the major streams can be adequately expressed. As Horton originally demonstrated, geometric relationships exist between the number of the tributaries, the lengths of the tributaries and the area of the basins drained by tributaries of different orders. These relationships have been clearly discussed by many geomorphologists and will not be reviewed here. For example, excellent reviews have been provided by Strahler (1964) and Abrahams (1984). Primarily the idea of network

Rivers in the tropics

120

(a)



Fig. 7.13

(b)

(c)

The same channel network according to the ordering systems of (a) Horton, (b) Strahler and (c) Shreve and ordering is very helpful in understanding the discharge of water and sediment along a drainage basin. Horton’s rules work well for river basins in the tropics. The comparative efficacies of two joining streams are important and determine the nature of sedimentation and erosion at their confluences. Similarly the geomorphic coupling at the contact between hillslopes and drainage channels to a certain extent determines the nature of the landform (Harvey, 2002). Deposition of material accumulates at confluences and on foot slopes if the main stream is not effective at the time the sediment arrives. In many cases, a time-based accumulation of material is cleared in an episodic large event like a flood (Fig. 6.8). The landscape is a reflection of many factors including elapsed time since the last such clearing (Fig. 7.14). The characteristics of the tropical climate act as a control factor.

7.6╇ River systems of the humid tropics Do tropical rivers look and behave differently to those of temperate regions? Governed by a different hydrologic environment many of them do. The form and function of tropical streams, however, are also determined by other aspects of the drainage basins:€geology, relief and land use. Such factors may locally override hydrology. We will examine the characteristics of streams in the humid tropics in this section. Rivers of the temperate regions demonstrate a relationship between bankfull discharge and channel size and shape (Leopold and Maddock, 1953; Leopold and Wolman, 1957; Leopold et al., 1964). This relationship can be explained by the magnitude–frequency concept of Wolman and Miller (1960). The most effective force on these rivers is bankfull discharge, which is powerful enough to determine the size and shape of the river and frequent enough to maintain their dimensions. A large flood may temporarily enlarge the channel

121



Fig. 7.14

River systems of the humid tropics

Accumulation of coarse sediment at a coupling point, Tiger Leaping Gorge, Jinsha Jiang (Upper Yangtze), Yunnan, China. From Gupta, 2011. By permission of Royal Academy for Academy for Overseas Sciences, Brussels. Photograph: A. Gupta but within months or years it would revert back to its original size. The river returns to its former size with in-channel deposition from common post-flood flows carrying a suspended load of silt and clay and also transporting a substantive amount of sand and small pebbles. This has been repeatedly demonstrated in the field (Costa, 1974; Gupta and Fox, 1974). In contrast, the size of a river channel in arid areas is controlled by floods that occur at long intervals of time (Wolman and Gerson, 1978) and in between the channel is usually dry. The channels of arid areas are essentially flood-controlled (see Chapter 12). The rivers in the seasonal tropics fall between these two cases. Extreme floods commonly occur on rivers whose basins are located between about 10° and 30° of latitude (Gupta, 2000). Such floods usually result from tropical storms occurring in the wet season; a number of these storms are strong enough to be recognised as tropical cyclones. Accounts of the resulting high shear stress and unit stream power, enhanced stream competence, sediment transport and storage, and channel forms are now available for a limited number of streams in Australia, South Asia and the Caribbean, and generalisations are possible. In such floods, high stream power operating over a moderate to high duration (Costa and O’Connor, 1995) are able to erode the bed and banks of the channel and transfer sediment for the time period. Flood effects are preserved if the sediment is coarse and the relief is high, as in such environments flood effects are difficult to undo (Gupta, 1988). The geographical locations where tropical cyclone-driven floods (and smaller storms) could be expected to affect river systems are • North and Northeast Australia • Parts of Southeast Asia affected by tropical cyclones, e.g. Myanmar, Viet Nam and the Philippines • Parts of the Indian subcontinent • Madagascar and neighbouring coastal areas of East Africa



122

Rivers in the tropics

Fig. 7.15

Annual hydrographs of four stations on the Mekong River (Total drainage area 795 000 km2):€Chiang Saen, Thailand (14% of the basin area) Chiang Khan, Thailand (36%), Pakse, Lao PDR (69%) and Stung Treng, Cambodia (80%). Data from Mekong River Commission, Lower Mekong Hydrologic Yearbook, 1997. From Gupta, 2007b. By permission of Wiley • The Caribbean Islands, the coastal areas of the Gulf of Mexico, and tropical and subtropical North America. The effects are especially strong where storm tracks are intercepted by mountain ranges. Taiwan in the subtropics is an excellent example. Strong seasonality is associated with much of the humid and subhumid tropics (see discussion on monsoon rainfall in Chapter 3) irrespective of the annual total rainfall, which may range from below 1000 to above 4000 mm. A very high percentage of the annual rainfall (≥75 per cent) arrives in the wet season which may extend for 4–5 months. The hydrographs of the rivers of the seasonal tropics therefore demonstrate a seasonal pattern (Fig. 7.15). The wet-season discharge is much higher than that of the dry season. Almost all the work of the river (erosion, sediment transport, etc.) is carried out during the wet season. Large tropical storms arrive over these river basins episodically, usually in the wet season when the river is already high and the soil saturated. The effect of seasonality is also reflected in sediment transport in many tropical rivers. In a recent publication, Billi and el Badri Ali (2010) reviewed suspended sediment transfer in the Blue Nile for three years:€1967–69. Although this is a short period, they could conclude that any significant transfer of sediment only happens during the wet monsoon period, July to October. Their curves for suspended sediment plotted against time showed considerable variation even within the wet period. During the three years of record, the maximum and minimum suspended daily sediment discharges at Khartoum where the Blue Nile meets the White Nile was 2 242 588 and 2484 t. A good correlation (R2 = 0.84) exists

123



Fig. 7.16

River systems of the humid tropics

Cross-section of the Auranga showing the 3-step physiography of the channel. From Gupta and Dutt, 1989. (www.borntraeger-cramer.de) between water and sediment discharges but not with sediment concentration. Peaks in sediment concentration appear to precede that in discharge. Most of the sediment is fine sand with an unusual proportion of coarse sand being present. The clay fraction is also high, but unlike fine sand, it tends to decrease with a rise in low discharge. Overall suspended sediment concentration increases with flow velocity. This is a description which probably holds true for many seasonal rivers in the monsoon tropics. Rivers in the seasonal humid tropics use a set of channels of different size, progressively smaller ones nesting inside the next bigger one in order to adjust to floods and seasonality. Working in Australia, Pickup and Rieger proposed that channel forms result from a series of discharges rather than a single dominant one (Pickup and Rieger, 1979). This nested pattern has been reported for other rivers of the seasonal tropics. The Auranga, a fifth-order stream that flows across the northeastern corner of the Indian Peninsula is an excellent example (Fig. 7.16). It is a seasonal stream carrying sand and fine pebbles, dependent on the monsoon system of South Asia. During the dry season, the Auranga is a braided channel in sand, and water flows only through the lowest part of its channel. Well-developed point bars occur above the level of the dry-season braided channel. At some of the bends, even higher bars, termed flood bars, exist, displaying an even wider channel. The dry-season braided channel, point bars and flood bars are all inside the riverbanks marking the entire channel. During the wet season, the river rises, fills part of the channel and meanders around point bars. The discharge starts to fall at the end of the wet season, and the river becomes braided with small sand bars separating channels that are only several centimetres deep. This kind of braiding has been described in detail for the Platte River in the Great Plains of the subhumid western United States (Smith, 1971). When high-magnitude floods associated with tropical storms arrive in certain years, the river needs the entire cross-section that consists of the flood bars,

124

Rivers in the tropics

point bars and low channel to carry floodwaters. The channel is cleaned of sediment during the floods but as the floods recede, flood bars are deposited at some of the bends with the slowing down of the river velocity. Subsequently, new point bars are rebuilt in the old places as determined by channel bends and in the dry season the braided low-flow channel reappears below the level of point bars (Fig. 1.2c). The Auranga therefore flows through a nested channel that is adjusted to the low flows of the dry season, the high flows of the wet season and occasional very high floods (Gupta and Dutt, 1989). A nested set of channel-forming discharges explains the size, shape and sedimentary characteristics of stream channels in the seasonal tropics. These rivers have been reported from different areas:€the Narmada in India (Gupta, et al., 1999), Yallahs in eastern Jamaica (Gupta, 1975), the Burdekin in northeastern Australia (Wohl, 1992), and the Mekong in Southeast Asia (Gupta and Liew, 2007). Both bed shear stress and unit stream power are strikingly different for the low flows of the dry season, the high flows of the wet season and high-magnitude floods. Wohl (1992) examined the effect of large floods and highly variable discharge regime in the seasonal tropics of northern Australia. During the dry season of April to November, the Burdekin carries a discharge of 60 m3s−1 and transports very fine silt and clay. The discharge rises to an average of 1260 m3s−1 during the wet season, and the sand and silt transport increases significantly. The sand comes from the weathering of granitic rocks on the hills and plateaux of the upper catchment. High-magnitude floods with a discharge of 15 000–30 000 m3s−1 are related to certain aspects of the morphology of the rocky gorge of the Burdekin, e.g. location of boulder bars and flood levees. These features form where the shear stress and unit stream power of the Burdekin are reduced due to channel widening and the passing high-magnitude floods tend to deposit some of the sediment. Only such huge floods have the competence to move the coarse material of the boulder bars of the Burdekin. Fairly frequent wet-season floods are capable of transporting a number of boulders found on bars, but extreme floods control the location of boulder bars which lesser floods may only rework slightly. A channel-in-channel topography is common in these rivers and the overall channel shape is box-like with high banks. The width–depth ratio of the channel decreases strikingly with the arrival of high-magnitude floods that use the entire channel. Quantities of very coarse sediment are moved during high-magnitude floods which are capable of dramatic erosion. The Narmada at Dardi, central India (Fig. 7.7) excavates and transports up to 2-metre cuboid blocks of orthoquartzite, stacking them like upstream-imbricated cushions (Gupta et al., 1999). In flood, the river probably clears out all the sediment in the channel except such huge boulders. Subsequent high flows of the wet season rebuild an inner channel by depositing sand and pebbles around boulders left by the high-magnitude floods at appropriate locations in the big channel. This leads to the channel forms, even floodplains, being inset within the bordering high banks that bound the high-magnitude floods. This happens both in rock gorges and alluvial sections of the Narmada; except in rock gorges, the deposited material is seldom smaller than pebbles, whereas in wider alluvial sections sand builds up kilometre-scale point bars with megaripples on top (Fig. 7.17). Clearing out the channel of all but very coarse sediment in hurricane floods and rebuilding depositional features after the passage of the flood peak also occurs in the Caribbean mountain streams (Gupta, 1975).

125



Fig. 7.17

River systems of the humid tropics

Sequence of channel form construction in the Narmada River. From Gupta et al., 1999. By permission of Wiley In the case of the Narmada, which flows inside 10–15 m high banks, no floodplain exists above the channel. A discontinuous and temporary floodplain-like feature, built of silt and clay, is seen inside the banks in places. In the case of large rivers like the Ganga, as described in Chapter 9, a shallow wide braided channel is flanked by riverbanks that are filled during the wet monsoon, with a higher floodplain surface and flood banks beyond. The entire cross-section is submerged during high-magnitude floods. Wolman and Miller (1960), while opting for bankfull discharge, stated that as the threshold stress increases catastrophic events become more important in shaping a channel. Variations from the general picture occur if seasonality in flow is not pronounced or the river carries only fine-textured sediment that can also be moved by flows of high frequency. Local geological characteristics too may override the described pattern by forcing the river to flow through sagged depressions or fault-guided gorges or across hard rock barriers. Bedrock rivers also are different. They are modified only during floods, although as bedrock sections are usually steep, these sections are not associated with significant



126

Rivers in the tropics

Fig. 7.18

Kaveri River, Srirangapatnam, Karnataka, India. Note structural control of the riverbed topography. Photograph A. Gupta sediment deposition, almost all sediment being flushed out downstream. Bedrock reaches may have an inner gorge and shoulders in bare rock (Fig. 7.7) but the size of the bars will be small and they tend to be in material coarser than sand. Sediment may also build up in the channel behind cross-channel barriers in bedrock. Thus the effect of geological structures and tectonics may be locally important, overriding the hydrological component. The river may be fault-guided like the Irrawaddy passing through the Sagaing Gorge (Fig. 7.11), but even inside the channel, smaller faults and joints control channel relief, well illustrated during low flow conditions (Fig. 7.18). Rivers in the seasonal tropics have their own style of adjusting to the magnitude–frequency concept. With the onset of aridity, rivers in basins of the subhumid and arid tropics tend to have longer periods of low flow or no flow. Most of the work is carried out during a brief wet season or the periodic passage of short-duration floods. Not many of these rivers have been studied. In the arid tropics, the form and behaviour of rivers depend solely on the passage of floods of varying magnitude and frequency. An account of such rivers is given in Chapter 12. A number of generalisations can be made regarding the rivers in the seasonal tropics: • A very high percentage of annual rainfall (75 per cent or more) occurs in the wet season, which may extend up to 4–5 months. • Large storms, when they occur, usually arrive in the wet season when the river is already high and the soil saturated. • It is difficult to measure discharge and sediment transport in high-magnitude floods. • Calculated bed shear stress and unit stream power show striking differences among (a) low flows during the dry season, (b) high flows during the wet season and (c) highmagnitude floods. • A large number of rivers in the seasonal tropics tend to braid during the dry season, but not so in the wet period when they may have a meandering appearance. • A channel-in-channel topography is common. • The overall channel shape is box-like with high banks.

127

Questions

• The width–depth ratio of the channel decreases strikingly with the arrival of high-magnitude floods when the entire channel is in use. • Quantities of very coarse sediment can be moved during the high-magnitude floods. • Readjustment of part of the sediment occurs during the high flows of the wet season. • Channel forms are usually inset within the bordering high banks associated with highmagnitude floods. • Variations from the general trend may occur depending on the strength of seasonality and the texture of sediment being transported. This in turn depends on basin location and modification. • Bedrock channels and alluvial channels look different and behave differently. • The effect of geological structures and neotectonics may be locally important. Rivers in the seasonal tropics have their own style of adjusting to the magnitude–frequency concept. Their forms and behaviour should also be remembered in conceptualising and managing the tropical environment.

Questions 1. Compute the discharge of a stream with 100 m width, 1.5 m mean depth and 0.0002 slope, given that n is (a) 0.02, (b) 0.035, and (c) 0.05. What effect does a changing n have on velocity and discharge? 2. What kind of a river would you expect to go with each of the n values? 3. A river flows through a 2 km gorge. After a large flood on the river, the elevation of the flood levels marked on the rock walls was 500 m at the beginning of the gorge and 480 m at the end of it. The average depth of the flood was 8 m in the gorge. Calculate the shear stress and unit stream power for the flood in the gorge. Assume n = 0.03 and use γ = 9807 Nm−3. 4. When you finally came to the gorge, the flood has abated and the water has gone down to an average depth of 1.5 m. What would be the shear stress and unit stream power? You have to guess a new n. 5. When the suspended load in a flood is plotted against discharge as in Fig. 7.3, the rising limb of the graph shows more sediment concentration than the falling limb. Why? Under what conditions could the opposite be found? 6. Bed load is usually not measured for most streams. Why? Explain under what conditions the bed load of a river would contribute a significant proportion to the total load. 7. (a) The river Mekong passes through a rocky gorge upstream of the Lao capital of Vientiane. Calculate the shear stress and unit stream power for a flood of 22 900 m3s−1 in the gorge. γ = 9807 Nm−3, depth = 10 m, slope = 0.002 and width = 335m. 8. What type of sediment can the Mekong move in the gorge? Williams (1983) found that a 1.5 m boulder can be moved under conditions of at least 1063 Wm−2 unit stream power, 255 Nm−2 shear stress and 2.52 ms−1 mean velocity. Use these figures as general indicators.

128

Rivers in the tropics

9. The Mekong is wider and shallower at Vientiane, where the river flows through an alluvial reach. Would the character of the sediment change between the gorge and Vientiane? If no tributary flows into the river between the gorge and Vientiane, and the height of the flood at Vientiane is 3 m and the slope 0.0015, what is the width of the flood? You may need to assume n from a table. Is flooding extensive in the alluvial reach? 10. You are standing in the midstream of a meandering river looking at the cutbanks. The river mostly carries silt, clay and fine sand as suspended load. In flood it also transports large quantities of coarse sand and pebbles. What kind of stratigraphy would you see in the cutbanks? Try the most likely scenario. 11. Most of the sediment in a river comes from its headwaters in the mountains. The valleys in the mountains are also areas where dams are constructed. How would a river adjust itself after a dam is built across its main channel in the mountains? Answer this question before you read Chapter 17. 12. A graded profile of a river is a smooth profile. Is this statement correct? Justify your answer. 13. You have two rivers with the following at-a-station hydraulic geometry measures. (a)╇ b = 0.44, f = 0.29, m = 0.27 (b)╇ b = 0.14, f = 0.45, m = 0.41 What would the two rivers look like? What else can you say about them? 14. Figure 7.11 shows the Irrawaddy in upper Myanmar entering a fault-guided gorge. What transformation does the river undergo at the head of the gorge? What transformation should you expect as the river emerges from the gorge to an alluvial valley again? This transformation is not in the figure. 15. Which channel pattern would have islands with better-established vegetation and why? (a) a braided stream (b) an anabranching stream 16. How are the four characteristics of a river channel (slope, size, shape and channel pattern) interrelated? 17. Find a topographical sheet of your area with contours and at a scale of at least 1:63â•›360. Trace the drainage lines of a river basin from the map to a sheet of tracing paper. Mark the drainage lines according to Strahler. Tabulate the number of streams for each order. Divide the total number of streams of each order by the total number of the next higher order. This is the bifurcation ratio. What range of ratio do you get for the basin? What is the average of all such ratios? 18. (a) In which part of the humid tropics would you expect nested channels? (b) Under what environmental conditions in these areas would nested channels not occur? 19. How do the rivers of seasonal humid tropics adjust to the magnitude–frequency concept of Wolman and Miller (1960)?

8

Alluvial valleys

Of golden sands, and crystal brooks John Donne

8.1╇ Fluvial depositional environment A river tends to deposit sediment in certain locations in a valley. Intuitively we expect sediment to be deposited as stream velocity drops, starting with the coarsest fraction. When that happens, the preferred locations for depositing sediment are • on floors of valleys in the mountains, frequently near confluences or barriers in the channel • at the foot of the mountains or plateaux (alluvial fans) • along an alluvial valley in the plains (channel bars, floodplains and terraces) • deltas at river mouths. The valley floors in the mountains and alluvial fans are discussed later in Chapter 13 and the deltas in Chapter 11. Here we review alluvial deposition in plains. The current alluviation in a river valley happens on top of an earlier sequence, which is mostly bedrock in the mountains and the upper course of the river but could be an older layer of alluvium in lower course. Meade described alluviation in the largest river basin in the world, the Amazon, as follows: As they leave their narrow mountain gorges, the streams coming off the Andes flow into river beds that are self-made. Massive amounts of sediment brought down from the mountains become the substrates over which and through which the flowing waters, with their accompanying load of even more sediment, must make their ways. Once off the Andean slopes and out of the confinement of mountain-girt channels, the riverborne sediment is likely to endure many episodes of confinement and occupy many rest stops and storage compartments before it reaches its ultimate destination. (Meade, 2007:€46)

This is for the largest river in the world, but the pattern remains the same for most rivers when scaled down appropriately. The steep and possibly tectonic upper parts of a river system carry high stream powers, and erosion and transfer of sediment is common (Blum, 2007). The depositional forms in the highlands are fragmentary and sediment grains tend to have a brief residence 129

Alluvial valleys

130



Fig. 8.1

Source-to-sink passage of a river time. The situation changes in the upper plains where the stream gradient is low and valleys are usually described as mixed bedrock alluvial valleys. The alluvial forms such as floodplains become continuous, sediment grains have longer residence time, and the entire valley acts as a conveyor belt of sediment being transported to the sea. Alluvial valleys and channels entirely in alluvium occur further downstream near coastal plains and deltas (Fig. 8.1). We examine the material, forms and sediment transfer in alluvial valleys in the following sections.

8.2╇ The alluvial valley In the upper part of their course in the plains beyond the mountains, rivers flow in a mixed bedrock alluvial valley. Downstream, alluvium in the valley thickens and bedrock is rarely visible. The valley alluvium is a three-dimensional body which fills a linear depression. This depression could be erosional, cut into rock below the surface alluvium, or a tectonic downwarp. It is difficult to visualise the base of the alluvium unless a number of boreholes are sunk or a geophysical survey is carried out which allows mapping of the rock floor (Schumm, 1977). It seems likely, from the limited amount of data available, that (1) the rock base is usually uneven and (2) a series of alluvial deposits rest on the bedrock, not all of which have necessarily been deposited by the current river. Figure 8.2 is a generalised diagram of a valley cross-section. It indicates that the upper level of the alluvium may not reflect all the morphology at the base. The present river may not be located directly over the deepest part of the valley. There may be an older alluvium below the alluvium of the present river. Past hydrology and river geomorphology can be derived from the texture and sedimentary structure of the past alluvium. Its nature and organic or carbonate contents may provide information about the past

The alluvial valley

131



Fig. 8.2

Generalised diagram of a valley cross-section



Fig. 8.3

Rock ribs in the Mekong, downstream of Chiang Saen, northern Thailand. Photograph: A. Gupta climate. Such indicators are known as climate proxies, and they have been heavily used for studying palaeoclimate during the Quaternary (see Chapter 16). The alluvium of the current river lies on top and is reworked from time to time in high flows. The depth to the rock base and its geological structure may still influence the form and function of the present river. This has been demonstrated through the thick alluvium of the Amazon River (Mertes and Dunne, 2007). Several structural highs cross the Amazon valley below the thick alluvium. Where the river flows over such rock arches, the river narrows in spite of the thick alluvium, its gradient steepens, and there is less deposition in the channel. After crossing a structural high, the Amazon widens and sediment is deposited in the channel (Dunne et€al., 1998). Where alluvium is thin or nearly absent, the older layers and the bedrock are visible. This is common for rivers near the mountain front. Given the regional geology, bedrock at shallow depth controls the form of the channel or even outcrops to form rapids and a line of rocks across the channel, known as rock ribs (Fig. 8.3). This is common even in a large river like the Mekong for the first 4000 km of its course (Gupta and Liew, 2007). Alluvial morphological features that fill a river valley can be listed (Table 8.1), although not every feature is present in all valleys.

Alluvial valleys

132

Table 8.1╇ Alluvial valley forms and sediment characteristics Location of deposition

Terms

Characteristics

In channel

Alluvium resting on bed

In channel but higher forms

Point bars Marginal bars Midchannel bars

Overbank

Floodplain

Back of floodplain

Terrace

Valley margins

Alluvial fans and screes

Bed load in storage between transport; period of storage usually directly related to texture Fills in abandoned or aggraded channels Point bars and marginal bars are usually formed by channel shifting or found below bank collapses. Midchannel bars of various size are formed in braided and anabranching streams Formed by river shifting (lateral accretion) or overbank floods (vertical accretion); in braided streams by merging of midchannel bars and islands Abandoned floodplain at a higher level Sloping triangular bodies formed by deposition by headwaters, �tributaries or mass movements

8.3╇ The channel alluvium The bed of a river is commonly floored with coarse alluvium. This alluvium is referred to as bed material, the gravelly (pebbles, cobbles, boulders) part of it being carried as bed load and the finer part (sand) as both bed load and suspended load. If the bed material is coarse, gravel accumulates at intervals forming shallow riffles and rapids. They are separated by deep pools where finer material accumulates. In tropical rivers that experience high-�magnitude floods, such accumulations of coarse gravels probably start by deposition of boulders in places dictated by channel geometry or exposure of bedrocks at the falling stage of the flood hydrograph. Cobbles, pebbles and sand then accumulate round the boulders and rock features to complete the riffles. Sand is deposited in the pools later and finer suspended material is transported downstream. This type of deposition, which requires availability and transportation of very coarse material, is usually seen in the upper river courses near the mountains, downstream of a rocky gorge between two alluvial reaches, and where a tributary contributes coarse material to the main stream. Sand is probably the commonest material deposited in an alluvial channel. The deposited sand may mask the channel base or form bars of different types.

The channel alluvium

133



Box 8.1

Critical, subcritical and supercritical flows The velocity of a surface wave moving over water is expressed as √(gy) where g is acceleration due to gravity and y is the vertical distance to the water surface. At critical flow, water moves at the same velocity as that of a surface wave moving relative to the flowing water. When the velocity of the flow is less than critical, it is considered as subcritical flow. Similarly, supercritical flow occurs when the channel velocity is higher than the critical velocity. The velocity of a river almost always stays subcritical. Imagine a surface wave being created by a piece of rock falling into the river. At subcritical flow, waves from this disturbance would move both upstream and downstream from the point of impact. We are familiar with this scenario. If the flow of the river is critical, the upstream point of the wave would appear to be stationary to an observer standing on the bank who would also see the downstream front moving at twice the velocity of the river. If the river flows very fast and the flow is supercritical, then the entire disturbance from the falling rock will move downstream and no wave front will propagate upstream of the point of impact. This happens rarely, only during very high floods or at waterfalls. An analogue can be drawn with subsonic and supersonic effects with the velocity of sound replacing critical flow. The ratio of stream velocity to wave velocity,√(gy), is the Froude Number (Fr). If Fr < 1, flow is subcritical; if Fr = 1, flow is critical; and if Fr > 1, flow is supercritical. A river is almost always at a subcritical stage when sandy ripples and dunes are created on the bed of the channel. Plane beds are formed when the velocity of the river approaches the critical mark. Antidunes form rarely, only at the supercritical stage. Sand is easily eroded and transferred (Chapter 7), and the surface features and the internal structure of a sand body deposited in the channel reflect the nature of the depositing flow. Sediment accumulates on the channel bed as discrete grains or aggregates. The coarser variety (sand-boulder) behaves as discrete grains. Sand and pebbles may form bars or a layer of sediment on the channel bed. Pebbles tend to imbricate against each other with individual pebbles sloping upstream. In extreme floods, as described for the Narmada in Chapter 7, such a characteristic is extended even to metre-scale boulders (Fig. 7.7). Pebbles and coarser material may lie buried under sand. Aggregates of sand and silt create bedforms under flowing water. Such bedforms appear as asymmetrical or symmetrical waves and vary in size from ripples (in centimetres) to large dunes (tens of metres). Materials coarser than sand are usually found in bars or forming riffles as described in Chapter 7. Three factors determine the types and size of bedforms: • flow velocity • flow viscosity • size distribution of bed material. Bedforms of different types depend on the velocity of flow at which they are built. They are grouped into lower and upper regime forms and are associated with critical and supercritical flows (Box 8.1). Each type is associated with a characteristic mode

Alluvial valleys

134

Table 8.2╇ Types of bedform in sand and silt Type of flow

Type of bedform

Manning’s n

Relative sediment transport

Lower flow regime Subcritical flow, Fr < 1

Ripples Dunes with superposed ripples Dunes Plane bed

0.018–0.030

Small

0.020–0.040 0.010–0.013

Large

0.010–0.012

Large

Upper flow regime Critical flow, Fr = 1 Upper flow regime Supercritical flow, Fr > 1

Antidunes:€Standing waves Antidunes:€Breaking waves Chutes and pools

0.012–0.018 0.018–0.030

Note:€Values of Manning’s n from Chow, 1964

of (a) sediment transport, (b) flow resistance and (c) amount of material transported. Different bedforms may occur next to each other in the same cross-section of the channel or replace each other over time as the discharge and velocity of the river change crosschannel (Table 8.2). In the lower flow regime, ripples and dunes cause high resistance to flow primarily by form roughness. The sediment transfer is carried out by individual grains moving up the back of ripples or dunes and sliding down the steep front face. The surface of water is out of synchronisation with the lower regime bedforms. This creates a zone of separation downstream of the crests resulting in surface eddies (Fig. 8.4). In contrast, resistance to flow is small and sediment transport large in the upper flow regime, which is commonly seen when rivers are in flood. As the water velocity rises, dunes are replaced by plane beds, although both dunes and plane beds may occur simultaneously in a channel. Antidunes are seen in extreme cases with very high flood velocities. Chutes and pools are rare and only seen in waterfalls and rapids. The resistance to flow is due not to form roughness but grain roughness. In plane beds, grains form a sheet which has the thickness of a couple of grains, and continuously roll downstream. In antidunes, in spite of sediment grains moving downstream, the forms move upstream. When breaking antidunes occur, much bed material is briefly suspended in water and then is partially stored on the bed. The surface of the water stays in phase with the bedforms (Simons and Richardson, 1971). Such bedforms are rarely seen in gravel, except as poorly developed dunes and ripples. The commonest form is a flat gravel bar, which is often seen after the passage of a flood which has winnowed out all the fine material. Individual pebbles, cobbles or boulders on this bar may be imbricated upstream.

135



Fig. 8.4

Bars

Alluvial bedforms and the nature of the flow in the river. After Simons and Richardson, 1971

8.4╇ Bars Bars are transitional channel-features built by the deposition of sediment carried by the stream. They are located at a level higher than the bed but lower than the bank top, and are exposed during periods of low flow. The tops of the bars are therefore visible most of the time.

136

Alluvial valleys

The location, size, shape and structure of bars depend on a range of channel properties including channel geometry, gradient, variation in flow, and volume and texture of sediment. Different types of bar therefore tend to be associated with variations in channel pattern. Alternate bar are seen in straight channels, located opposite channel pools. Meandering rivers display point bars (as discussed in Chapter 7) on alternate sides, following meander bends. Braided streams mainly have bars in mid-channel but, in many locations, a scouring of the banks and widening of the channel also lead to the deposition of bars attached to a bank. The location of the bars in a rock-cut channel may also be associated with a rock protrusion on the floor or the accumulation of a pile of boulders forming the core of the bar and accelerated sedimentation in finer material round it. Bars are usually formed at the falling stage of a hydrograph and carry the history of their formation in the texture and structure of their sediment. They also reflect post�depositional changes in the form of shallow flood channels or sand ripples on their top surface. As expected, bars are common in rivers where the discharge fluctuates and sediment is available, often from the channel and the floodplain. During the dry season, the beds of many seasonal rivers are an assemblage of barforms with one or several narrow channels separating them. Such bars may disappear entirely during the wet season to reappear in a renovated form in the next dry season.

8.5╇ Floodplain A floodplain is a strip of land with very little relief that borders a river and is inundated at times of high water. As the extent of the inundation depends on the size of the flood, floodplains are described according to flood probability, for example, an n-year floodplain. The rarer the flood, the more extensive is the floodplain. Floodplains comprise a number of distinctive features: • the river channel • point bars • scroll bars • depressions and risers on bar tops • oxbow lakes (cut off sections of a meander bend) • natural levees • crevasse splays • backswamps. The first four of these features generally occur inside the channel of the river. Abandoned and degraded portions of the former channel with these features may occur away from the river’s present position to form part of the floodplain. Levees are raised linear features that run parallel to rivers. As a river overflows, hydraulic conditions change at the top of the banks and active sedimentation in fine sand, silt and clay takes place, building levees along its channel. Levees are generally the highest feature in a floodplain, disrupting drainage to and from the river, and thus swampy conditions form over the rest of the floodplain. These are the backswamps that cover most of the floodplain. They are bounded on the side close

137

Floodplain

to the river by levees and on the other side by hillslopes and small screes or alluvial fans of tributary streams. The smaller tributary streams could be blocked by a levee, forcing them to flow parallel to the main stream for some distance until they could merge. Such tributaries are known as yazoo streams, the term originating from the valley of the Lower Mississippi. Such streams are common on the wide floodplains of large rivers, as on the north bank of the Ganga between its confluences with the Gandak and Kosi. The presence of levees on both banks also causes the main river to deposit sediment inside the channel. Given high sediment availability and an efficient levee system, this may cause the bed of the river to rise and, in extreme cases, to reach a level higher than the backswamps. A collapse of part of the levee in a flood at this stage is a disaster leading to widespread and fast inundation of the backswamps. Such a widespread levee collapse happens rarely. Instead cracks in the levee may be enlarged in floods and sediment-laden floodwater may erode its way through. As a result, a cone of sediment often breaks through the levee to spread down the slope towards the backswamps. The cracks are called crevasses and the cones are called crevasse splays. A number of crevasse splays occur along the levee of a large river. They are identifiable by their relative coarseness, as most of the crevasse splay material is sand which spreads over the silty and clayey floodplain. Coleman (1969) described the crevasse splays of the Brahmaputra (locally known as the Jamuna) in Bangladesh. Most of the floodplain, therefore, is dissected by tributaries and abandoned and overflow channels of the main river. The low gradient across the floodplain creates an environment of stagnant or slow-flowing water and deposition of fine material. A number of these abandoned channels, especially the cutoffs, tend to be filled with clay. Floodplains can be formed in three ways: • lateral accretion from movement of meandering rivers • vertical accretion from overland flows • island formation and their consolidation and channel abandonment by braided rivers. The classical model of floodplain formation by lateral movement and building of point bars was illustrated by Leopold and Wolman (1957) and Leopold (1973), who for years monitored surveyed sections across a small stream, the Watts Branch in Maryland, USA. In this model, a meandering stream moves laterally across its valleyflat, eroding on the convex side and building point bars on the concave bank. Thus over time the entire floodplain is built and the river maintains its dimensions as it shifts. An account of vertical accretion, the second process, is provided by Schumm and Lichty (1963), who recorded the reconstruction of a floodplain previously eroded in a large flood on the Cimarron River, Kansas, USA. Many floodplains are perhaps formed by a combination of both processes, although the lateral movement of a meandering channel usually forms most of them. This is reflected in the sedimentary structure of the floodplain, which has a layer of gravel at the base (bedload), a layer of sand (some cross-bedded) on the gravel (indicating lateral movement and point bar formation) and a thin layer of fine sediment on top (overbank deposition). This, however, is a simplified description. Floodplain sediment also includes deposition from crevasse splays or abandoned channels, etc. on a minor scale. Braided rivers build floodplains by consolidating the mid-channel islands by filling the intervening channels as the stream shifts.

138



Fig. 8.5

Alluvial valleys

Narmada floodplain inside the high banks of the large flood channel in alluvium. Note mudcracks in the clay. Photograph: A. Gupta Dietrich et al. (1999) summarised the floodplain deposition of the Fly River in the wet equatorial hilly environment of Papua New Guinea as due to (1) advection of sediment with overbank flows, (2) lateral diffusion from sediment-rich water of the Fly River and (3) sediment-laden water travelling upstream into tributaries or small floodplain channels. In general, the Fly has a low rate of channel migration and a limited current volume of sediment deposition in the floodplain. Dunne et al. (1998) also mentioned the transfer of sediment-laden water up secondary channels into the floodplains filling water bodies. A floodplain is a place for storage. A grain of sediment is stored in the floodplain for years until the main channel or a tributary comes across, erodes the alluvial fill and picks up the same grain of sediment. In certain rivers, lateral movement of sediment may be comparable in volume and speed to downstream transfers. Descriptions of such movements have been provided for the Amazon by Dunne, et al. (1998) and Mertes and Dunne (2007). Sediment grains move downstream in a number of steps, interrupted by long periods of storage. It may take a sediment grain centuries, even millennia, to reach the end of the journey at the mouth of the river (Leopold et al., 1964; Madej and Ozaki, 1996). In the humid tropics, this long storage results in chemical weathering of the stored grains of sediment, ending in the formation of clay minerals and a proportional increase in quartz. The sediment that reaches the sea is compositionally not the same sediment that came off the mountains. Meade (2007) discussed this for the Amazon and Singh (2007a) for the Ganga. Rivers in the seasonal humid tropics that are also affected by large floods have been described in Chapter 7. Such rivers have a nested appearance with high banks enclosing the entire channel required to transfer the discharge of high-magnitude floods. As the river cannot normally climb out of the high banks, deposition related to annual high flows occur between them. Figure 8.5 illustrates a floodplain built by clay and silt within the high banks of the Narmada River. A similar pattern of deposition, except in coarse material, is shown for the Yallahs River at Mavis Bank in the Blue Mountains of Jamaica (Fig. 8.6). The



139

Terrace

Fig. 8.6

Valley of the Yallahs River, Mahogany Vale, Blue Mountains, Jamaica. Note deposition of coarse sediment derived from upper basin. Photograph: A. Gupta

photograph was taken in 1971:€seven and a half years after Hurricane Flora (1963) caused tremendous flooding in these rivers. The channel was cleared of all except perhaps very large boulders. Some of the coarse material was deposited to form a high terrace (under young trees). The Yallahs then rebuilt its floodplain and a smaller channel as required by its common pattern of discharge within the big channel. Its valley now demonstrates a river channel, a floodplain and a hurricane terrace (or a hurricane floodplain) in coarse material. This is the channel-in-channel physiography.

8.6╇ Terrace Terraces are fragments of former floodplains which are no longer flooded as they are currently at a level higher than the present floodplains. Terraces are generally located away from the river, behind floodplains, rising sharply from the floodplain level. The rise (also called the scarp) and the flat top (termed tread) together constitute a terrace. Terraces are formed when a river downcuts, lowering its bed to a level from which the old floodplain cannot be inundated. The old floodplain becomes dissected and fragments of it survive as a terrace; a new floodplain is deposited below it and above the riverbed. A river lowers its bed elevation due to several causes. A drop in sea level, or a lake level which was operating as the base level of a river, would expedite downcutting. This leads to terrace formation along the lower course of a river. These are eustatic terraces. In tectonically active areas, almost always in the mountains, tectonic uplifts steepen the river gradient, contributing higher energy for deepening its channel and forming tectonic terraces. Flights of such terraces are seen where a river emerges from a gorge section in a tectonically active

140



Fig. 8.7

Alluvial valleys

Terraces along the Jinsha Jiang, southwestern China; the river is emerging from the Tiger Leaping Gorge. Photograph: A. Gupta area (Fig. 8.7). Climate changes during the Quaternary have been suggested to explain the formation of terraces in alluvial material by the downcutting of the Upper Godavari and Krishna Rivers in southwestern India (Kale and Rajaguru, 1987). Extensive low terraces may occur in the valley of large rivers. Singh (2007a) has described two large alluvial terraces above the floodplain of the Ganga in its middle course. The first terrace (termed T1) is at a level 5–10 m above the floodplain and the next one (T2) is 5–10 m higher. Abandoned channels are common on terrace surfaces. Multiple terraces are seen following repeated tectonic movements or negative base level changes. Such terraces will occur on both sides of a river and could be paired or unpaired and at alternate heights. Terraces at approximately the same level on both sides of a river are known as paired terraces. These are formed when the river downcuts rapidly. If the terraces are not at the same level on either side, they are called unpaired terraces. These are formed with a slow downcutting, giving the river time to meander and destroy the old floodplain on both sides alternately (Fig. 8.8). Multiple terraces, between four and six in number, occur in the tectonically active valleys of the Himalayan headwaters of the Ganga River (Khan et al., 1982). Such terraces have also been reported from the Spiti Valley, Himachal Pradesh, India (Sah and Virdi, 1997). Multiple terraces are quite common in tectonic mountains. Terraces can be entirely alluvial or underlain by hard rocks, leading to their identification as alluvial or bedrock (strath) terraces.

8.7╇ Valley margins The edge of an alluvial valley is usually marked by hillslopes, small screes from mass movements and alluvial fans (Chapter 13) built by tributary streams. Unless the hills are

Questions

141



Fig. 8.8

Paired and unpaired terraces tectonically active and/or drained by an eroding tributary, not much sediment is transferred across the valley to the main channel. Instead, the transfer of sediment happens downstream along the main stream and also laterally into its floodplain from the main channel.

8.8╇ Sediment transfer along the valley axis Changes in sediment deposited along a valley axis are related to the size of the river. A long river will show the difference between the steep upper reaches, the low-gradient mixed bedrock-alluvial valley and the true alluvial valley near the coast. However, as large rivers (a) tend to be polyzonal (different parts of the basin contributing water and sediment via tributaries at different volumes and rates); (b) flow across varying geology; and (c) are long-lived, they carry a history of adjustment to climate changes from the past, the final pattern may be complicated. There should be an overall trend but local deviations from it should be expected. The sediment in channel bars, floodplains and terraces reflect all this (Blum and Törnqvist, 2000; Blum, 2007).

Questions 1. Sand is a very common material deposited in an alluvial channel. Do you agree? Why?

142



Fig. 8.9

Alluvial valleys

Flood probability analysis of Baitarani River at Akhuapada, Orissa, India. Plotted from data in Ahuja and Majumdar, (1959). From Gupta, 1988, by permission of Wiley. 2. What are the effects of the rock ribs in the Mekong as shown in Figure 8.3 on sediment accumulation in the channel? 3. Can you explain the sedimentary structures shown on the side of the bar in the Jamuna (Fig. 9.11)? 4. Define a 50-year floodplain. Can two 50-year floods occur in successive years? 5. Why is there a progressive increase downstream in the proportion of quartz grains in river sand? 6. Identify the alluvial valley features in Figure 8.6. Rank such features in the order of length of survival. 7. What is a polyzonal river system? Provide several examples. 8. Figure 8.9 displays the flood potential of the Baitarani River at Akhuapada, Orissa, India. The data run from 1874 to 1957 (Ahuja and Majumdar, 1959). Gumbel Extreme Value Distribution has been used to plot the regression line. (a)╇Using the fitted line, find the flood discharge for the recurrence intervals:€1.1, 1.5, 2.33, 5, 10, 50 years. (b)╇Would it be acceptable to use this diagram to determine the 100-year flood? The 500-year flood? Justify your answer. (c)╇Calculate the unit discharges for the same floods. Unit discharges are calculated by dividing the total discharge by the drainage area. Upstream drainage area at this station is 11 360 km2. Can you identify any trend in the unit discharge figures?

9

Large rivers in the tropics

A huge river is the Padma, just like the sea Buddhadev Bose, translated from Bengali

9.1╇ Introduction A number of large rivers drain the tropics, acting as huge conduits for transferring water and sediment to the oceans. These include the Amazon, Orinoco, Magdalena and Paraná in tropical South America; the Nile, Niger, Congo and Zambezi in Africa; and the Indus, Ganga, Brahmaputra, Irrawaddy and Mekong in tropical South and Southeast Asia. These rivers play a very important role in tropical geomorphology. We review a selected number of large rivers in this chapter to illustrate their importance.

9.2╇ Characteristics of a large river A proper definition of a large river is elusive. Potter (1978b) listed four essential properties for large rivers:€size of the drainage basin, length of the main river, magnitude of water discharge and volume of sediment transported. He then listed 50 largest rivers that collectively drain about 47 per cent of the land surface, excluding Greenland and Antarctica. All except one of these rivers are more than 103 km long and the smallest drainage basin is 105 km2. Good data on the other two properties, discharges of water and sediment, are not available for all major rivers, especially regarding their bed load. It is very difficult to measure the bed load of a large river. Table 9.1 lists selected characteristics of the rivers of the tropics that are longer than 2000 km. Meade (1996) provided two lists, one of which ranked the top 25 rivers with the biggest discharge and the other included the top 25 with the largest suspended sediment load. The two lists do not coincide, as certain rivers such as the Zambezi carry a large water discharge but a low sediment load. Sediment also accumulates behind impoundments in several rivers with only a fraction reaching the sea. Such a river no longer functions as a natural conduit controlling a continental-scale drainage basin. The Nile below the Aswan Dam is a good example. It is expected, however, that a large river is long, drains a large basin, carries a very big discharge and commonly, but not always, transports a large volume of sediment. The great length of the rivers allows them to flow across a range of environments. For 143

Large rivers in the tropics

144

Table 9.1╇ Selected properties of large rivers in the tropics (ranked according to discharge)

River Amazon Congo Orinoco Ganga– Brahmaputra Mekong Paraná-Uruguay Irrawaddy Zhujiang Salween Indus Magdalena Zambezi Niger

Annual average water discharge (109 m3) 6300 1250 1200 870 470 470 430 300 300 240 240 220 190

Length (km)

Drainage area (106 km2)

6000 4370 770 B-2900 G-2525 4880 3965 2010 2197 2820 3000 1540 2575 4100

5.9 3.75 1.1 G-1.06 B-0.63 0.79 2.6 0.41 0.41 0.27 0.97 0.26 1.32 2.27

Annual average suspended sediment load (106 t) 1000–1300 43 150 900–1200 150–170 100 260 80 about 100 50 220 20 40

Note:€Drainage areas are rounded off to 106 km2 to avoid discrepancies between various sources. Figures from Gupta 2007a:€Table 1.1. For primary sources see references therein. The table does not include the Nile because, in spite of a length of 6500 km, its current water and sediment discharges are not comparable with figures in this table. Source:€Gupta, 2007a and references therein

example, the Mekong flows in and out of narrow rocky valleys and wider alluvial basins. Hydrological conditions also vary along their courses. Different parts of their large basins contribute water and sediment in different fashions to the mainstream. The end part of the river has to adjust to all such variations including changes in sea level. Large rivers are therefore recognised as polyzonal. The origin and sustenance of big rivers are usually associated with orogenic belts, as the Amazon is with the Andes and the Ganga with the Himalaya. Part of their upper course may follow a geological lineation like a long section of the Upper Indus that remains pinned to a suture zone. Certain large rivers, however, start in rift valleys and cratonic settings. Neither the Congo nor the Zambezi, which rise on the opposite sides of a common divide, originates in an orogenic belt. Structural controls on a continental scale, like a downwarp, may determine the location of the river even in lowlands. The Ganga, for example, flows through a downwarped foreland parallel to the Himalaya which is related to the tectonic origin of the mountains. The Amazon connects the Andes to the South Atlantic Ocean through a tectonic downwarp. Even away from the steep headwaters, smaller crustal deformations and transverse belts of resistant lithology may regionally alter the low gradient and valley width of the rivers, resulting in repeated changes in channel form and behaviour at the scale of tens and hundreds of kilometres. Rivers flow over a steeper gradient in a narrow valley in such

Characteristics of a large river

145

Table 9.2╇ Large rivers:€tectonic settings and morphological variability Tectonic setting

Description

Tropical example

Continental collision belt:€rivers: Longitudinal foreland basin

Rivers with headwaters in the mountain range at converging plate boundary and the main stream flows axial or longitudinal to the mountain range Rivers with headwaters in the mountain range at converging plate boundary but the main stream flows transverse to the mountain range across a stable platform or craton Rivers that flow within mountain ranges parallel to major structural trends almost to their mouths River basins formed by creation of new relief in rift and rift-shoulder settings River systems formed by cratonic tectonic events unconnected to plate tectonics, e.g. large-scale igneous doming

Ganga, Paraná

Continental collision belt rivers: Transverse systems

Continental collision belt rivers: Intrabasinal setting Rift systems

Cratonic settings

Amazon, Orinoco

Brahmaputra, Irrawaddy, Mekong

Lower Paraná, Orange

After Tandon and Sinha 2007

locations, and are likely to deposit less and transfer more sediment. Tectonics, climate and eustasy combine to control a long river in its entirety, but any of these factors may dominate part of its course. In sum, large rivers tend to have their headwaters in the uplifted large mountains, but not always. They tend to drain continental-scale cratonic areas (Tandon and Sinha, 2007) usually through wide low-gradient alluvial plains. Tandon and Sinha (2007) summarised the tectonic setting and morphology of large rivers (Table 9.2). The age of large rivers varies but some have existed for a long time. The extra-tropical Mississippi is at least as old as the late Jurassic. The entire present course of a river, however, may not have come together at the same time. It could be an assemblage of several subunits of different age and appearance (Gupta, 2007a). The nature of the river also changes over time. Goodbred (2003) has demonstrated the changing pattern of the Ganga over the last 50 ka in response to climate and sea-level changes. Rivers of this size are supported by huge amounts of precipitation falling on their basins. This requirement is met by heavy rainfall of the equatorial and monsoon tropics. Snowmelt forms a substantive part of the baseflow for rivers originating in the high mountains. The mainstem discharge is augmented cumulatively by a number of tributaries draining different parts of the basin, but the sediment supply often comes predominately from the source-mountains of high relief and the tectonically fractured rocks and oversteepened

Large rivers in the tropics

146

(a)

(b)



Fig. 9.1

Average discharges of (a) suspended sediment and (b) fresh water in the Orinoco River and its tributaries. From Meade, 2007. By Permission of Wiley slopes (Milliman and Syvitski, 1992). The rest of the basin usually provides little sediment but augments the discharge of the main river cumulatively at tributary junctions. This is illustrated by Meade (2007) for the Orinoco, which receives its sediment from the Andes and its foreland, but water from all parts of the basin (Fig. 9.1).

147

The Amazon

Large rivers with headwaters and tributaries in the high mountains contribute large volumes of sediment to the oceans (Table 9.1) unless inhibited by impoundments, as in the case of the Indus. The transfer of this sediment along a long route occurs in pulses and is characterised by short transfer times and long storage periods, especially when sediment could be stored in floodplains. A number of rivers − the Amazon, Ganga, Brahmaputra − went through large shifts in discharge following climate change and meltwater floods at the end of the Pleistocene and in the early Holocene. Sediment of this time may still exist in the channel or floodplain of the river. Over the last several millennia, anthropogenic activities such as deforestation, impoundments and land use changes have altered the characteristics of many river basins and channels. Several case studies in this chapter illustrate the way water and sediment reach the mainstream in large basins and are carried out to the coastal waters and beyond, building deltas in the process. A very large part of the tropical continents is drained by large rivers and the dynamics of their drainage basins is an essential component of geomorphology in the tropics. These basins integrate a variety of structure and processes on a sub-Â�continental scale.

9.3╇ The Amazon 9.3.1╇ The physical setting The Amazon is well studied, in spite of its size (Fig. 9.2). It still remains in natural conditions and is not modified by anthropogenic activities. Plate tectonics control the location and lithologic and topographic frameworks of the Amazon (Potter, 1978b). The basin was delineated following the Miocene uplift of the Andes due to the subduction of the Nazca plate below the South American. It is generally thought that the rise of the Andes created this huge eastward drainage system. The tectonically and volcanically active Andes Mountains slope down towards the east to a foreland basin. The Amazon flows from the Andes across the foreland eastward through an axial trough, which is bounded to the north and south by the crystalline rocks of the Guyana and Brazilian Shields. A fair amount of knowledge exists regarding the underlying structure along the channel and floodplain because of exploration for petroleum (Mertes and Dunne, 2007). Evidence from deep cores indicates a 6000-m deep east–west crustal sag below the basin axis that ends below the mouth of the Amazon in a graben, the Marajó Rift. Precipitation over the Amazon Basin is influenced by the annual shifting of the Inter Tropical Convergence Zone (ITCZ) and the South Atlantic Convergence Zone (SACZ) over the Andean region. Annual average precipitation is about 2000–2500 mm, approximately uniformly distributed over most of the basin with the maximum of 7000–8000 mm on the lower eastern slopes of the Peruvian Andes. The extreme northeastern and southern parts of the basin are drier. The rains arrive first over the southern basin in November to December, and then move north. The annual hydrograph of the river (Fig. 9.3) is unimodal, damped and delayed by the sheer size of the basin, length of the drainage network

148



Fig. 9.2

Large rivers in the tropics

The Amazon. Top diagram:€generalised map of lithological zones and structural features. Note the river crossing four intercratonic arches and the Monte Alegro ‘intrusion’. Top: from Dunne et al., 1998. Bottom: from Meade, 2007. By Permission of Wiley and storage of water in the enormous floodplains, the cumulative size of which exceeds 100€000 km2. Both precipitation and river discharge are affected in certain years by the El€ Nino Southern Oscillation (ENSO), low flows of the Amazon being associated with the El Nino years (Mertes and Dunne, 2007). Such climatic fluctuations affect flooding and sedimentation (Aalto et al., 2003) but the effect on the morphology of the rivers of the Amazon Basin is yet to be understood.

The Amazon

149

(b)

(a)

(c)



Fig. 9.3

Monthly rainfall and river discharge (1972–1996) at three locations on the Amazon River. From Mertes and Dunne, 2007. By permission of Wiley The axial trough of the central Amazon Basin exhibits a remarkable suite of fluvial landforms (section 9.3.3) under an extensive forest cover. The forest is interspersed with savanna and recently deforestation has occurred towards the southern and eastern margins of the basin.

9.3.2╇ Transfer of water and sediment along the Amazon About 80 per cent of the Amazon’s sediment is derived from the Andes and its foothills, although they cover only 800 000 km2 of this 6 million km2 basin. The rocks in the Andes are primarily volcaniclastic and marine sedimentary rocks, metamorphosed to various degrees. These mountains have been supplying the Amazon with 500–600 million tonnes of sediment throughout the late Cenozoic. Almost the entire sediment arrives from the Peruvian Andes by the Amazon mainstem or from the Bolivian Andes via the Madeira (Meade, 2007). Rivers in the Andes Mountains are steep and flow on bedrock or gravel. In contrast, the river gradient across the subsiding Andes foreland is only a few centimetres per kilometre. The Amazon has very low stream power, about 12 Wm−2, even at average peak discharge, and only sediment finer than about 0.5 mm (sand, silt and clay) is carried into the lowland Amazon Basin. Rivers start to meander in the downwarped foreland zone, and

150

Large rivers in the tropics

high sediment deposition leads to the building of bars and shifting of channels. Tributaries draining land below the Andes and the Guyana and Brazilian Shields bring in very little sediment but contribute high volumes of discharge. In fact, two of the biggest tributaries of the Amazon, the Madeira and Negro, contribute more water to the Amazon than is discharged by any other river except the Congo, Orinoco and the Changjiang. Meade has further observed that the freshwater discharge at the mouth of the Amazon is five times that of the Congo, six times that of the Orinoco and 12 times that of the Mississippi (Meade, 2007). All these are huge rivers and the comparison indicates the scale of the Amazon and its importance in global hydrology. The average annual sediment of the Amazon at Óbidos is about 1200 million t, at present exceeded only by the combined flow of the Ganga–Brahmaputra. Far higher sediment, however, is transferred laterally in the Amazon between the channel and the floodplain than downstream along the channel as calculated by Dunne et al. (1998), who studied sediment transport through the 2010 km of the lowland Amazon in Brazil. The lateral exchange involves bank erosion, bar deposition in the main channel, settling from overland flow on the floodplain and sedimentation inside channels within the floodplain. Much of the sediment that leaves the channel in suspension during floods and enters the floodplain is deposited there before clearer waters return to the Amazon during the falling stage of the annual hydrograph. This is particularly demonstrated in the lower Amazon where the large floodplain lakes are silting slowly but progressively. Sediment grains are stored in the floodplain of the Amazon for centuries and millennia. Mertes et al (1996) Â�estimated the mean recycling time to be between 1000 and 2000 years for the floodplain sediment in the reach between the confluences of the Jutaí and Madeira. This allows chemical weathering of lithic fragments, production of clay minerals and a progressively quartz-rich Â�sediment. Ecologically, with annual inundation alternating with exposure, the floodplain of the Amazon is a remarkable landscape (Junk, 1997). The pattern of sediment transport is influenced by (1) the hydrology of the Amazon and (2) geomorphic characteristics of the channel and floodplain. Regional and local tectonics determine the second to a considerable extent. The pattern that persists is the deposition of floodplain sediment and its subsequent removal by the shifting river (Meade, 2007). The tidal limit is approximately up to Óbidos, almost 1000 km from the mouth of the Amazon, due to the very low gradient of the river. It is difficult to measure sediment fluxes in the end part of a river, but a large amount of the river sediment is apparently being deposited on the floodplain, mobilised inside the channel of the Amazon, and also entering the Atlantic Ocean. About half of the sediment that passes Óbidos settles on the continental shelf beyond the mouth of the Amazon (Kuehl et al., 1986). A substantial part of the sediment is carried northwards and deposited along the coast up to the edge of the Orinoco Delta. Meade (2007) suggested that the amount reaching this delta, about 1600 km from the mouth of the Amazon, could be as high as 100 million tonnes. This is not a precise number, but anything in that range indicates the size of the Amazon and its importance in the geomorphology of the American tropics. A different pattern of sedimentation seems to have operated on the continental shelf near the mouth of the Amazon during the Last Glacial Maximum. It has been proposed that the sea level dropped 120 m during the last regression and the Amazon incised to 12–25€m below its present bed and flowed on a steeper gradient. This allowed the transfer

151

The Amazon

of sediment offshore to build the Amazon Cone (Fig. 9.2). Subsequently, as the sea level rose and accommodation space became available, considerable sedimentation happened in the lower river and its floodplain. Less material reached the continental shelf. The reduction in the number of lakes between the middle and lower courses of the Amazon supports this interpretation.

9.3.3╇ Morphology of the Amazon River Beyond the Andes foreland, the Holocene Amazon floodplain is incised more than 10 m below a landscape of low hills under thick forest cover. The hills had been dissected out of Tertiary and Quaternary lacustrine and alluvial deposits that fill a central depression. The Holocene Amazon has incised its channel and floodplain in the middle of this depression and the surface of the floodplain displays a complex pattern of channels of different size, scroll bars, levees and various kinds of depressions such as cut-off lakes and backswamps. The floodplain is commonly flanked by discontinuous terraces, 5–15 m above the usual inundation level. The Amazon in Brazil has a remarkably straight and anastomosing channel within its floodplain. Avulsions related to flow switching are common, and where the channel bends it usually undercuts the cohesive terrace material above its Holocene floodplain. Over a measured length of 2000 km in the downstream direction, the low water width of the river increases from 2 to 4 km and the corresponding depth from 10 to 20 m. The floodplain contains a dense network of channels, some of which are as big as large tributaries. The large channels stay connected to the Amazon all year but smaller ones may become detached during the dry season, although they continue to hold water. The main channel, anabranches and floodplain channels all migrate and, as a result, numerous scroll bars and depressions occur on the surface of the floodplain. Mertes and Dunne summarised the Amazon as an entrenched river that is confined to its valley, remains straight and is relatively immobile over a distance of hundreds of kilometres (Mertes and Dunne, 2007). In general, the low-gradient Amazon flows through a 40–50 km wide floodplain with considerable complexity. Anabranches, levees, scroll bars and large-scale channel migration or avulsion are common. A complex mosaic of lakes, lake deposits and overbank sediments are also characteristic features in a wide floodplain. The smaller of these lakes could be due to small low-sediment tributaries being dammed by rapid alluviation on the floodplain of the Amazon. Larger lakes, up to 65 km across, occur downstream. The Amazon, however, in spite of flowing on top of thick sedimentary layers along its central valley axis, is affected by a set of buried transverse structures (Fig. 9.2). Four major structural arches (Iquitos, Jutaí, Purús and Gurupa) and the Monte Alegro Intrusion, further to the east, modify the river in their vicinity. The Amazon straightens its course as it crosses the structural highs. On a slightly steeper gradient of the water surface, the floodplain narrows, the river tends to be at the foot of the terraces, scroll bars are limited to channel margins and channel migration becomes limited. Gravity measurements for the lower Amazon region relate direction change and form of the river with gravity anomalies (Nunn and Aires, 1988), indicating that the Amazon crosses its floodplain only in specific places. Even the biggest river of the world is structurally controlled.

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Large rivers in the tropics

Other large-scale fracture patterns control the tributary network of the Amazon. For example, Franzinelli and Igreja (2002) associated the alignment of the lower Negro River with a NW–SE trending tectonic lineament. Here intersecting sets of faults have created sunken crustal blocks and depressions along a half-graben that is inundated to a depth of 20 m across a 20-km wide reach of the river. The Negro carries very little sediment and this small amount is flushed through the fault-lined depression with accumulation of only little sand bodies along the margins of the river. The banks of the Negro are in cohesive cliffs of lacustrine deposits and bedrock is exposed in the river channel due to a lack of sediment to cover it up (Franzinelli and Igreja, 2002). Latrubesse and Franzinelli (2002) related straight reaches of the Solimões (middle Amazon) to recent activities along a set of NW–SE and SW–NE trending fractures. These fractures were originally mapped in 1952 from river and lake alignments by Sternberg and Russell. Latrubesse and Franzinelli (2002) interpreted the confluence of the Purús with the Amazon also as tectonically controlled. It coincides with a wide V-shaped sunken tectonic block. Three factors interacted through the late Cenozoic and particularly the late Quaternary to determine the form of the Amazon channel and its floodplain. These are the basin tectonic setting, climate and sea-level fluctuations. Together these factors controlled erosion, sediment transport and deposition in the Amazon valley.

9.4╇ The Zambezi 9.4.1╇ The physical setting Unlike other large rivers described in this chapter (the Amazon, Ganga–Brahmaputra and Mekong), the Zambezi does not rise from a tectonic mountain system. In contrast, it rises from a small spring on a relatively low divide that separates it from the northwest-flowing headwaters of the Congo. Its extreme upper course is characterised by several directional changes and supported by steady drainage from a series of small pans (swamps) within broad lowlands. None of these pans extends for more than 4 km in diameter, and they are either perennially waterlogged or fed by marginal springs. Discharge to the upper Zambezi is also augmented by base flows from shallow grassy valleys with high water tables and anaerobic soils, locally known as dambos. Moore et al. (2007) have described these dambos as sponges that support the Zambezi. The river flows through a number of gorges and wide basins, a pattern which has been attributed to an evolutionary history of river capture and course changes that started before the disruption of Gondwana about 120 ma ago. Moore et al (2007) have divided this 2575 km long river that drains a 1.32 million km2 basin into three major sections, each with characteristic geomorphology. The upper Zambezi extends from the source to the Victoria Falls, the middle river from the Falls to Cahora Bassa Gorge at the edge of the Mozambique coastal plain and the lower Zambezi flows over the coastal plain to the Indian Ocean (Fig.€9.4).



153

The Zambezi

Fig. 9.4

The Zambezi drainage system. Rift basins:€G, Gwembe Trough; MP, Mana Pools basin; C, Chicoa or Cahora Bassa Basin. Gorges:€B, Batoka; K, Kariba; M, Mupata; CB, Chora Bassa. Rapids and Falls:€1, Chavuma; 2, Ngonye; 3, Katima Mulilo; 4, Mambova; 5, Katombora; 6, Victoria. Floodplains:€Bar, Barotse; MM, Mulonga-Matable; OK, Okavango Delta. Pans:€MP, Makgadikgadi. From Moore et al., 2007. By permission of Wiley

The upper Zambezi starts as a small channel cut into Precambrian basement rocks before opening out into a low-gradient shallow stretch through the Barotse floodplain that ends at the Ngonye Falls. Downstream of the Ngonye Falls, the gradient of the river begins to steepen and the river downcuts into a set of rapids to flow over the 1700 m long and more than 100 m high Victoria Falls, also known as Mosi-a-tunya (the smoke that thunders). The middle river starts in a zigzag gorge banked by low escarpments before turning abruptly east into the 100 km long Batoka Gorge inside which the Zambezi narrows to about 25–30 m in the 6 m high Chimamba Rapids over basalt. The difference between the upper and the middle Zambezi is perhaps best appreciated by its change in average gradient, from 0.00024 above the Victoria Falls to 0.0026 below it. Downstream of the Batoka Gorge, the Zambezi flows through a succession of broad basins separated by narrow defiles, altering its form and behaviour at every change. These basins, often seen in popular television documentaries as important wildlife habitats (such as Mana Pools), have been designated world heritage sites and the gorges (such as Kariba) have been used for power generation with large lakes inundating the canyons. The middle Zambezi flows into the last of the gorges (Cahora Bassa) to emerge onto the coastal plain as the lower Zambezi. The lower Zambezi is a different river. It forms a 100-km long floodplain–delta system of oxbows and swamps and meander bends. Old distributary channels are found along a 280-km stretch of the Indian Ocean coast between the cities of Quelimane and Beira. A

154

Large rivers in the tropics

6.4 m high tidal range, a 40–50 km incursion of tides up the river and coast-parallel low ridges found up to 30 km inland and indicating former high sea levels mark this delta-face (Moore et al., 2007). A number of attempts have been made to reconstruct the history of the river. Such reconstructions are based on a multitude of characteristics: • alternating pattern of gorges and wide basins • changes in flow directions of the tributaries (several of which are big rivers themselves) forming elbow-bends • sedimentary sequences on the coastal plain and offshore • palaeocurrent directions in fluvial sedimentary rocks dating back to the upper Cretaceous • alignment of buried valleys under fluvial sediments • distribution pattern of fish, fauna and antelopes in the basin. In sum, the process of drainage evolution started in the early Cretaceous following the disintegration of Gondwana and the opening of the Atlantic Ocean. The drainage system that currently terminates in the inland Okavango Delta, different sections of the present Zambezi that probably flowed as separate systems earlier, and the proto-Limpopo River south of the current Zambezi basin, all were involved at a certain stage of its history. The current Zambezi system evolved over time following headward erosion of sections of the proto-Zambezi, a succession of river captures, updoming of parts of the continental surface, the presence of low relief over much of the area, and tectonic activities along faults with a NE–SW orientation. This history of drainage development has been discussed in detail by de Wit (1999), Moore and Larkin (2001), Moore and Blenkinsop (2002) and Moore et€al. (2007). This explains the changing characteristics of the Zambezi along its course, the elbow bends in tributary rivers, the occurrence of past river sediments, the presence of huge palaeo-lakes and the pattern of current speciation of plants and animals.

9.4.2╇ Water and sediment The average annual discharge of the Zambezi is 220 m3, befitting a major river, but it carries a rather low sediment load, on average only 20 million tonnes of suspended sediment (Meade, 1996). A number of dams and reservoirs (Fig. 9.4) undoubtedly impound a significant proportion of its sediment load, but basically it lacks the presence of an unstable tectonic mountain system within its basin, unlike the other large rivers reviewed in this chapter. The pattern of discharge is seasonal (Fig. 9.5), with most of the rain arriving in the southern hemispheric summer due to the meeting of the northeast monsoon and southwest trade winds across the Inter Tropical Convergence Zone over the middle and lower Zambezi, and that of the southwest monsoon and southeast trades over the upper basin across what is known as the Congo Air Boundary. The two air boundaries move north in winter resulting in drier months. Most of the annual rainfall of 600–1300 mm falls over the northern basin, which contributes a higher runoff to the Zambezi (Moore et al., 2007). Like its sediment load, the downstream passage of the Zambezi’s water is modified by the number of dams and reservoirs, such as the Kariba and Cahora Bassa. This certainly influences



155

The Ganga-Brahmaputra system

Fig. 9.5

Mean monthly discharge of the Upper Zambezi showing seasonality in flow. Drawn from data in Table 15.2 in Moore, et al., 2007 the floodplains and marshes along the Zambezi and presumably has altered their ecological conditions. These shallow-water, low-gradient, natural sediment traps are important wildlife refuges, including fish, avifauna and mammals that require a wetland habitat such as hippopotamus, and lechwe and sitatunga antelopes (Moore et al., 2007). These also could be important for farming and settlements as in the Barotse floodplain. Large floods periodically occur on the Zambezi. Moore et al. (2007) referred to a 100year flood in 1958 with approximately 9000 m3s−1 of water passing over the Victoria Falls. This is three times the volume of the biggest monthly discharge (April) over the falls. Large floods have occurred in this century in 2000, 2003 and 2009, and evidently are expected to occur periodically on this seasonal river. A cyclical pattern of wet and dry years on the Zambezi has also been suggested by McCarthy et al. (2000).

9.5╇ The Ganga–Brahmaputra system The two rivers rise on the opposite sides of the Himalaya Mountains and travel a couple of thousand kilometres before they meet and build a huge delta into the Bay of Bengal.



156

Large rivers in the tropics

Fig. 9.6

The Ganga Basin:€physiographic divisions and major rivers. Megafans of the large tributaries shaded. (b) after Blum, 2007

Together they drain over a million square kilometres, and discharge the fourth highest Â�volume of water (970 × 108 km3) and second highest average annual suspended sediment discharge (about 1000 × 106 t) to the sea. These numbers are reached because the rivers drain a high and tectonically active mountain and intense monsoon rain falls over their basins. The rivers are separately discussed first, followed by an account of their delta Â�conjoint with that of the Meghna River (Fig. 9.6).

9.5.1╇ The physical setting of the Ganga River The main headwater of the Ganga, the Bhagirathi, starts from the Gangotri Glacier at Gaumukh at 3800 m. The river is called the Ganga after its confluence with the Alaknanada at Devprayag. The Ganga descends to 290 m in 300 km at Haridwar, where it leaves the Himalaya to enter the alluvial plain (Figs. 2.6 and 9.6). After descending the steep Himalayan slopes, the Ganga and its tributaries flow through the low-gradient extensive alluvial plains of the Himalayan foreland, a subsiding continental interior foreland setting.

157

The Ganga-Brahmaputra system

The south-flowing Ganga turns first southeast downstream from Haridwar and then east to flow down a wide alluvial plain built by the river and its tributaries. All along its course, it is joined from the north by large Himalayan tributaries with high discharge at intervals of hundreds of kilometres. These rivers and the Ganga have built huge fan-shaped alluvial deposits at the highland–lowland contact, known as megafans (see Chapter 13). Several of the Himalayan megafans have been described in detail, e.g. the Kosi (Gole and Chitale, 1966; Wells and Dorr, 1987; Singh et al., 1993). The tributaries from the south drain the old rocks of the northern edge of the Indian Peninsula. A number of these tributary streams flow into the Yamuna, the largest tributary of the Ganga. The Yamuna is also a Himalayan river that joins the Ganga at Allahabad. Nearly 59 per cent of the combined discharge of 130 × 109 m3 at the confluence is from the Yamuna (Das Gupta, 1984). At the eastern margin of the large alluvial plain of the Ganga valley, the river passes through a gap in the basaltic low hills of Rajmahal and within a short distance enters its delta. One of its two major distributaries, the Bhagirathi or the Hughli, flows directly south within India, collecting drainage from the northeastern corner of the Indian Peninsula. The other major distributary, the Padma, carries most of the discharge into Bangladesh where it meets with the Brahmaputra, and then the Meghna, another major river. The three major rivers have built the combined Ganga–Brahmaputra delta.

9.5.2╇ Hydrology The Ganga is a rainfed seasonal river, although it does receive some summer snowmelt from the Himalaya. More than 70 per cent of the annual rainfall, in certain locations about 80€per cent, comes from the wet monsoon between July and early October. The annual total decreases from east to west, from 1600 to 500 mm. The southwestern basin is also drier with an annual total between 500 and 700 mm. As expected, the rainfall rises on the Himalayan slopes, reaching 1500–2300 mm (Singh, I., 2007). The rain often falls intensely, and tropical storms, some reaching cyclonic status, periodically affect the lower basin leading to floods. The flow of the Ganga reflects both the seasonality of rainfall and the stepwise increase in discharge where the major tributaries such as the Yamuna, Gomati, Ghaghara, Gandak, Son and Kosi come in. Some increment is due to the baseflow that seeps in through the deep alluvial cover. According to Das Gupta (1984), 50 per cent of the annual rainfall enters the river as surface runoff, 30 per cent is lost by evaporation and 20 per cent seeps to the subsurface. During the dry season, part of this subsurface water flows through the high banks of the Ganga into its channel as baseflow. The mean discharge of the Ganga at Farakka before it divides into deltaic distributaries is 70 547 m3 s−1. About 60 per cent of this arrives from the Himalaya and the northern plains (Das Gupta, 1984).

9.5.3╇ Sediment As expected, the sediment load of the Ganga comes mostly from the tectonic Himalayan Mountains. Chemical weathering is not important in the Himalaya and the solution load of the river is low, being diluted even further when the seasonal discharge is high. The



158

Large rivers in the tropics

Fig. 9.7

Changes in grain size of the bed material of the Ganga from Haridwar to Ganga Sagar. The middle reach coarsening is due to the contribution from the southern rivers draining the Indian Peninsula. From Singh, 2007a. By permission of Wiley

suspended and bed load of the Ganga are very high; the suspended load being the second highest after the Amazon. Milliman and Syvitski (1992) have estimated the annual suspended load of the Ganga as 520 × 106 t. About 90 per cent of the sediment travels during the months of the wet monsoon (Singh, I., 2007). The bed load of a large river is difficult to measure but Wasson (2003) estimated that 600–2500 million tonnes of bed load reaches the delta each year. Most of the sediment comes from the Himalaya via the large tributaries. The source of the tributaries varies, as they may originate from the mountains, the foothills or the plains below the hills (Sinha and Friend, 1994). Rivers starting in the northern mountains and foothills contribute most of the sediment. The southern tributaries drain a craton and contribute coarse sediment. Figure 9.7 demonstrates the downstream change in bed material of the river from Haridwar at the foot of the Himalaya to Ganga Sagar where one of the distributary channels, the Hughli, flows into Bay of Bengal. Measured from bar samples, it indicates a general downstream fining characteristic interrupted by periodic coarsening of the bed from contributions by large tributaries (Singh, M., 1996). The bar sediment of the Ganga is dominated by sand, the mineralogy of which is essentially quartz with minor amounts of feldspars, micas and rock fragments. Material from the weathered source rocks undergoes further alteration when the grains form part of the floodplain alluvium and remain in storage between being transported in high flows.

9.5.4╇ The valley morphology The Ganga and its headwaters in the Himalaya flow in narrow, deep, almost gorge-like Â�valleys flanked by small discontinuous patches of floodplains and terraces. The rivers in the wide alluvial plain between the Himalaya and the peninsular craton are entrenched below the surface of the plain. The channel of the Ganga remains within a 10–25 km wide elongated lowland, which is bounded by several metres high alluvial cliffs, the lowland being referred to as the Ganga River Valley (Singh, I., 2007). The cliffs enclose

159

The Ganga-Brahmaputra system

the channel with braid bars and meander scars, the floodplain, terrace-like features and wetlands. The channel and the floodplain together could measure up to 3 km in width and large sand bars, kilometres in dimension, are common in the channel. The Ganga is confined within the valley and even the large floods rarely overtop the cliffs. The floodplain and terrace-like features, however, are periodically inundated, the frequency of which depends on their height above the channel. Gullies and small ravines are common, and abandoned channels and wetlands appear on top of the cliffs on the alluvial plain. Like the seasonal rivers discussed in Chapter 7, the channel of the Ganga becomes braided with multiple channels and huge kilometre-scale bars consisting of braid bars, lateral bars and point bars at low flow. The low flow effect is enhanced by the large-scale transfer of water into irrigation canals or to meet other demands. Point bars develop where the river displays a meandering pattern with local narrowing of the channel due to extensions of peninsular lineaments under alluvium. The meandering pattern also appears in high flows when nearly the entire channel is under water, submerging the mid-channel bars. The bars occur at several levels related to the frequency of their inundation. The higher bars are vegetated and usually farmed. During the dry period, sediment transfer is confined to the deeper sub-channels. During the wet monsoon, sediment travels across the entire channel width and occasionally also over the floodplain. At high flow, several metres of sediment is scoured from temporary storages on top of the floodplain (Shukla et€al., 1999). The general channel pattern remains the same, but the location and geometry of the bars vary over time. The river was mobile in the past, but currently its lateral shifting is limited to a scale of several kilometres within the high cliffs.

9.5.5╇ The setting of the Brahmaputra River The Brahmaputra rises from the southern slopes of the Kailash Mountain north of the main Himalayan Range, and flows eastward along the Indus–Tsangpo suture through the Tibetan Plateau on a relatively gentle gradient, about 0.001. In Tibet, it is known under various names, the most common of which is Tsangpo. After traversing a distance of 1200€ km on the Tibetan Plateau, the Tsangpo takes a large U-turn at 95°E in a 5075 m deep gorge around the Namche Barwa Peak of the eastern Himalaya, a part of the Himalaya known as the Eastern Syntaxis. This sudden bend and the deep gorge are fascinating but yet unresolved features of the regional geomorphology. The river then turns south to enter the Arunachal Pradesh of India as the Sihang or Dihang. The huge river in the gorge flows over a very steep gradient (0.03) with a very high velocity and is extremely turbulent. Further downstream it reaches the Assam Plains and, after merging with the Dibang and Lohit rivers flowing from the northeast, becomes the Brahmaputra, a wide, deep braided stream flowing west–west–south in the reversed direction of the Tsangpo. The Brahmaputra maintains this course until near Dhubri, where it turns south to enter Bangladesh. In Bangladesh, it is known as the Jamuna, and joins the Ganga at Arichaghat. From its source to the confluence with the Ganga, it is 2900 km long. Figure 9.8 maps part of its basin and the major tributaries.

160



Fig. 9.8

Large rivers in the tropics

The Brahmaputra River and its main tributaries upstream of the confluence with the Ganga. From Singh et€al., 2005. With permission from Elsevier

9.5.6╇ Hydrology After flowing through the dry Tibetan Plateau in the rain shadow of the Himalaya, the Brahmaputra passes through the extremely wet eastern Himalaya Mountains, collecting drainage from the wet southern slopes of the Himalaya and the hills of eastern India and India–Myanmar border. The annual precipitation in the upper basin ranges from about 300 mm in Tibet to about 5000 mm in the Eastern Syntaxis. Like the Ganga, the Brahmaputra’s discharge regime is seasonal, dependent on the wet southwest monsoon. In Tibet, snowmelt, groundwater and rainfall are about equal contributors to the Brahmaputra’s discharge. In India, the monsoon rainfall predominates, with about 70–80 per cent of the annual discharge of the Brahmaputra in India arriving between June and September. The bankfull discharge of the Brahmaputra is about 35 000 m3s−1 at Pandu near Guwahati (Fig. 9.9). The mean annual flood is about 50 000 m3s−1, and the 25-year flood about 60 000 m3s−1. Large discharges arrive in the Brahmaputra via the Himalayan tributaries to fill the river and inundate its floodplain. Annual floods are expected on the Brahmaputra, given the seasonal rainfall, rapid change in relief to a very flat gradient, obstruction of the channel by many sand bars and the local

161

The Ganga-Brahmaputra system



Fig. 9.9

The annual peak discharge of the Brahmaputra River at Pandu, near Guwahati, India. Overbank flooding is common. From Singh, 2007b. By permission of Wiley



Fig. 9.10

MODIS/NASA image of the Brahmaputra Valley and the Ganga–Brahmaputra Delta. See also colour plate section narrowing of the Assam Plains between the Himalaya Mountains and the Shillong Plateau (Fig. 9.10). Floods on the Brahmaputra also occur due to seismic events. Massive earthquakes of 1897 and 1950, both with magnitudes of 8.7, partly blocked the river, extensively flooding the Assam plains.

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Large rivers in the tropics

9.5.7╇ Sediment A huge amount of sediment and solute is discharged by the Brahmaputra:€1000 million tonnes of clastic sediment and 100 million tonnes of dissolved matter (Singh, S. K., 2007 and references therein). This, however, is an estimate, as accurately determining the sediment discharge in the high flow of a large turbulent river like the Brahmaputra is nearimpossible. Bed load could be an important component of sediment transfer in a river like the Brahmaputra (Galy and France-Lanord, 2001). S. K. Singh (2007) has commented that given the scale of the runoff and the lithology of the eastern Himalaya, both physical and chemical erosion rates should be higher for the Brahmaputra Basin than that of the Ganga. Galy and France-Lanord (2001) determined that the total erosion in the Brahmaputra Basin is about 1.5–2 times higher than that in the Ganga. Isotopic compositions of the sediment of the Brahmaputra indicate that parts of the basin in Tibet, the Eastern Syntaxis (around the Namche Barwa Peak), the hills at the eastern end of the basin and the Himalayan slopes contribute 5, 45, 10 and 40 per cent of the total sediment, respectively (Singh and France-Lanord, 2002). The Eastern Syntaxis Zone only occupies 4 per cent of the basin, but according to S. K. Singh (2007) the erosion rate in this zone is 14 mm per year, among the highest in the world. This is an area of tectonics, high relief and intense rainfall. It has been surmised that such a rate of erosion leads to high rates of isostatic rebound, which is responsible for the very high peaks, like Namche Barwa (7750 m) and Gyala Peri (7150 m), in this area.

9.5.8╇ The channel of the Brahmaputra The Brahmaputra displays a surprisingly wide range of variations in its channel and bedforms. Unlike most rivers, its steepest slope is not at the headwaters in Tibet but in the middle of its 2900 km length where it crosses the Eastern Syntaxis. The gradient of the river is variable on the Tibetan Plateau but averages round 0.001, interrupted by a number of knickpoints. Downstream the river has cut a 5000 m gorge through the mountains of the Eastern Syntaxis where the gradient steepens to 0.03. Its slope drops to 0.0001 in the middle of the Assam Plains (Singh, S. K., 2007). The average channel of the river also widens to 8 km, but varies between 1 and 20 km in the tectonics-controlled lowland of the Assam with associated changes in gradient and bedforms. The upper Tsangpo in Tibet changes from a freely meandering channel to a braided one. It then flows in a single straight channel that becomes a winding channel flowing through gorges, before curving round the peak of Namche Barwa in a very deep gorge and emerging on the Assam Plains as a wide braided river. The highly braided channel of the Brahmaputra, which is full of sandbanks and high islands, has been described in detail by Goswami (1985) in Assam, and Coleman (1969) and Best et al. (2007) in Bangladesh. It is a wide mobile channel between narrow stable nodes (Coleman, 1969). Avulsions have been recorded. Huge braid bars, lateral bars and bars formed by backwaters at the mouths of tributaries are built by the material eroded from the mountains and at various levels within the channel. The Majuli Island in Assam, which is between two channels of

163



Fig. 9.11

The Ganga-Brahmaputra system

A bar in the low season Brahmaputra (locally known as the Jamuna) at Sirajgunj, Bangladesh. Note the high volume of sediment in episodic transport and storage and the sedimentary structures. Photograph: A. Gupta

the Brahmaputra, is 600 km2 in area, the largest river island in the world. The course of the river has changed in the past, following earthquakes in this tectonically active basin (Goswami, 1985). Further downstream in Bangladesh, the river, known locally as the Jamuna, receives the large Himalayan tributary of Tista from the west. The Brahmaputra in Bangladesh has changed course in the past by avulsion, possibly following earthquakes, indicating the mobility of large rivers in low flat terrain. The braidplain of the Jamuna is 11 km wide on average, the flow depth is about 5 m and its Brice Braiding Index is 4–6. The braidplain may widen to 15 km wide in floods, when scour depths of 40 m have been measured (Best et al., 2007 and references therein). It is a huge river flowing through a dynamic and highly variable alluvial environment (Fig. 9.11). Goswami (1985) gave the mean depth of the 31€July 1977 high flow at Pandu, Assam, India as 13.4 m. The river would be several kilometres wide at this stage and that measure of mean depth indicates several very deep channels for this wide braided river. Coleman (1969) described the short-term channel migration of the Brahmaputra in Bangladesh as drastic. The movement of major channels, modification of bed and bar forms, and bank slumping are conspicuous during high flows, and especially in floods. The huge combined delta of the Brahmaputra and the Ganga is the final product of the sediment eroded and transported from the Himalaya. Goodbred (2003) has described the history of the delta evolution and the changing nature of the source-to-sink sediment transfer and discharge variation for the Quaternary, including the passage of the sediment to the submerged Ganga Fan on the floor of Bay of Bengal as summarised in Chapter 16.

164



Fig. 9.12

Large rivers in the tropics

The Mekong and its basin. From Gupta, 2007. With permission from Elsevier

9.6╇ The Mekong The 4880 km long Mekong River drains a northwest–southeast trending pan-shaped basin in Southeast Asia to the South China Sea (Fig. 9.12). The handle of the pan is in China where the river starts at an elevation of 5000 m on the Tibetan Plateau. The river then flows in a partly confined channel through the mountainous terrain of southwestern China, Myanmar, Lao PDR and Thailand to the alluvial lowlands of Cambodia. Its delta stretches from Cambodia

165

The Mekong

to Viet Nam. The Mekong is a remarkably structure-guided river for most of its course; it is strongly seasonal in nature. The 785 000 km2 basin is primarily rural with a generally low density of population. It has been significantly affected anthropogenically only since 1993, when the first of a planned cascade of dams was closed across its upper course in China.

9.6.1╇ The physical setting The river runs on rock through narrow valleys in mountainous regions for the first 3000€km of its course. The next 1000 km are on mixed rock and alluvium, although the valley is almost always enclosed by neighbouring mountains. Only for the last 600 km of its course, can this huge river move freely through a 500-km wide alluvial lowland that ends in a large delta. The geology of the Mekong Basin is patchily known. Apparently, as the Indian Plate collided with the Eurasian Plate to create the Himalaya Mountains, part of the stress was transferred southeast to open up river valleys such as the Mekong (Tapponier et al., 1982). The upper basin in China is only a valley cut into granitic and sedimentary rocks, several kilometres wide. The river flows over Palaeozoic and Mesozoic rocks and local igneous intrusions downstream, in places along faults. The Mekong flows in a narrow valley with steep side slopes on rock; the water flow is steep and turbulent; and the channel is partly filled with sediment, probably from slope failures on the steep valley slopes. Gorges, terraces and rapids are common. South of China in the mountainous terrain of Lao PDR and north Thailand, the river flows through a landscape of narrow, sharp-crested ridges separated by deep valleys, with a local relief between 500 and 1000 m. Near Vientiane, a shallow layer of Quaternary alluvium overlies the rocks. Downstream of the alluvium, the river continues on rock and then a mixed bed of rock and alluvium until it reaches northern Cambodia, where it starts to flow over an alluvium of variable thickness, the thickness increasing downriver (Gupta, 2007b). The Annamite Mountains between Lao PDR and Viet Nam form the higher and steeper divide to the east, with peaks reaching to 2000 m. The 200–500 m Korat Highlands of Thailand form a lower but steep divide to the west. South of the Korat Plateau the western divide is formed by the small but steep Cardamom and Elephant hills that reach the sea abruptly at the northwestern corner of the Mekong Delta. The basin is rugged, but the elevation of the valley floor drops below 200 m about 1500 km from the sea. The river, however, is still structure-controlled, and only becomes free to move laterally 600 km from the sea in the Cambodian plain. This plain rapidly widens to about 500 km between the two divides and has a western extension where the regional drainage is collected in the large lake of Tonlé Sap. This lake is connected to the Mekong by the Tonlé Sap River, whose direction of flow reverses seasonally. In the wet season, water from the Mekong flows down the river and extends the lake. In the dry season the lake drains via the link river to the Mekong and shrinks in size. The Tonlé Sap River joins the Mekong near Phnom Penh, where about 330 km from the sea, the first deltaic distributary, the Bassac, splits from the Mekong marking the apex of a very large delta.

166

Large rivers in the tropics

9.6.2╇ Hydrology The annual average basin precipitation is 1672 mm (MRC, 1997) but it varies geographically. The annual total drops below 1000 mm in China in the north and in the Korat Highlands, Thailand in the west. To the east, over the northern and eastern basin, especially on the slopes of the Annamite Mountains, the annual rainfall is 2000–4000€mm. The strongly seasonal rain comes from the southwestern monsoon, with 85–90 per cent falling between June and October. This is reflected in the flow regime of the Mekong. The river rises a little in May following the summer snowmelt over the Tibetan Plateau and mountains, but 80 per cent of the discharge occurs between June and November (Fig.€7.15); 20–30 per cent may arrive in a single month. Large floods tend to occur late in the wet season, tailing off very slowly. Large floods, that inundate thousands of square kilometres in the lower basin on the plains of Cambodia late in the wet season, are often triggered by the arrival of tropical storms over the Annamite Mountains when the Mekong is already high. The tributaries, especially the smaller hilly ones, are extremely seasonal and rise sharply after rainfall. A disproportionately high amount of water arrives from the steep northern hills of Lao PDR and the northern Annamite Mountains via the Nam Ngum and Nam Theun systems and the southern Annamite slopes via the large tributaries of the Kong, San and Srepok (Gupta, 2007b). These are the major source areas for both the normal discharge and floods of the Mekong.

9.6.3╇ Sediment Sediment data on the Mekong are not of the same quality as the water discharge, but certain generalisations can be made. Significant sediment contribution is from the narrow, steep part of the basin in China, the hills of northern Lao PDR and the Annamite mountains. The high difference in river stage (30 m or more) between the wet and dry seasons results in a wide difference in shear stress, unit stream power, channel erosion and sediment transport. With the falling stage, sand and gravel accumulate around cross-channel protrusions in the rock and form bars, fill depressions in the rock bed of the river and form insets against vertical rock banks from massive episodic beds to plasters over the rock in small quantity. Material ranging up to boulders is exposed on the river bars during low flow, indicating the high competence of the Mekong in floods. The Mekong River Commission (1997) mentions 75–85 × 106 t arriving from the part of the basin in China and 150–170 × 106 t being discharged into the South China Sea annually. This is comparable to Mead’s estimate of 160 × 106 t and Ta et al’s figure of 144€± 36 × 106 t derived from past sedimentation rates in the delta (Meade, 1996; Ta et al., 2002). But the Mekong carries far less sediment than other large rivers of South and East Asia, such as the Ganga or the Changjiang. Sediment transport is episodic in the Mekong, mostly carried out during the high flows of the wet monsoon when the water depth may rise by 10–30 m. Most of the sediment discharge in the South China Sea thus happens later in the wet season, with the discharged sediment drifting west across the face of the delta (Gupta et al., 2006).

167

The Mekong

9.6.4╇ The channel of the Mekong The morphology of the channel of the Mekong changes several times along its course (Fig. 9.12). As expected, the river on rock is entirely different from the river on alluvium. The upper Mekong in China is a structure-guided river flowing over a very steep gradient, nearly 0.002. Straight sections are joined by short sharp bends in rugged gorges. The channel of the Mekong here is steep, straight and laden with coarse sediment contributed by short steep tributaries and mass failures on steep side slopes. Up to northern Cambodia, the Mekong is essentially a rocky conduit, frequently fault-guided, and with very little option to shift its course in the narrow deep valley. Accommodation space is limited in the valley and the general direction of transport for both water and sediment is downstream. The river, however, is not the same morphologically and Gupta and Liew (2007) identified eight different sections in the 2000 km of the river south of China. In cross-section, the river is either trapezoidal in shape or has a deep inner channel bounded by rock benches. Scour pools and rock protrusions occur on top of the rock benches and within the inner channel. The channel is marked in wider reaches by a profusion of tens of metres high cross-channel rock ribs, isolated transverse rock piles and lines of rapids. Local relief inside the channel of the Mekong controls sediment accumulation and coarse sediment is visible during low flows, forming bars on rock benches and deposited against rock ribs. The banks are in rock with insets in silt and sand. The channel pattern is straight with sharp bends but a reach with six entrenched meanders in rock also occurs (see Gupta, 2007b:€Fig. 20.10). Near Vientiane, where the Mekong flows across an area of Quaternary alluvium, the channel shallows and widens and a braided pattern develops during the dry season. This is a near-straight part of the river with elongated rock-core islands located askew inside the channel. Downstream of the Vientiane alluvium, the river is on rock with deep scour pools and entrenched U-bends (Fig. 7.8). Rock-cored alluvial islands are the common inchannel depositional features, where sediment starts to accumulate against a high feature on the rocky bed. In southern Lao PDR, the Mekong travels over a 50 km reach where it is anastomosed in rock before flowing over a series of rapids and waterfalls into the alluvial lowlands of Cambodia. The river is still structure-controlled as reflected by its four 45–50 km long straight reaches joined by right-angled bends. The Mekong flows freely in alluvium for the next 50 km, overflowing into its wide floodplain during the wet season and more strikingly during times of high-magnitude floods. Beyond is the delta of the Mekong, one of the largest in the world. The subaerial delta can be divided into an upper and lower part dominated by fluvial and marine processes, respectively. The upper delta is an assemblage of levee-bounded tidal channels and backswamps, whereas the lower delta is characterised by a series of parallel beach ridges and inter-ridge swamps and flat plains (Nguyen et al., 2000). The current river and its basin owe its present location and geometry to the collision of the Indian Plate with the Eurasian Plate and the strengthening of the monsoon system in the Early Holocene. This probably explains the narrow upper basin, the control on river morphology by the regional pattern of folds and faults, the straight river segments

Large rivers in the tropics

168

(a)

(b)



Fig. 9.13

The Mekong River. (a) River channel in rock near Luang Prabang, Lao PDR. Note the alluvial bank in the background and the sand deposited in the shelter of the rock to build a bar. (b) River channel in alluvium from the air near Savannakhet, Lao PDR. Note the large midchannel bar. Such bars commonly have a rock core around which alluvium is deposited. Photographs: A. Gupta

interrupted by sharp turns and U-bends, the inner channel and the rock ribs, and the restriction on Â�channel movement for about 95 per cent of the river length. Other regional big rivers such as the upper Changjiang, Sông Hóng (Red) and Salween exhibit similar features. The shifting of the coastline across the South China Sea during the Quaternary also affected the Mekong’s lower course. The current Mekong could possibly be an integration of several different river systems due to the formation of the Himalaya and the sea-level change. The Mekong certainly is an unusual river, both in its morphology and behaviour (Fig. 9.13).

169

Questions

9.7╇ The importance of major tropical rivers Large rivers operate as huge systems for transferring water and sediment to the oceans. Their nature and behaviour determine the morphology of their drainage basin to a large extent as the entire drainage network is connected to these massive conduits. They are not only important in the humid tropics with large rainfall but their passage, as in the case of the Nile or the Niger, is also crucial for dry regions. It is therefore essential to study these massive river systems. The majority of large tropical rivers have been interrelated with human habitation for a long time because of water, fertile fine-grained sediment, extensive floodplains of low relief and ease of navigation. The floodplains and deltas have been anthropogenically modified for a long time in human history and several of these rivers are now impounded and altered from their natural form and function. An understanding of large rivers is therefore essential for the sustainable development of their drainage basins, which cover a significant part of the tropical land and ocean.

Questions 1. What is a large river? 2. Why does the Amazon have a very low stream power on average in spite of its huge size? 3. Describe sediment transfer and storage in the valley of the Amazon. Is it all in the downstream direction? 4. The Amazon is a huge river and so are many of its tributaries. Do they still exhibit �evidence of control by geological structure? 5. Does the channel area of the Zambezi increase uniformly with its length? 6. Describe the change in grain size of the bed material of the Ganga, using Figure€9.7. 7. Explain the form and function of the Brahmaputra River in India and Bangladesh from the MODIS image of Figure 9.10. 8. Does the Mekong receive water and sediment uniformly from all parts of its drainage basin? If not, identify and explain the major sources. 9. In what way is the Mekong in rock different from the Mekong in alluvium?

10

The tropical coasts

Stately Spanish galleon coming from the Isthmus, Dipping through the Tropics by the palm-green shores John Masefield

10.1╇ Introduction The coast is where geomorphic processes operate in combination. Waves and currents of the sea play the dominant role but coastal forms are also shaped by fluvial processes, mass movements and wind action. For example, the face of a coastal cliff is denuded by weathering, rainwash, rill action and mass movements, while the sea erodes its base (Fig.€10.1). Sediment derived from a coastal cliff is transported along the coast, as is sediment brought down by rivers flowing into the sea and material moved landwards by waves from the sea floor. The term 'coastal processes' includes all activities that take place at or near the sea. Coastal processes, however, are restricted to a narrow vertical range primarily determined by the sea level, and confined within the extreme tidal limits. The reach of the sea,



Fig. 10.1 170

Coastal cliff and shore platform, Australia. Note the effect of sedimentary structure in cliff erosion and the operation of multiple processes on the cliff shown. Photograph: A. Gupta

Introduction

171

however, extends in huge storms, e.g. tropical cyclones, or gigantic waves generated by tsunamis, but these are low-frequency events. Certain distinctive characteristics of the coastal zone are:

1.╇Coastal forms occur within a vertical range defined by the height of the highest cliff or dune back of the coast and the depth of the lowest tidal limit. 2.╇The horizontal range of the coast is defined by the local slope. Steep offshore slopes give rise to narrow coasts, gentle slopes to wide ones. 3.╇This vertical range and the absolute elevation of its boundaries vary over time as the sea level fluctuates due to climate change or tectonic activities. 4.╇Coastal forms change swiftly, being different between the wet and dry seasons of the year. 5.╇The sediment in the coastal zone may be derived from both land and water and contributed to by a number of geomorphic processes. 6.╇A tropical coast often carries a rich biological diversity that creates a special environment:€mangroves, salt flats, coral reefs. 7.╇ Many coastal stretches are heavily populated.

The words shore and coast are used both interchangeably and definitively, the usage depending on the context. Box 10.1 includes a list of definitions, including that of shore and coast.

10.2╇ Types of coast Inman and Nordstrom (1971) proposed a classification of the coasts of the world, based on the premise that coasts are primarily controlled by plate tectonics. This classification is based primarily on large coastal forms with linear dimensions at the scale of 103 km, offshore–onshore (coastal plains to continental shelves) measurements at 102 km and vertical dimensions of about 10 km from the ocean floor to the top of the coastal mountains. According to their classification system, first-order features occur at this scale. They envisaged coasts primarily as (1) collision-edge coasts (occurring at the active plate margin of a continent); (2) trailing-edge coasts (on the passive side of a continental plate); and (3)€marginal sea coasts (where island arcs separate seas from the open ocean). Their coastal classification is hierarchical. Second-order features are river deltas, sand dunes or glaciated valleys. Smaller coastal forms such as beaches or bars are recognised as the third-order features. Apart from coastal tectonics, Inman and Nordstrom (1971) also divided the coasts of the world as wave-eroded (45 per cent), glaciated (36 per cent), and wave-deposited (10.7€ per€ cent). Rivers are primarily responsible for the remainder. Life forms such as mangroves and corals thrive on the hot and humid coasts. Large salt flats tend to develop along arid coasts. Large deltas form in protected physical environments, provided enough sediment is available. Mangroves, salt flats and coral reefs are discussed in detail later in this chapter and deltas in Chapter 11. In the tropics we should expect coastal processes to

The tropical coasts

172



Box 10.1

An introduction to coastal terminology Shore: The part of the land–sea interface bounded by the extreme high and low tide levels. Coast: Coast is shore plus the cliff or dune at the landward margin of the shore and, in some locations, a small sector of land behind the cliff or the dune. A coast is wider than the shore, but both are usually limited to several kilometres, generally much less. The shore is further divided according to tidal levels. Foreshore: It is the intertidal zone between the mean high and mean low tide levels. It is partially exposed or covered with water, depending on the tidal conditions. Backshore: Backshore is the zone landward of foreshore, inundated only during large storms; it includes higher parts of the beach backed by cliffs or dunes. Inshore: This part of the shore lies seaward of the foreshore and is commonly submerged. Offshore: This is the furthest seaward zone, beyond breaking waves. See Figure 10.2 for illustration.

Fig. 10.2â•… Coast and shore

be defined by tectonics, but modified by waves, currents, life forms or the presence of a huge amount of river sediment in appropriate locations.

10.3╇ Moving water:€tides, waves and currents Tides, waves and currents control the movement of water near coasts. It is therefore necessary to introduce these phenomena prior to discussing coastal forms and processes. An in-depth discussion is found in a book on beach processes by Komar (1998).

10.3.1╇ Tides The tidal range controls the width of coast over which wave action operates. Strong tides not only create a wide coast but also result in water flowing fast in and out of bays and

173

Moving water:€tides, waves and currents

lagoons, keeping such inlets open and carrying sediment out to the open sea and up inland rivers. Tidal heights vary in different parts of the coasts of the world. Davies (1964) used tidal heights to classify shores of the world into three classes: • macrotidal (spring tidal range above 4 m) • mesotidal (spring tidal range 2–4 m) • microtidal (spring tidal range below 2 m). Tides moving inland negotiate progressively narrower and shallower passages along rivers and the inlets of coastal lakes. This leads to the front edge of the tide oversteepening and resembling a wave front. This is the tidal bore. Off the Ganga–Brahmaputra Delta, where the range is macrotidal, high tidal bores are formed as the rising tide moves into the delta through funnel-shaped estuaries. The highest tidal bores probably occur on the Amazon where they resemble upstream-moving waterfalls, 5 m in height above the level of the river and advancing at a rate of about 10 ms-1 (Komar, 1998). In large storms, the sea level is raised by wind pressure and huge waves approach the coast. This phenomenon is known as storm surge. Storm surges are common on many tropical coasts, especially those affected by tropical cyclones. Storm surges are extremely effective in eroding and transporting coastal material and are destructive of life and property. For example, the 29 October 1999 super cyclone from the Bay of Bengal raised a 7–8 m storm surge that caused widespread destruction on the coast of Orissa, India.

10.3.2╇ Waves Waves are oscillatory movements that cause the water surface to rise and fall. A number of terms are used to describe the physical properties of waves (Box 10.2). Waves are generated by wind in the open ocean and travel for considerable distance to reach land where they steepen and break against the coast, resulting in erosion, sediment transfer and deposition. As the energy of waves is derived from the wind blowing on the surface of the sea, the following properties simultaneously determine the size of the waves and their function: • wind velocity • wind direction • fetch (the distance over which a wave-generated wind blows). The first step in the generation of waves usually happens in the open sea in an area of strong winds. This area where waves start is known as the sea and its surface is complex with waves of varying sizes and shapes. Some old waves disappear and are replaced by new waves, and wave crests interfere with each other. Waves tend to move out of this source area and travel downwind, and in the process the complexity is resolved into a number of separate wave groups (also known as wave trains) called swells. The process of transformation from sea to swell is known as wave dispersion. Swells tend to have waves of similar heights and wavelengths and may travel huge distances across oceans (Fig. 10.4).

The tropical coasts

174



Box 10.2

Properties of waves wave crest: highest point of a wave wave trough: lowest point of a wave wave height (H): vertical distance between a wave crest and the following or preceding wave trough wave length (L): horizontal distance between two similar points (e.g. crests) on successive waves �measured perpendicular to wave height wave period (T): time taken by two successive wave crests to pass a fixed point wave set up: average water level raised by a wave above the stillwater level of the sea wave steepness: ratio of wave height to wave length (H/L) wave velocity (C): travel speed of wave motion in a specific direction (L/T) See Figure 10.3 for illustration.

Fig. 10.3â•… Wave properties



Fig. 10.4

Wave formation A number of relationships exist between wave properties

L = gT2/2π C = gT/2π, where g is acceleration due to gravity and the other symbols are as in Box 10.2.

(10.1) (10.2)

Moving water:€tides, waves and currents

175



Fig. 10.5

Wave refraction Such relationships indicate that waves with similar periods travel together and a confused sea is transformed into a regular swell. Longer and faster waves outdistance themselves from shorter and slower ones over time. In a swell, individual waves rise from the back of the group, travel through the group and disappear towards the front. Thus individual waves disappear but the group characteristic remains. Swells are wind-driven and therefore they tend to travel in deterministic paths in long and narrow bands across the open water. Very little energy is lost once the swell is formed and running. The period of these waves is usually shorter than 20 seconds (Komar, 1998). Waveforms and dynamics change as waves approach a coast. The shallowing and irregularity of the coastline affects the wave trains in three ways:€refraction, reflection and diffraction. A line of waves refract as they approach the coastline. The water shallows and the base of the waves slows down progressively. On a coast with alternate headlands and bays, the section of the line of waves facing the headlands feels the effect of shallowness early and slows down. The section of the wave facing the bay is still at a distance and therefore continues to move fast into the bay. This bends a straight wave crest until it becomes parallel to the coastline. This is wave refraction (Fig. 10.5). The bending of the wave results in part of the wave converging on the headlands and part of the wave diverging into the bays. This results in the concentration of wave energy per unit of wavelength on the headlands and the dissipation of energy in the bay. Thus high waves and erosion occur on headlands, whereas low waves and deposition occur in bays. When waves approach the coast through deeper water, such as in front of a cliff, a steeply inclined beach or a seawall, waves are reflected back. If both the incoming waves and the reflected waves are of similar dimensions, then standing waves or clapotis may develop off the coast. If the reflected waves are of dissimilar dimensions, then partial development of clapotis is seen. This kind of periodic oscillation is also known as seiches, a term originally used for enclosed water bodies such as lakes, but it is now also used for this phenomenon in harbours and open waters. Wave diffraction occurs when waves approach a natural or engineered barrier, such as an island or a breakwater, at an angle. This leads to a shadow zone in the lee of the barrier where the waves are propagated with a change in direction of the crests (Fig. 10.6). A wave approaching the coast moves into shallower water and at the depth where the wave height is about half the wave depth, the base of the wave begins to be impeded and

The tropical coasts

176

Wave crest



Fig. 10.6

Wave diffraction in the lee of an island



Fig. 10.7

Transformation of wave orbital motion near coast

a differential develops between the top and bottom of the wave. This difference in velocity increases as the water shallows and ultimately the waveform becomes so steep that it breaks. The point where the wave breaks is determined by the ratio (γ) between wave height (H) and water depth (h), γ = H/h. Waves of different heights therefore break at different distances from the coast, giving rise to several lines of breakers offshore. A wave of a given height breaks closer to the shore on a steep coast than on a gentle one. The nature of a wave changes as it moves from deep water to shallower depths near the coast. Waves are oscillatory in deep water, where the waveform moves over the surface but individual particles of water instead of travelling undergo an orbital motion (Fig.€10.7). The orbit gets smaller with depth and stops beyond a threshold. The circular orbit becomes elliptical as the wave moves into shallow water, as its base is impeded, and finally the impediment causes the wave to break and the water at the crest moves forward in a linear fashion. These are now translatory waves where the water particles move forward instead of

Moving water:€tides, waves and currents

177



Fig. 10.8

Types of breakers

oscillating, and impact the shore with the possibility of erosion. Instead of the low, rounded swells of deep water, the waveform is now peaked with flat troughs between waves. At the same time, the waves may undergo refraction, reflection or diffraction. The velocity€(V) and wavelength (L) of a translator wave are given by V = √(gh) and L =€ T√(gh). The Â�notations are the same as before. The incoming wave finally breaks on the shore, giving rise to one of the three types of breaker:€spilling, plunging or surging. The particular type formed depends on the local conditions. Spilling breakers occur on beaches with low gradient. The top of the wave becomes unstable and flows down the front of the waves with bubbles and foam, while the wave travels for several wavelengths before breaking. Plunging breakers have near vertical wave fronts that curl downwards and impact the base of the wave. These are therefore violent breakers and, unlike the other two types, erode an impressive amount of sediment. Surging breakers are seen on very steep beaches. The wave crest does not collapse; the wave top curls over, but it is the base of the wave that travels up the beach causing the wave to collapse (Fig.€10.8). Different breakers may operate at different locations on a given shore, and the types may also change over time. Plunging breakers erode material out of the shore. The other two types tend to deposit. The area where waves break is termed the breaker zone. The shallow front part of the breaker zone is the surf zone. Water from breaking waves runs up the surf zone as shallow sheets known as swash. The swash travels at high velocity and falls in the upper-flow regime. The deposited sand on the beach therefore has a sub-horizontal lamination. As swash runs out of energy, it flows backwards down the slope, known as the backwash, leaving the beach temporarily exposed to the atmosphere. The zone over which swash and backwash operates moves up and down the shore, depending on the height of the tide and storm surges.

10.3.3╇ Currents Waves move water at an angle to the shore. Nearshore water also flows in a direction parallel or sub-parallel to the coast as a current. Two main types of current can be identified: 1. Longshore currents 2. A circulatory cell system.

The tropical coasts

178



Fig. 10.9

Rip currents Longshore currents are created by waves coming at an angle to the shore. A wave coming in at an angle to the shore may meet the retreating water from the previous wave which returns downslope and orthogonal to the shore, and move it obliquely further down the shoreline. In the process water and sediment may move longitudinally along the shore. The velocity of longitudinal currents usually is 0.3–0.6 ms−1, although higher values have been reported and a pulsating nature of velocity is common (Carter, 1988). A considerable amount of sediment may move along the beach in this fashion, and longitudinal currents tend to move sediment from headlands or river mouths to downcurrent bays to form the depositional forms described later. These currents are visible at low tide when the surf zone topography is visible and water can be seen flowing through troughs and channels close to the shore. A circulation system of water can be set up in a different way near the shore. As waves move shoreward, the heights of waves change along the crest causing water to flow along the shore from the highest towards the lowest breakers. This movement is often helped by the presence of troughs running parallel to the shore in the surf zone. On reaching the position of the lowest breakers, water flows seaward as a strong current in a narrow zone through the surf. This is the rip current, which is very dangerous to swimmers. Velocities approaching 2 ms−1 have been observed (Sonu, 1972), indicating that rip currents are capable of moving coarse material. Water is returned to sea by rip currents and then moved shoreward in waves, completing the circulation (Fig. 10.9). This circulatory system,

179

Rocky coasts

especially the location of rip currents, is enhanced by suitable bottom topography of the surf zone, especially if rips are located over a trough. Rips, however, also occur off straight beaches with little topographic variations in the surf zone. Tides, waves and coastal currents erode part of the shore and deposit sediment �elsewhere. A coast exhibits evidence of erosion and deposition simultaneously, the exact location being dependent on the local environment, such as the position of headlands and bays, location of river mouths, offshore gradient and presence or absence of rocks on the coast. It is convenient to examine the rocky and non-rocky coasts separately, as their forms and processes differ. Usually, however, a coastline includes both, and the forms and processes transform from one type to the other.

10.4╇ Rocky coasts Cliffs marked by rock-cut platforms at their base characterise rocky coasts (Fig. 10.10). A coastal cliff is a steep slope whose base is eroded by wave action. The gradient of the cliff and its appearance are determined by its geology and the nature of the waves that break at its base. Certain properties can be highlighted. 1. Cliffs located at tectonically active coasts as along convergent or transform plate �margins are usually steep and pounded by large waves at their base because of steep offshore gradients. These cliffs are referred to as plunging cliffs. 2. Cliffs reflect their lithology, especially in tectonically passive areas. For example, cliffs in granite or basalt are steep but those in soft, less indurated rocks tend to be low, gentle and actively eroded.

Fig. 10.10

Base of a rocky cliff, shore platform and stacks exposed at low tide, east coast of western Malaysia, facing the South China Sea. Photograph: A. Gupta

180

The tropical coasts

3. Structural variations are important. Waves tend to erode cliffs along lines of weakness:€bedding planes, faults, joints. Cliffs in seaward-dipping sedimentary rocks tend to break up in blocks which slide to the base. Cliffs in landward-dipping rocks are steeper and more stable. If hard and soft rocks are bedded alternately, cliffs are often unstable. 4. If the strike of the rocks is at an angle to the coast, a headland-and-bay coast is formed, with hard rocks at headlands and easily eroded softer varieties behind bays. If the strike is parallel to the coast then hard rocks form a long line of cliffs, and bays are formed only where such rocks are breached. The sea then can erode softer rocks at the back to form near-circular bays with a narrow opening through the hard rock. Waves erode the base of the cliffs, but most of the cliff is usually beyond the reach of waves. The upper part of cliffs is therefore denuded by weathering, slope failures and rill action. A coastal cliff is therefore formed by a combination of geomorphic processes whose operations are influenced by the nature of regional and local geology. As a result, many cliffs have an upper overhanging part as they are efficiently undercut by waves. This is followed by collapse, and cliffs tend to retreat inland over time. The erosion of cliffs by wave action takes place by any of the following sub-processes or by a combination of multiple processes: 1. attrition:€disintegration of rock particles when waves bring them into forceful contact with a rock or sand surface 2. corrasion:€erosion of coastal bedrock, commonly by the steady grinding action of sand grains or rock fragments brought in by waves and coastal currents 3. corrosion:€dissolution of soluble coastal rock such as limestone by seawater 4. hydraulic action:€erosion by the waves themselves; caused by the water pressure of the pounding waves or by the pressure exerted by collapsing air bubbles trapped between sea and cliff (termed cavitation). Sand and rock fragments are major tools for erosion and they could be derived from the cliff itself. A coastal cliff therefore partially provides the tool for its own destruction. However, the accumulation of too much material at the base may save the cliff from further erosion. This may happen if material from slope failures on the upper part of the cliff falls to its base or a large part of the lower cliff collapses because it has been undercut by waves. Estimates of the annual retreat of coastal cliffs are usually very low and dependent on rock type. Annual rates of cliff erosion have been estimated by laboratory experiment as 10 m in volcanic ash, 10 cm–1 m in sedimentary rocks, 1 mm–1 cm for limestone, 1€mm in granite and even less in basalt (Carter, 1988; referring to Sunamura, 1983). As expected, the rates of erosion and steepness are indirectly related. Near-vertical cliffs are formed in basalt. A retreating sea cliff is replaced by a near-horizontal or slightly dipping rock platform, the level of which is a little below the high tide level and probably within a depth of 10 m (Bradley, 1958). This is known as both a wave-cut platform and a shore platform. The gentle slope and the limited vertical range of the wave action determine the width of shore platforms. On an active plate margin coast, however, the platform may drop off steeply or the cliff may plunge to depths without much of a shore platform. Waves move back and forth across the platform in a tidal action, both trimming it and enlarging the irregularities in the rock.

181

Fig. 10.11

Non-rocky coast

Sea arch at low tide, Rawa Island, Malaysia. From Kale and Gupta (2001), © Orient Blackswan Pvt Ltd, India 2001 The retreating cliff leaves vertical erosion remnants in many places, which are Â� distinctly visible on the shore platform, especially at low tide. A narrow rock spur may run down from the headland. If its nose is eroded at its base exposing a roofed passage, it is known as a sea arch (Fig. 10.11). A sea arch frequently stands isolated from the line of cliffs. When the roof of the arch collapses to form one or two separate vertical bodies of rock, stacks are formed. Over time, stacks are eroded by the sea and disappear. If soluble rocks like limestone are present, caves may develop into cliffs and solutional features on the shore platform (Viles and Spencer, 1995). Even non-soluble rocks like sandstone with well-developed zones of weakness may be eroded to form tunnels and vertical holes through which seawater spurts in high tide. These are known as blowholes. About 80 per cent of the coasts of the world are lined with cliffs (Emery and Kuhn, 1982), although their height and appearance vary. A very long stretch of the subductionderived coast of South America is backed by high cliffs, running south from Peru. The west coast of the Indian Peninsula, except the extreme north and south, is backed by coastal cliffs. The northern part of this long line of cliffs is in flood basalt and the southern part in gneissic material. Many mid-oceanic volcanic islands carry cliffs.

10.5╇ Non-rocky coast A non-rocky coast is usually formed of material deposited either along a straight coastal stretch or at the heads of bays in headland and bay type coasts. The material is sand, or in less common instances, gravel. The depositional forms could be higher than the mean sea level (beaches, spits, forelands, etc.) or submerged most of the time (bars).

182

The tropical coasts

10.5.1╇ Beach A beach is probably the most common and striking form of coastal deposition (Fig. 10.12). It is a cover of wave-deposited sand or gravel over rock or a pre-existing sediment backshore. The gentle surging breakers move sediment up the shore to build a beach. The top of the leading wave approaches the beach at an angle, curls, breaks and its impact on the shore carries a thin sheet of water upslope, depositing sediment. The water (generally with very little sediment, if any) then runs down the steepest slope, orthogonal to the beach, as backwash. Repeated swash and backwash thus cause water and sediment to drift parallel to the coast; the process is known as beach drifting. The result is the same as the longshore drifting mentioned earlier: water is transferred parallel to the coast and sediment is transferred with it, extending and building the beach. Beaches are generally made of sand which comes from three sources:€rivers, headlands and the sea floor. The available volume can be very large. The Ganga–Brahmaputra system, for example, annually brings about 1000 million tonnes of sediment to the Bay of Bengal. Large beaches, as expected, occur close to large sources of sand. The river sand can travel long distances. Sediment carried to the South Atlantic by the Amazon is dispersed northwestward along the northeast coast of South America. Much of this sand goes onshore within the first thousand kilometres of coastal wetlands and estuaries, but about half the amount, which Meade estimated as in the order of 100 million tonnes per year, reaches much further to build the beaches of the Orinoco Delta, 1600 km away (Meade, 2007). This of course is an extreme example; the journey of the sediment is usually much shorter. The composition of sand varies. Commonly, it is silica sand, but where the material is derived from a coastal limestone or a coral reef, the material is texturally sand but

Fig. 10.12

Beach and berm, east coast of Malay Peninsula. Note vegetation at the back of the berm. A spit extends to the right from the beach enclosing a lagoon. Photograph: A. Gupta

Non-rocky coast

183

Fig. 10.13

Ridges and runnels and bars compositionally calcareous. Such beaches are frequently dazzlingly white and found on coral islands of the Caribbean Sea and Pacific and Indian Oceans. In contrast, sand derived from volcanic rocks is dark. Pyroclastic material derived from the slopes of the volcano Merapi has built black sandy beaches and dunes at Parangtritis on the south coast of Java as discussed later in Chapter 13. If a beach receives sand from more than one source, then different parts of the beach may have different colours. If coarser material is available, the beach could be made of pebbles. The presence of calcium carbonate near the water table may lead to induration of sand as the water table fluctuates with tides or seasonal changes. The resulting consolidated material is known as beach rock. As expected, beach rock is common in a lime-rich environment such as the coasts of the Caribbean Sea. The alignment of the beach depends on the sediment supply and angle of the swash. With a high supply of sediment, longshore transport extends the beach in the direction of the drift. If sediment supply is low, the beach is extended following the angle of the swash. Beaches can be long and straight, stretching for kilometres or they could be curved round the head of a bay between headlands. Bay beaches may resemble the Greek letter zeta, and are known as zetaform or log-spiral beaches. They acquire this shape because the direction of swash is changed along the bay by wave refraction. Grains are sorted along beaches in the direction of drift. Very small curved beaches are infrequently found attached to headlands. These are appropriately called headland beaches. A rock cliff or a sand dune occurs at the back of the beach. A line of vegetation marks the landward end of the beach in rare cases when cliffs or dunes are not present (Fig. 10.12). Beaches slope towards the sea but not uniformly. The upper part of the beach is generally flat, called a berm. This ends at a scarplet, below which the lower part of the beach slopes down to the sea. The level of the berm is reached only by waves that are significantly high, and berm levels are often markers of heights of storm waves. The berm is thus a depositional form, modelled by sand left by the swash. Erosion by smaller waves gives it a vertical step (Fig. 10.12). Waves from large storms, however, tend to flatten the beach and recovery, with re-steepening of the beach surface, happens during inter-storm periods. Beaches, however, do occur with no berms or multiple berms. The lowest part of the beach, exposed only during low tide, exhibits lines of bars and troughs between bars. A sequence of small parallel bars and troughs occur nearest to the beach and are termed ridges and runnels (Fig. 10.13). Bigger bars occur in deeper waters, except where the offshore profile is too steep for bars to form. Lines of breakers often indicate the location of bars. Bars form as far into the sea as the limit of the wave base disturbs the sediment on the sea floor, roughly up to 10 m depth. Even bigger bars are formed by storms in deeper waters, which remain in position between storms. The geometry of bars varies. They could be parallel or transverse to the shore; some bars are even crescent-shaped.

184

The tropical coasts

Certain beaches display special characteristics. A series of cuspate forms may develop along the foreshore, consisting of uniformly spaced crescentic piles of sand or coarser material that taper towards the water. These are known as beach cusps. Parallel ridges of sand separated by linear depressions, called lagoons, may occur at the back of the beach. These are known as beach ridges, formed either from old spits which have been welded later to land, or as piled-up coarse sediment and shells by storm waves. Thus the presence of beach ridges indicates either a former high sea level or a stormy coast. Many beach ridges were formed in the last 5000 years, and are associated with the highest sea-stand after the last glaciation. Excellent examples of this occur along the south Aceh coast in Sumatra and the Kelantan and Trengganu coasts of the Malay Peninsula. Well-developed beach ridges on a marshy coast are called cheniers. A beach is an extremely ephemeral form and its size and shape change frequently in storms or between wet and dry seasons. For example, beaches in the monsoon tropics are eroded during the wet monsoon and rebuilt during the dry ones. Komar (1998) has observed that beaches are eroded and bars are built in storms. Beaches with berms return at other times when bars tend to cover less of the area of a beach.

10.5.2╇ Spit A spit is an elongated depositional form like a very low embankment that extends from land into the sea (Fig. 10.14). It is made of the same material as the beach and is built by longshore drifts carrying sediment. Longshore drifts tend to continue in the same direction even if the coast bends following the outline of a bay, and spits are extended straight into the bay as a continuation of the beach. A spit extending into open water is characterised by a welldeveloped beach on the seaward side and a partially enclosed body of shallow water on the other, known as a lagoon. When lagoons become filled with fine material and aquatic vegetation, they become tidal flats. Spits are common on the downcurrent side of river mouths, and in places they divert river mouths or even block them, except for a small tidal inlet between the sea and the river. Spits and their effect as barriers are common on the east coast of the Malay Peninsula, the coasts of the Bay of Bengal and coastal Kerala (India). Contact with a second set of longshore drift, running obliquely to the first set, curves the distal end of a spit. This type is called a recurved spit or hook, a name borrowed from the northeastern Atlantic coast of the United States where this feature is common. Recurved spits are also formed from wave refraction round the distal end of the spit. Multipronged recurved spits also occur, and the shallow water between prongs commonly changes to tidal mudflats. Tombolos are spits that join two islands or an island with the mainland. Spits are linear but triangular depositional features called forelands also are formed by the accretion of material on both sides of a spit, or from deposition in a sheltered location such as the lee of an island.

10.5.3╇ Barrier island A barrier island is an elongated depositional feature, which appears above water as a long line of sand separated from the mainland by a narrow and shallow body of water, called the

Non-rocky coast

185

Fig. 10.14

Spits, recurved spits and tombolos lagoon or bay. These are much bigger features, up to 100 km in length and 5 km in width, but they are low, commonly below 6 m in elevation (Ritter et al., 1995). As these are large features, people have settled on them, but the low elevation makes them, and anyone who lives on them, vulnerable to tropical storms and high waves. Their growth requires an assured supply of sand, limited wave energy, a low offshore gradient and a limited tidal range. Beaches tend to develop on the open side of a barrier island. Their tops carry low sand dunes and flat areas under grass and shrubs. Higher barrier islands may even have trees. Marshes and tidal flats occur between the island and the proper lagoon, and narrow

186

The tropical coasts

tidal inlets separate individual barrier islands, which may stretch one after another for a long distance parallel to the mainland. The inlets, however, are temporary and often shift or silt up following events such as a large storm. The origin of these islands is much debated, but they are generally seen as relict forms from a lower sea level during the Quaternary which have migrated landward with the rising sea level and been modified by current coastal processes.

10.6╇ Coastal sand dunes Wide sandy beaches are commonly backed by sand dunes. Two properties differentiate coastal dunes from the dunes in the desert discussed in Chapter 12. First, although coastal dunes are formed from wind-deposited sand, they are also found on humid coasts. In Southeast Asia, for example, coastal dunes occur in areas that receive 2000 mm of rainfall a year or more. Second, vegetation plays an important role by anchoring sand to develop coastal dunes. Wind blows sand inland from the beach. This implies that coastal dunes develop best in areas with a wide beach, a large supply of fine sand, a strong wind blowing onshore and the presence of low vegetation at the back of the beach. The width of the beach is usually determined by the tidal range; a high range exposes a wide stretch of beach at low tide. Coastal dunes, however, vary in their dimensions. In an earlier study of coastal dunes, Smith (1954) stated that these dunes vary in width from 30 m to several kilometres and in height from under 3 m to 600 m. Multiple lines of dunes may occur. The older dunes may be semi-indurated and immobilised from the precipitation of CaCO3 from circulating groundwater. The consolidated dune sand is called aeolinite. Smith (1954) provided a very useful classification of coastal dunes (Table 10.1).

10.7╇ Coastal tropics The general coastal characteristics and processes as discussed so far are the same in and out of the tropics. The tropical coasts, however, have certain properties which are added to the general ones. First, wetlands and offshore forms associated with climate and biological activities provide tropical coasts with a distinct appearance with the presence of mangroves, dry salt flats (sabkhas) and coral reefs. Second, certain tropical coasts, because of their location, are periodically modified by tropical storms. We deal first with the coastal wetlands of the tropics. Coastal wetlands appear as flat swampy areas behind bars or spits or as infilled depositional basins such as deltas and lagoons. These areas are characterised by very little relief, standing water, brackish water and depositions of fine-grained sediment. Tropical coastal wetlands are rich in life forms, and in many places prevent serious coastal erosion by

Coastal tropics

187

Table 10.1╇ Types of coastal dune Types

Description

Foredune

Piles of sand not more than 3 m in height, running parallel to the beach, immediately behind it. The form changes rapidly with a shift in the wind unless anchored by vegetation. Curved sand ridges that open towards the beach. They are formed when the middle part of a dune is removed by wind or people. A former elliptical mound of sand thus is transformed into two linear arms anchored by vegetation. Typical barchan (see Chapter 12) with the slip-off slope towards the sea and with horns pointing inland. Sand ridge parallel or at a low angle to the shore and �perpendicular to the direction of the prevailing wind. The side towards the beach is gentle but the one towards land is steep. Highly oblique or perpendicular ridges to the shore but parallel to the wind. They tend to be symmetrical in form. Hollows cut into a longitudinal dune with sand mounds developing in the lee of the large hollows. Accumulation of sand around obstacles, mostly vegetation.

Parabolic or U-shaped dune

Barchan or crescentic dune Transverse dune ridge

Longitudinal dune Blow-out Attached dune After Smith, 1954

�functioning as a barrier to waves. Mangrove swamps are associated with the humid tropics, although some occur in arid areas. Arid coasts are characterised by salt flats. Although deltas occur on coasts, they are formed by a combination of multiple processes (fluvial, coastal, wind), and are discussed separately in Chapter 11. Deltas are the most important type of coastal wetland.

10.7.1╇ Mangrove swamps Mangroves form a vegetation community, which extends from about 32°N to 38°S, and thus fringe a very large part of the tropical coast, especially the humid variety. The Â�vegetation community commonly occurs in the zone between mean sea and mean high spring tide levels (Fig. 10.15). Mangroves develop best in sheltered areas where a large supply of silt from rivers as well as sand from the sea is available. However, this vegetation may grow on various foundations:€rocky coast, sandy coast, sheltered mudflats. Spatially they extend from saline conditions next to the sea to the brackish waters of inland river mouths. Plants embedded in mud tend to slow down the flow of water and a basket-like pattern of mangrove roots assist deposition. Mangroves therefore tend to accelerate sedimentation and build up coastal wetlands, especially in shallow sheltered areas such as tide-dominated deltas, river mouths and sheltered coasts behind a line of coral reefs. Fine-grained sedimentation associated with mangroves occurs mostly by flocculation in tidal waters. Clay-size particles (flocs) brought down by rivers as suspended load carry a positive charge on their

188

Fig. 10.15

The tropical coasts

Mangroves, South Andaman Island, India. Note the breathing root systems in the foreground. Photograph: A. Gupta

outer surface which results in flocs mutually repelling each other. However, where the river water mixes with saline seawater, the dissolved salts react with the flocs neutralising the charge. The flocs then coagulate and grow bigger (flocculation). They can no longer remain suspended in water and are deposited at the bottom of the channel. In a tidal river, flocculation spreads over a wider area, spreading the zone of sedimentation. Such sedimentation raises the land, which is then colonised by mangroves. Once mangroves are in place, sedimentation is accelerated. Mangroves also tend to protect coasts from the erosive action of large waves. The surface of a mangrove swamp is low and flat, dissected by small creeks and �commonly tidal. The levees of such creeks and bigger streams tend to form linear high grounds that may rise to several metres above the general surface and could be the only location not inundated in high tide. Small low mounds are also formed by the swamp fauna, such as mud lobsters. The soils tend to be fine-grained, rich in organics, poorly drained and saline. Mangroves cope with high salinity and anaerobic soil conditions by physiological adaptations such as the modification of leaves, sap and root systems. Species of mangroves, however, vary from each other in physiological adaptation and seed propagation strategy.

189

Coral reefs

A mangrove community is very rich in biodiversity and mangroves have been compared in this respect to tropical rain forests and coral reefs. Different species tend to adapt to variations in the micro-environment and different mangroves tend to exist in zones roughly parallel to the shoreline at the genus level. Such zonations are characterised by belts of Avicennia, Rhizophora, Brugeria, Ceriops, Sonneratia and Laguncularia. A dense community of mangroves is known as a mangrove forest or mangal. Individual plants could be shrubs or trees that range up to tens of metres in height. The community may also include non-mangrove plants such as nypa palms or ferns. Mangroves are common along the tropical coasts but are best developed in tidal deltas. The Ganga–Brahmaputra delta is possibly the biggest mangrove community in the world.

10.7.2╇ Sabkhas Salt flats, resembling coastal deserts, are common in the arid tropics. They are known as sabkhas, from Arabic. These barren salt flats can be extensive. In the Arabian Gulf, a set of sabkhas on the United Arab Emirates coast is more than 300 km long and locally up to 24 km wide. These sabkhas are about 7000 years old, and are currently located landward of coral reefs and beach ridges. Coastal salt flats are common in arid coastal areas, and they have been studied in detail in other parts of the Gulf, the Red Sea coast of Egypt, Baja California and Mexico. Generally, sabkhas are flat, salt-encrusted, geomorphologically stable, sandy plains (Fryberger et al., 1984). The origin of sabkhas is associated with a cycle of sedimentation related to a regressing sea level displaying a simple upward change. Previously deposited subtidal carbonate sediments are replaced over time, first with intertidal and then supratidal sediment. This is reflected in the subsurface by algal mats, carbonate sands and gypsum being replaced vertically first by gypsum and anhydrite and then nodular anhydrite. On an arid coast, secondary minerals grow by rapid evaporation leading to a concentration in the subsurface fluids and the precipitation of gypsum, anhydrite and halite, as secondary minerals in the interstitial space within the sabkha sediment. These evaporites fill the pore space within the deposited grains and then expand to raise the surface level of the sabkha. Wind action and rare floods scour the surface and keep it at a very low gradient.

10.8╇ Coral reefs Coral reefs are biological structures built offshore by a number of organisms, the principal members being corals and algae. This implies the requirement for a favourable habitat. Corals, for example, need light and a narrow temperature range, the optimal being 25–29°C. Coral reefs therefore grow only in tropical seas. Three types of reef are built by the corals. 1. Fringing reefs, which grow next to land.

190

The tropical coasts

2. Barrier reefs, which occur within a short distance of the land and are separated from it by a shallow body of water called the lagoon. The seaward side of a barrier reef may drop abruptly to deep water. Small patches of reef may occur in the lagoon. 3. Atolls, which are near-circular or elliptical reefs, although some atolls have strange shapes. These are found in the open sea, even in the middle of an ocean, and they build upwards from a relatively shallow underwater basement such as a sunken volcano or a seamount. The outer sides of an atoll drop off steeply to considerable depths, but a shallow lagoon occurs inside. A large number of atolls occur off the north shore of Java near Jakarta. The tropical Pacific Ocean and the Indonesian seas are dotted with atolls. All types of reef vary in size, usually from metres to tens of kilometres. The Great Barrier Reef off the eastern coast of Australia stretches for 2000 km from the Torres Strait (about 10°S) to a little beyond the tropic of Capricorn. It is separated from mainland Australia by a lagoon which is 80 km wide towards the north and 300 km towards the south. A number of small reefs and islands occur inside the lagoon. The Great Barrier Reef itself consists of about 2500 individual reefs. Of course, other coral reefs are not on this scale. The basic building block of a coral reef is the coral polyp enclosed inside a calciumbearing skeleton. The coral grows by adding new material to the outside of the skeleton. Parts of reefs are also built by algae. Reefs have a basal platform of calcium carbonate with vertical ribs growing out of it. The final reef is flat-topped with vertical sides. Part of the reef could be exposed, especially at low tide, but most of it would remain under water. This submerged part is a wonderful architecture of walls, flat platforms and canyon-like passages. Towards the top, corals build a hemispheric structure or a number of upwardextending branches. Different species of coral inhabit different parts of the reef, as do varieties of algae, fish, clams, sponges and foraminifera. Like mangroves and the tropical rain forest, a coral reef is a species-rich ecosystem. Coral reefs may provide some defence to the nearby coast from waves in tropical storms or small tsunamis. Coral reefs have been divided into an Atlantic province centred on the Caribbean and an Indo-Pacific province that includes most of the reefs of the two oceans. Specieswise, the Indo-Pacific province is much richer. Morphologically, reefs on the windward side of an island may differ from those on the leeward side. The windward reefs are straighter and better developed and may show a relationship between geomorphological zones and the ecological composition of reefs (Woodroffe, 2003). Figure 10.16 shows generalised sections across reefs from the two provinces. Reefs usually have three zones:€reef front (towards the open sea), reef crest, back reef. The steep reef front is a dynamic area with active corals. The reef flat is exposed at low tide, and at other times waves break over it. The back reef is a sheltered environment, rebutting either on land or, in the case of an atoll, sloping down to a central lagoon. Broken bits of reef form ramparts of rubble on the reef flat in areas where storms are frequent. The locations of coral reefs are Â�generally influenced by changes in sea level during the Pleistocene (Â�section€10.9), and modern reefs often start on a base of Pleistocene limestone or submerged volcanic material.

191

Fig. 10.16

Tropical coasts and time

Diagrammatic section through Indo-Pacific and Carribbean (West Indian) coral reefs. Figure reproduced, with permission, from Woodroffe, 2003:€a) based on Emery et al. (1954), Stoddart (1969) and McLean and Woodroffe (1994); b) based on Shinn et al. (1982) and Macintyre (1988)

10.9╇ Tropical coasts and time Two time-based observations are crucial for understanding coastal geomorphology. First, strictly speaking, all coasts are geologically young, not older than the last sea-level change at the end of the Pleistocene about 10 000 years ago. The loss of ice has led to the isostatic recovery of a number of extra-tropical coasts; the land is still being raised slowly against the sea level. Tectonic movements have episodically raised coasts in active areas, disrupting the ongoing coastal forms and processes, so that a new coast has been formed several times (Fig. 10.17). Second, the coastal forms, especially the forms of deposition, tend to change rapidly. A number of sea-level changes occurred during the Quaternary, as discussed in Chapter€16. The peak of the last glaciation was about 18 000 years ago. At that time, the sea level was about 120 m lower and a long distance away from the present coastline, unless the offshore gradient was very steep. This led to coastal changes. For example, in Southeast Asia a large part of the shallow South China Sea was exposed and major rivers flowed on this land. As glaciation ended, the sea level started to rise, the old coastline was

192

Fig. 10.17

The tropical coasts

Raised coral reef, eastern Jamaica. Such reefs provide information about past sea levels and tectonic uplifts. Photograph: A. Gupta submerged, and the sea advanced over the then exposed part of the South China Sea. The sea level continued to rise until about 6–7000 years ago when it stood several metres above its present position, after which it came down to the present level (see Chapter 16). None of the coastal forms in the tropics is older than several thousand years, even in areas not affected by tectonics. Even on an annual scale, coastal forms change rapidly. Beach profiles tend to change between seasons and in storms. A shift in wind vector alters the shape and size of dunes at the back of the beach, even if they are anchored by vegetation. Features in rock last longer, but they change more rapidly than most landforms. Cliffs extending as spurs on the shore platform change to arches and stacks and finally disappear. This may happen on a decadal scale. Large storms such as tropical cyclones (see Chapter 3) alter a coastline rapidly. This happens in the Caribbean, East Africa, South Asia, parts of Southeast Asia and Northeastern Australia. Tropical cyclones not only erode coastal forms disastrously, but also remove large quantities of sediment in order to deposit this material in lagoons, bays and other low areas, at times completely filling them. Large storms also transfer sediment to form beach ridges. The deltas on the Bay of Bengal coast are eroded by tropical cyclones at a span of

Tropical coasts and time

193

Fig. 10.18

(a)

(b)

(c)

(d)

(e)

(f)

Tsunami erosion and recovery, northern Aceh coast. (a) location of the image used for the Aceh coast; (b) coast on January 2003; (c) December 2004 (tsunami); (d) February 2006; (e) January 2007; (f) April 2008. Note removal of the beaches, destruction of the wetland and erosion along swales. Within 13 months, beaches have returned, although the wetland in the centre is under water. Vegetation has returned to the swales. The effect of the tsunami is not generally perceptible even in the 2006 image. The next two images show the continuation of the rebuilding process. IKONOS satellite images © Centre for Remote Imaging, Sensing and Processing, National University of Singapore, reproduced with permission. From Liew et al., 2010. See also colour plate section

several years. In between, more frequent, low-magnitude coastal processes tend to bring back the depositional features. A coast or a shore may be reviewed as a collection of transitional features that change every wet season, high storm or tsunami, but recover quickly to resemble their former appearance. An excellent example of this destruction and recovery comes from the 26 December 2004 tsunami on the Aceh Coast, Sumatra (Liew et al., 2010). The same 175 km of coastline was mapped from IKONOS satellite imagery at a 1 m resolution five times:€2003 (before the tsunami), late December 2004 and early 2005 (immediately after the tsunami), 2006, 2007 and 2008. The Aceh coast was eroded back for about 500 m in the tsunami, except at the rocky headlands, and almost the entire suite of depositional landforms (beaches, low sand dunes, wetlands) was removed. A new coast started to appear within weeks, closely resembling the pre-tsunami version (Fig. 10.18). Normal coastal processes returned to operate on the post-tsunami coast, and as the environmental constraints remained the same, the new coast is very much like the pre-tsunami one, although still several hundred metres back from its previous location.

194

The tropical coasts

The response and recovery of a coast to a geomorphic process is fast. The only two evidences of a powerful process such as tsunami or a hurricane storm that survive are (1) masses of corals or boulders detached from the source, carried a short distance and left as exotic blocks or (2) a buried layer of sand or coarser material in the subsurface, deposited during the tsunami or the storm and subsequently buried under fine sediment.

Questions 1. Explain how multiple geomorphic processes simultaneously operate on a coast. 2. What happens geomorphologically when waves approach the coast? 3. Explain the movement of sediment along a coast. 4. What controls the nature of a coastal drift? 5. Describe the usual sources of sand for a beach. Is this sand transported for a short distance or brought from afar? 6. Can a coast or shore be seen as a collection of short-term features? 7. Describe a mangrove swamp. Does it have any control on the geomorphology of the local coastal area? 8. In what way may coral reefs act as indicators of geological history? 9. How long-term are the effects of a particular tsunami or tropical cyclone on a coast?

11

Deltas in the tropics

This is a land half-submerged at high tide Amitav Ghosh

11.1╇ Introduction Deltas are partly subaerial and partly subaqueous accumulations of riverborne sediment deposited at the mouth of the river, with the sediment reorganised by tides, waves and currents. The characteristics of a delta reflect both the river basin and the dominant geomorphic process operating in the coastal region. Deltas therefore vary in geometry, morphology, operating processes and sedimentary characteristics. They develop best when the river contributes a large amount of sediment on a tectonically inactive, wide and shallow continental shelf. Deltas can form in both seas and lakes. The prime morphological division of a delta is between the subaerial and subaqueous parts. The limit to which tidal processes operate divides the subaerial delta further into a non-tidal upper part and a tidal lower part. In shape, deltas are approximately triangular or lobate, the apex of the triangle starting where the main river first splits into more than one channel. Channels then divide repeatedly and the surface of the delta is marked by a very large number of such distributaries, their number generally increasing towards the depositional basin, usually the sea. The main flow of the river uses only some of these channels; other channels carry less water. The subaerial delta is thus also divided into an active and an abandoned sector. Most of the discharge and sediment of the master stream flows through a number of channels in the active part, and the abandoned part includes channels that are old and moribund, and do not carry much water or sediment (Fig. 11.1). Towards the end of its course, a river usually carries fine-grained sediment:€sand, silt and clay. This is the material that generally builds deltas. Deltas tend to show a coarseningupward sedimentary sequence, with sand overlying clay and silt. River sediment flowing into the sea is reorganised by tides, waves or currents, depending on local conditions, and delta-building continues under water by sedimentation in the depositional basin. As expected, the surface of a delta is nearly at sea level and has very little relief or gradient. The highest areas are usually levees running parallel to the distributary channels, with low alluvial basins separating them. In many deltas, almost the entire surface except the levees becomes inundated during the rainy season or after a large storm. Under natural conditions, freshwater swamp vegetation covers the upper part of tropical deltas and mangroves flourish over the saline lower section. 195

Deltas in the tropics

196



Fig. 11.1

Landforms of a delta Many deltas are highly populated and cultivated, because of their flat gradient, richness in aquatic resources, wide biodiversity, availability of water and navigability along channels. On the other hand, deltas are extremely vulnerable to floods, storm damage and large-scale channel avulsions. Tropical deltas are hazardous places during hurricanes and typhoons, even during the average wet monsoon. As about a quarter of the world’s population live at or near deltaic coastlines or wetlands (Giosan and Bhattacharya, 2005 referring to Syvitski et€al., 2005), at times this translates into huge losses of life and property.

11.2╇ Distribution of deltas in the tropics Rivers draining tectonic mountains with high relief and shattered rocks generally transfer a large volume of sediment along their lengths to construct deltas at their mouths. Sediment accumulation is faster if the coast is shallow and of low gradient. Large deltas are therefore associated with the passive edge of continents and shallow seas, e.g. the Niger Delta. Figure 11.2 shows the location of large deltas in the tropics. In sum, deltas reflect the complex interaction between sediment supply from rivers, sediment accommodation space on the coast and coastal energy. Apart from rivers of significant size, smaller rivers also build substantive deltas for their size, if they carry large volumes of sediment and discharge onto a shallow sea. A number of rivers draining the islands of Southeast Asia, such as the Solo in Java, have impressive deltas. The large sediment load is derived from young steep

Age and evolution of deltas

197



Fig. 11.2

Distribution of major tropical deltas

volcanoes and intensive cultivated land. The annual rainfall on these islands could be as high as several thousand millimetres and the derived sediment accumulates on the broad continental shelves of the South China Sea. Woodroffe (2005) stated that most Southeast Asian deltas display a prominent anthropogenic impact. This effect is probably more dominant for smaller deltas, as sedimentation rates have remained similar for 7000 years for the Ganga–Brahmaputra (Goodbred and Kuehl, 2000a), 3000 years for the Mekong (Ta et€al., 2002) and for the entire period of delta formation for the Chao Phraya (Tanabe et€al., 2003). Hilly islands and the collision side of continents do not have large deltas but are drained by short steep streams. Such streams may carry a large amount of sediment, especially in floods, which is deposited in small deltas with a steep gradient. For example, streams tens of kilometres long drain the earthquake-prone Blue Mountains of eastern Jamaica, an area which is also impacted by tropical storms, a number of which reach hurricane level. A number of deltas are found on the northeastern and southeastern coasts of Jamaica:€steep, built of coarse gravelly material and drained by braided streams. Such features are known as fan-deltas (Wescott and Ethridge, 1980).

11.3╇ Age and evolution of deltas Deltas are geologically young worldwide, as they expanded in area only in the Early Holocene after the slowing down of the post-glacial sea-level rise, which provided accommodation space for rivers to deposit sediment and build deltas higher rather than forward (Stanley and Warne, 1994). Stanley and Warne dated the beginning of modern deltas at 8500–6500 years before present (BP), based on radiocarbon dating of material collected from boreholes sunk in the deltaic sediment. The dates, however, may vary regionally. For example, Goodbred and Kuehl (2000a) dated the base of the present Ganga–Brahmaputra Delta as 10–11 000 years BP. The delta of the Mekong started to prograde 6–7000 years BP. The large deltas of South, Southeast and East Asia, Ganga–Brahmaputra, Irrawaddy, Chao Phraya, Sông Hóng (Red River), Mekong, Chang Jiang and others, were formed in the Holocene, following stable to slightly falling sea-level conditions, high sediment discharge from the Himalaya and the Tibetan Plateau (Ta et€al., 2005), and strengthening of

198

Deltas in the tropics

the monsoon system. The highest Holocene sea level probably occurred about 6000 years ago. It was several metres higher than the present sea, but varied in age and height in different locations. In some cases, deltas probably built up vertically during the sea-level rise in the Early Holocene and, after the stabilisation of the sea level, prograded outwards as observed for the delta of the Changjiang. In other instances, such as the Nile and the Ganga–Brahmaputra, the present sea level is the highest, and fluvial sediments are actively accumulating on the delta surface (Hori and Saito, 2007). Sedimentation occurs where the river meets the sea due to deceleration of the river by seawater and flocculation following the mixing of river and saline waters, as explained in Chapter 10. As a result, when a river enters the sea as a jet of sediment-laden water, sand and silt are deposited first followed further on by clay. The front of a subaqueous delta, its lowest part, is therefore clay and as a delta is progressively built into the sea, sand is deposited on top of the clay settled earlier. Delta sediments therefore show a vertical coarsening pattern. The river water floats above the saline seawater as it flows into the sea, unless it is very heavily laden with sediment, and sinks. That, however, is a rare case requiring a very high sediment concentration, which used to happen at the mouth of the Huang He before dams and reservoirs blocked the flow of its sediment downstream. Mixing between the two sets of water occurs at the edge of the intruding river jet and the deposition of sediment builds underwater ridge-like forms. Over time and deposition of more sediment, these ridges are exposed above water as levees. The seaward face of the delta becomes a mosaic of river channels bounded by high levees with low waterlogged depressions such as bays and backswamps in between. Breaks in levees give rise to distributary channels that radiate from the main river across backswamps and bays carrying fine-grained sediment, and sandy sediment splays out through crevasses. The advance of the delta into the shallow sea is thus followed by the filling of the low ground between levees, transforming the bays and swamps into the subaerial deltaic plain. Avulsions and the switching of the main river to a lower part of the delta front focuses delta-building activities to a different area, leaving behind an abandoned sector. River-dominated deltas expand aerially in this fashion. The Mississippi is an excellent example. Tides, waves and currents, if strong, reorganise the sediment across the face of the delta and determine the final form and sedimentary sequence (Fig. 11.3). In a strongly tidal zone, the riverborne sediment moves back and forth over a length of the channel where flocculation takes place. Tidal-dominated deltas are therefore distinguished by the following characteristics:€funnel-shaped river mouths resembling estuaries; bars and islands in the channels, elongated in the direction of tidal movement; and a number of small tidal creeks between and through such islands. In the humid tropics, salt-resistant vegetation, mainly mangroves, tends to grow on the islands and along the banks of coastal creeks, anchoring the fine-grained sediment (Fig. 11.4). The Ganga–Brahmaputra Delta exhibits all these characteristics. Where waves are strong, riverborne sediment may be pushed back to give the delta edge a straighter appearance; such deltas do not protrude into the sea. The deltas of the Godavari, Senegal, Nile and Sâo Francisco are good examples. This often leads to the development of wide beaches and, if onshore winds are strong, lines of beach ridges and

199

Age and evolution of deltas



Fig. 11.3

Different types of deltas. After Hori and Saito, 2007



Fig. 11.4

Tidal creek through mangrove forest, Sundarban, Ganga–Brahmaputra Delta. Note the steep bank in silt-clay. Photograph: A. Gupta dunes behind them. The lower Mekong Delta exhibits such beach ridges (Fig. 11.5). The upper part of the Mekong Delta was primarily built by fluvial processes where multiple channels are found bordered with levees and separated by backswamps. This part of the delta was built in a sheltered bay. As the delta extended out beyond the headland, it was exposed to waves and the lower delta became tide- and wave-dominated. Here divergent

200



Fig. 11.5

Deltas in the tropics

Generalised diagram of the Mekong Delta. The upper part is river-built; the lower one is wave and tide controlled. Simplified from Ta et al., 2005 and bifurcating beach ridges rise 3–10 km above the sea level, separated by low inter-ridge depressions (Ta et€al., 2005). A strong longshore current tends to remove and transfer sediment away from the delta face, limiting the size of the delta. The best example is probably provided by the Amazon itself. The Amazon discharges a large volume of sediment into the South Atlantic which is carried northwest parallel to the coast by the Guinea Current. For hundreds of kilometres this sediment comes ashore to form the mudflats and beach ridges of north Brazil, French Guiana and Surinam. Part of the Orinoco Delta is built with sediment from the Amazon,



201

Delta morphology

Fig. 11.6

Comparison of morphology and sedimentary environment of different types of delta. The top diagram shows a generalised pattern of angle of sedimentary beds in a delta. From Hori and Saito, 2007. By permission of Wiley that was originally eroded from the Andes but which has arrived by a very long and circuitous route. Sediment carried by the current has built mudcapes (linear, round-ended, fine-grained promontories), several kilometres wide and up to 100 km long, a characteristic coastal feature between the mouths of the Amazon and Orinoco (Warne et€al., 2002). If the deposition of sediment in a receiving basin is carried out mostly by a river, it is a river-influenced constructive delta. If the sediment is reworked and re-deposited by tides or waves, the delta is recognised as a tide- or wave-influenced delta. Tide- and waveinfluenced deltas can be both constructive and destructive (Hori and Saito, 2007).

11.4╇ Delta morphology Deltas have three components (Fig. 11.6): • a low and flat delta plain forming the subaerial part of the delta • the seaward-dipping part which extends offshore beyond the delta plain, called the delta front

202

Deltas in the tropics

• the subaqueous low edge of the delta in front of and below the delta front, termed the prodelta. When deltas prograde, these components move forward, overriding each other. A delta plain can be described as an extensive low flat area that includes a number of active and abandoned distributary creeks that leave the main channels at a high angle. The channels are bounded by levees, and an assemblage of bays, marshes, tidal flats and floodplains form the rest of the delta plain. These low areas are climate-dependent. Freshwater swamps and mangroves provide a luxuriant surface cover in the humid tropics, as in the Niger and Mekong deltas. Vegetation is scarce in the arid tropics and instead a saline crust of gypsum and halite is found on the surface. If the sand supply is plentiful, dune fields appear, as in the Sâo Francisco Delta. Channels in non-tidal, river-dominated deltas tend to be sinuous but distributary channels may braid or anastomose in arid areas with discrete high flows and coarse bedload (Elliott, 1986). The interdistributary areas are flooded during the wet season but emerge as dry land at other times. Overbank flooding and crevasse splays during high flows are common in the humid tropics; so is channel avulsion. The effect of the tidal passage of water is discernible in the lower parts of tide-dominated deltas. The channels have a funnelled pattern resembling estuaries, and are of low sinuosity. Bedforms and sand bars are common. For example, sand bars in low sinuosity channels have been reported for the Mahakam Delta in eastern Kalimantan, Indonesia. The interdistributary areas in a tide-dominated delta are dominated by mangroves. For example, Sonneratia, Avicennia and Rhizophora grow in the part of the Mahakam Delta next to the sea. An extensive zone of nipa palm (Nypa fruticans) and Heritiera is found behind these mangroves, which is replaced by a swamp forest landward (Woodroffe, 2005, and references therein). The delta front is the subaqueous area in which sediment-laden river water meets the saline seawater and sediment is actively deposited, as described earlier. Part of the sediment is also delivered by traction. Bars are often deposited in front of a distributary mouth and the channel bifurcates to flow past the bar. This is how the delta spreads into the receiving basin. Tidedominated deltas tend to have wide delta fronts (Nittrouer et€al., 1986). Waves redistribute the sediment unless they are strong and persistent, when the shape of the bar is remodelled by the direction of wave approach. Exposure to strong waves may straighten the delta face, building beaches and beach ridges landward in sequence. The delta front advances, usually with a gentle seaward slope, building a coarsening-upward pattern of vertical sequence. The prodelta lies seaward of the delta front as a layer of clay and silt on the floor of the receiving basin. It is the advanced edge of the delta. The sediment is intercalated with silt stringers and thin shell beds and is highly bioturbated. The presence of shell beds and coarse sediment in prodelta sediment is commonly attributed to storms.

11.5╇ Delta sediments and sedimentary structures A structural pattern in delta sediment was recognised by G. K. Gilbert around lake margins in the western United States in 1885. This concept indicated that deltas start with

203

The Ganga–Brahmaputra Delta:€a case study

a Â�near-horizontal deposit of fine material in the receiving basin called the bottomset. A clinoform called the foreset advances over the bottomset as a delta progrades. The foreset is overlain by the topset, another layer of near-horizontal fine sediment. The angle of the foreset or delta slopes depends on the coarseness of the deposited material. This concept works for deltas in lakes and fan deltas with coarse material, where foreset beds have an inclination of 10–25°. The majority of the deltas of the world are in fine sediment. Milliman and Meade (1983) determined that the major deltas of the world receive 80–90 per cent fine-grained suspended load and 10–20 per cent coarse-grained bed load. The subaerial delta plain, described earlier, carries a range of sediment that varies in texture and structure. For example, levees, crevasse splays and beach ridges are usually made of sand, whereas the interdistributary lowlands are of clay. In the humid tropics, the mud is associated with decaying mangrove and peat. Interlaminated or thinly interbedded alternate mud and sand layers occur on the delta front. The variation has been attributed to a change in energy from processes such as changing tides. Where the beds are relatively thick (about 2 cm) as in the delta of the Fly River, Papua, they have been associated with seasonal changes in wind pattern. The succession of prodelta, delta front and delta plain gives rise to a characteristic vertical succession. A coarsening-upward sequence is seen from the prodelta sediment to that of the delta front. This is overlain by a fining-upward succession from the top of the delta front to the delta plain. The coarsest and best-sorted sediment occurs where the transporting energy is the highest. This fining-upward sequence may be replaced by a coarsening-upward one on the delta plain if dunes or beach ridges occur. This is a generalised succession which is not necessarily found everywhere, as deltaic sedimentation is variable and complex (Hori and Saito, 2007). The next section illustrates this complexity.

11.6╇ The Ganga–Brahmaputra Delta:€a case study Huge quantities of sediment, about a billion tonnes, brought down annually by the Ganga and Brahmaputra Rivers (see Chapter 9) have built one of the largest deltas of the world, with a subaerial extent of 111 000 km2 in Bangladesh and India (Kuehl et€al., 2005). The current delta also includes a prograding clinoform, more than 250 km across the continental shelf and extending about 125 km from land. The total area, summing the subaerial and subaqueous parts, is about 140 000 km2. This large tidal delta in the humid tropics has a monsoon setting, is highly populated and carries dense mangrove forests near the deltaface (Fig. 11.7). A number of research programmes have recently expanded our knowledge of this delta.

11.6.1╇ The background Tectonics divided the Ganga–Brahmaputra Delta into subsiding basins and Holocene and Pleistocene uplands (Fig. 11.7) which influence sedimentation pattern. The Bengal Basin has been subsiding since the Eocene, providing space for a very thick accumulation of deltaic sediment towards the south and the east beyond a hinge line marking the boundary

204



Fig. 11.7

Deltas in the tropics

Physiographic and tectonic map of the Ganga–Brahmaputra Delta. From Kuehl et al., 2005, © SEPM Online between the Indian continent and the oceanic crust (Sengupta, 1966). The delta surface slopes from 10–15 m elevation at the northwestern corner towards the southeast to the sea level. It is always low in relief. The only exceptions are two uplands, known as the Barind and Madhupur Terraces, that rise to about 15 m above the alluvial surface of the delta. The surface dips gently offshore, except near Chittagong towards the east, where a series of north–south-trending structural ridges and lows occurs offshore. A submarine

205

The Ganga–Brahmaputra Delta:€a case study

canyon, known as Swatch of No Ground, that reaches within 30 km of the coast towards the west, opposite the India–Bangladesh border, is the other significant offshore relief feature. This canyon is believed to have been incised during the lowstand of the huge river in the Pleistocene. The canyon is a link to the Bengal Deep Sea Fan. The offshore surface otherwise is covered by the Holocene sediment of the Ganga–Brahmaputra (Kuehl et€al., 2005). The delta, in fact the Ganga and Brahmaputra drainage basins, is under the influence of the Indian monsoon system. Rainfall is concentrated in a few months (late June to early October) when the rivers are high and the joined flow of the Ganga–Brahmaputra to the Bay of Bengal rises from 104 m3s−1 to 105 m3s−1 (Kuehl et€al., 2005). Most of the sediment is also distributed during this period. The tropical storms, a proportion of which become tropical cyclones, also inundate the lower delta and move a huge amount of offshore sediment landward to raise the coast, especially the islands. This is a tidal area, with tides reaching up to 5 m in the estuaries (Barua, 1990). The erosional damage from tropical cyclones is accentuated if they arrive at high tide. The two major rivers have been described in Chapter 9. They join in the centre of the delta at a place called Aricha, and the combined waterway, known as the Padma, flows southeast. The Padma joins another huge river, the Meghna, which drains the Sylhet Basin (Fig. 11.7), and the combined mass of water flows into the Bay of Bengal past Noakhali.

11.6.2╇ Morphology Kuehl et€ al. (2005) divided the entire Bengal Delta into three broad morphological compartments: • upper delta plain and flood basins • lower delta plain and delta front • subaqueous delta. The upper delta plain covers a 200 km wide zone of the land part of the delta. Kuehl et€al. (2005) extended the seaward boundary of this unit to the limit of the dry season inland extension of salt water. This area is primarily controlled by fluvial processes, influenced by tectonics and downstream coastal evolution and sea-level change. The river floodplains are bounded by uplands. The upland areas are broad fluvially dissected surfaces of Barind and Madhupur Terraces, which are probably Pleistocene in age, and less distinct, terrace-like Chandina and Comilla surfaces, which are younger. These uplands partition the upper delta in separate sub-basins, which are either narrow alluvial corridors or low and wide flood basins such as the Sylhet. The floodbasins such as the Sylhet and Atrai remain inundated during the wet monsoon under several metres of water that trap fine material (Fig. 11.7). Both the Ganga and Brahmaputra flow through 40–80 km wide corridors with braidbelts, overbank flooding and channel avulsions. Channel sand is the dominant subsurface material (Fig. 9.11 and Fig. 11.8). These rivers transfer sediment through the corridors downstream but a significant part of their sediment load is trapped there, leading to the build-up of this very flat and low surface. Rivers have changed their course repeatedly on the upper delta plain. The best example is that of the Brahmaputra, avulsing alternately



206

Deltas in the tropics

Fig. 11.8

The Brahmaputra (Locally known as the Jamuna) at Sirajganj, Bangladesh in low flow. Section of a sandy mid-channel bar. For perspective, also look at Figure 9.11. Photograph: A. Gupta into channels east or west of the Madhupur Terrace. Such major avulsions have occurred in this enormous river at an interval of several thousands of years. The present course of the Brahmaputra flows through the western channel. The lower delta plain and delta front have been defined by Kuehl et€al. (2005) as the part of the delta plain affected by salt water in the dry season, a zone about 100 km wide and rising to hardly 3 m at its inland limit. This was originally under a mangrove forest, locally known as the Sundarbans (Fig. 11.4); the forest is now much depleted. The coastline is a series of peninsulas between major channels, the peninsulas are criss-crossed by minor tidal creeks. The surface sediment is silt to clayey silt with a little sand. Clay-rich peat basins occur, as internal drainage and standing water conditions are common. Interdistributary islands (islands between distributary channels) extend seaward and in time transform into peninsulas as the tidal channels separating them silt up. On the seaward side, the islands extend for tens of kilometres under water as shoals that merge into an advancing wide apron at a depth of several metres. Kuehl et€al. (2005) have referred to this zone as the delta front. The river system has built an accretionary subaqueous delta on the continental shelf. A complete set of topset, foreset and bottomset forms has been identified on this subaqueous delta. Apparently the Ganga–Brahmaputra delta, like a number of large river deltas, has two sets of clinoforms both of which can be identified as delta fronts. Beyond the subaqueous bottomset lies the outer shelf. Evidence of submerged former deltas dating back to glacial lowstands occurs on the outer shelf, along with oolitic ridges that indicate a low supply of silica-rich sediment due to climate change during the Pleistocene (Goodbred, 2003). The submarine canyon (Swatch of No Ground) that dissects the shelf and starts from about 30 km off the present coast is a prominent feature on the Bengal shelf. The canyon

207

Deltas in the tropics, a summary

is 20–30 km wide with a steeper eastern wall, and is believed to fill with deltaic sediment which is then periodically emptied by turbidity currents that transfer the sediment to the enormous Bengal Fan that extends for 2000 km.

11.6.3╇ Evolution The evolution of the delta (Goodbred, 2003; Goodbred and Kuehl, 2000a) is related to climate and sea level changes during the Pleistocene and Holocene, and is discussed in detail in Chapter 16. Briefly, the low stand of the sea occurred around 18 000 years BP, when both the Ganga and Brahmaputra flowed through incised valleys and the surface of the Bengal Basin outside such valleys consisted of broad lateritic uplands at about 45–55 m below the present sea level (Goodbred and Kuehl, 2000b). Strengthening of the summer monsoon started around 15 000 years BP and enhanced precipitation led to the supply of an enormous amount of riverborne sediment to the delta. The first batch of sediment reached the Bengal Fan but subsequently, as the sea level rose to transgress the low areas, it trapped this sediment on top of the delta which then aggraded at an enormous rate between 11000 and 7000 years BP. After 7000 years, deltaic deposition was progradational, extending it seaward. The surface of the delta was also selectively filled by the avulsions of large rivers, principally that of the Brahmaputra. The main course of the Ganga migrated or avulsed eastward about 5000 years ago.

11.7╇ Deltas in the tropics: a summary Tectonics and climate determine the location and characteristics of deltas to a large extent. The deltas in the humid tropics are further characterised by the availability of a large amount of sediment, the occurrence of tropical storms and the presence of salt-resistant vegetation, primarily mangroves. The deltas of the arid tropics display salt flats and, in certain locations, dunes are formed on the delta plain. Many tropical deltas are densely populated and farmed. This investment is getting bigger, not only because of the increase in population density but also due to changes in land use. For example, a number of tropical deltas in the humid tropics were farmed for rice. They are still farmed, but the intensive management of soil and water has given rise to high-value crops such as vegetables for regional cities or aquaculture for the world. This is happening especially near the coastline. At the same time, the flat deltas continue to be extremely vulnerable to natural hazards such as river floods, channel avulsion and abandonment, tropical storms and wave surges, which not only destroy villages but also leave the paddy fields inundated with salt or brackish water. The damage inflicted by the tropical cyclone Nargis on the Irrawaddy Delta in the first week of May 2008 is a sobering example (Fig. 11.9). The spectre of such damages increasing due to climate change falls over the tropical deltas (see Chapter 19). We know from the history of adjustment of these deltas to climate and sea-level changes in the Quaternary and Early Holocene that such hazards are possible.

Deltas in the tropics

208

(a)



Fig. 11.9

(b)

The effect of a tropical cyclone on a major delta. (a) Part of the Irrawaddy Delta. (b) Same area after tropical cyclone Nargis. IKONOS satellite image. © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2009), reproduced with permission. See also colour plate section

Questions 1. List the major deltas in the tropics. How many of them carry mangroves? Group them into river-dominated deltas, tide-dominated deltas and wave-dominated deltas. 2. Why do tropical deltas carry a large population? 3. How old are the current deltas? 4. List the factors that determine the geometry, morphology and sediment of the deltas. 5. Construct a hypothetical vertical section for a prograding delta. 6. Describe the morphological characteristics of the Ganga–Brahmaputra Delta. 7. Figure 11.9 includes two IKONOS images showing the destruction of a village in the Irrawaddy Delta of Myanmar by tropical cyclone Nargis. Interpret the images to show the type and extent of the damage.

12

The arid tropics

The low and level sands stretch far away. P. B. Shelley

12.1╇ Arid areas The tropics are not green and humid everywhere. Semi-arid and arid climates prevail over about half of the tropics. Apart from a significant moisture deficiency, the arid tropics are also characterised by high temperature, a high diurnal range of temperature, a theoretically very high rate of evaporation (potential evapotranspiration > actual evapotranspiration) and extreme variability in rainfall. Ground cover is limited to low and scattered vegetation, and even this is absent from the extremely arid areas. A number of attempts have been made to define the boundaries of the arid region using various statistics on moisture availability. Grove (1977) defined semi-arid areas as receiving between 200 and 500 mm of rain annually. Even less rain was expected in the truly arid areas. Annual precipitation (P) and potential evapotranspiration (PET) are often used to construct an aridity index. For example, the aridity index of the United Nations Environmental Programme is the ratio P/PET. Semi-arid, arid and hyper-arid areas have aridity indices 0.50–0.20, 0.20–0.05 and <0.05, respectively (UNEP, 1992). A number of more complicated formulae have been used in the search for a better delineation of arid lands, but for our purpose we only need to identify regions where the temperature is high, the supply of moisture limited and the computed potential evapotranspiration unachievable. Aridity is caused by several environmental factors and the relative importance of these factors differs geographically. In general, aridity occurs at the latitude of subtropical high-pressure belts, where the descending air is heated, clouds are rare and rainfall is limited (Fig.€1.1). A long distance from oceans may also cause aridity, so continental interiors are usually dry. Orographic barriers may cause aridity on their leeward side where the wind descends and tends to pick up moisture. Lastly, cold ocean currents on the western side of southern continents contribute to aridity by causing atmospheric condensation before the winds reach the coast. There may be more than one factor behind the aridity of a place. Although arid areas thus reflect current climate belts and topography, some of these areas have been dry for a very long time. Certain areas in Africa and Australia have been deserts since the Tertiary, and the Namib in Southwest Africa has been dated back 80 million years to the Cretaceous (Thomas, 1997). Their sizes and boundaries have shifted over time, especially during the Quaternary (see Chapter 16). 209

The arid tropics

210



Fig. 12.1

Arid lands in the tropics Aeolian (wind-driven) processes and sand dunes are commonly perceived as associated with an arid landscape, but at least half of the arid tropics is not covered by sand. These are rock deserts, usually under low and scattered vegetation. After a rare occurrence of intense rainfall, they become subjected to brief durations of surface runoff and even flooding. Arid geomorphology therefore includes the examination of both fluvial and aeolian processes and related landforms.

12.2╇ Geological characteristics of arid lands Arid lands in the tropics tend to be on stable landmasses, mostly cratons (Fig. 12.1). Arid landscapes commonly consist of highlands bounding an intermontane basin or a number of mountain ranges or uplands rising out of low dry plains. Tectonics has given rise to a series of block-faulted, steep-sided mountain ranges and basins in the tropical and subtropical North and Central America over which the Sonoran and Chihuahuan Deserts have developed. The north–south-trending Chihuahuan Desert of Mexico is hemmed in between the Sierra Madre Oriental and Sierra Madre Occidental Ranges. In the Sahara, lowlands and dunes cover less area than rocky and stony mountains (Fig. 12.2). The Eastern Desert of Egypt is a rugged highland that runs south to the Ethiopian Plateau, where extremely arid areas occur towards the northern end of the Eastern Rift Valley of Africa at low elevations that locally drop to depressions below sea level. The Arabian Desert also has hilly and rocky surfaces. In Southwest Africa, the Namib Desert, located between the Atlantic coast and the Great Escarpment of Southern Africa, is a narrow strip of rugged upland along with dunes, plains and dissected hills. In contrast, the Kalahari Desert in the interior of Southern Africa, east of the Great Escarpment, is characterised by extensive plains of low relief and the huge swamps of the Okavango Delta, Makgadikgadi Pan and Etosha Pan. Areas of low relief are also common in the Thar Desert of India and Pakistan, but the land is also crossed by narrow ridges of ancient metamorphic rocks. The huge, arid interior of Australia exhibits extensive areas of riverine plains and sandy tracts, but ridges and rocky uplands are not uncommon (Fig. 1.2b; note the presence of vegetation). The Atacama Desert of South America, located on the lee side of the Andes, is somewhat unusual. It comprises a series of north–south-trending ranges and valleys that rise in elevation from the coast to the Andes.

211



Fig. 12.2

Arid hydrology

Topography in Sahara indicating the extent of rock desert and major sand seas. From Goudie, 2002. By permission of Oxford University Press

Deserts on old cratons, occurring in the Indian subcontinent, Africa, Arabia and Australia, have less relief than the Atacama, but are not necessarily a dune-covered land with low relief. The landform on such a geological framework is produced by processes of erosion and deposition carried out by both water and wind. The surface material includes bare rock, gravel, sand and riverine silt. A detailed account of the landforms of the arid lands of the world is in Goudie (2002), but even this brief description illustrates the variability of desert landforms. The mountains and uplands may get more precipitation and the relief may be shaped by rare instances of flowing water. In contrast, wind may take over as the dominant geomorphic process in the lowlands.

12.3╇ Arid hydrology 12.3.1╇ Desert rainfall Generalisations on rainfall in arid areas are difficult because rain gauges are few and far between; the records are seldom long-term; and the rainfall is sporadic by nature. Schick (1988) attempted a summary account of rainfall in the arid areas, based mainly on data from the instrumented small drainage basin of Nahal Yael near Eilat, and certain records from the Sahara. Averages may not mean much in such areas. For example, the mean annual rainfall in the Mount Sodom area near the Dead Sea is 50 mm but the maximum measured intensity is 50 mm in 30 minutes (Schick, 1988). The extreme 24-hour rainfall

212

The arid tropics

can exceed the annual mean by a factor of 3–4 once in 20–30 years in the Sahara. Most of the rain that falls is highly concentrated in time and Schick refers to one case in Eilat where the 24-hour rainfall was nearly twice that of the annual mean. The intensity of the desert rain is usually high. At Nahal Yael, intensities exceeding 14 mmhr−1 were associated with almost half the total rainfall over a period of 17 years. Desert cloudbursts over the Ahaggar Mountains in Sahara recorded about 1 mm per minute, which for partial periods of rain rose to 2 mm per minute. Grodek et al. (2000) described the 18 October 1997 rainstorm over Eilat in southern Israel. Three spells of rainfall occurred within 16 hours, showing near vertical rises in the cumulative rainfall diagram. The first rain spell reached an intensity exceeding 80 mmhr−1. The other two spells arrived with lower intensity, but even then, as Figure 12.3 shows, the figures were impressive. Grodek et al. also listed other storms, some of which had a very high short-term intensity. For example, the storm on 12 November 1973 recorded 5- and 10-minute intensities of 120 mmhr−1. The measures are usually not this high, but clearly rain can arrive with high intensity in the desert. The rain is usually from a travelling rain cell, whose velocities vary from nearly zero to tens of kilometres per hour. The cell boundaries are very sharply delineated. On the ground, the transformations from completely dry periods to blasts of intense rainfall are generally near-instantaneous. Rainfall records reported by Schick from Nahal Yael show an insignificant pre-‘rain’ amount, then a period of sharp and intense rainfall, followed by a tail of low intensity post-‘rain’ for individual rainfall events. Telemetric measurement of the frontal advance of a storm cell indicated that it took only 18 minutes to cross the 2 km long Nahal Yael basin (Schick, 1988). In general, rainfall is localised, intense and erratic over both time and space. Widespread rain occurs only when unusual meteorological incidents occur. A flood in 1976 in the Sinai and the southern Negev from a 72-hour rainfall was caused by the unusual combination of a Red Sea low-pressure trough and a Mediterranean frontal system that took an extreme southward deviation.

12.3.2╇ Flood generation Given this kind of intensity, storm rainfall from desert cloudbursts may cause brief periods of flooding, especially in places where infiltration to the subsurface is low. Thus, streams draining out of the mountains, where bare rock and thin soil and vegetation cover prevail, may rise briefly. For flooding in arid lands, antecedent moisture is usually not a consideration. Rainfall simulations in experimental basins show that the lag time from the initiation of rainfall to the onset of runoff is highly dependent on the rock and soils of the area. Schick (1988) provided useful data from infiltration experiments carried out in the Nahal Yael Basin. On a 34° hillslope on widely jointed granite, runoff started 72 seconds after the onset of rainfall. A total input of 1.05 mm of rain was required to stabilise the water loss, which was primarily due to infiltration at a rate of 2.8 mmhr−1. In comparison, on a colluvium underlain by schists and dykes and with a 95 per cent cover of angular 10–40 mm gravel, runoff started after 340 seconds. A total input of 6.4 mm of rain was required to stabilise infiltration, which was as high as 10.5 mmhr−1.

213



Fig. 12.3

Arid hydrology

Example of type rainfall in arid areas from Nahal Yael and Eilat. From Grodek et al., 2000. By permission of IAHS Following rain, floodwater may be trapped in the alluvial fills of the several kilometre long small headwater channels of arid drainage systems. Such fills tend to become saturated prior to the passage of the flood down-channel. Rapid infiltration occurs in these channels as the rain continues and, after the entire fill resting on the rocky floor of the channel is saturated, a flood wave proceeds down-channel on top of this saturated alluvium. Schick has suggested that the amount of moisture held in the alluvial fill of the headwater channel acts in the same way as the antecedent moisture that saturates humid



214

The arid tropics

Fig. 12.4

Example of high-intensity rainfall in arid areas and the corresponding discharge hydrographs with steep rise and fall. Data from the instrumented small drainage basin of Nahal Yael, near Eilat, Israel. Date of storm 18 October 1997. From Grodek et al., 2000. By permission of IAHS

areas for flood generation. The nature of flood hydrographs in arid areas depends on the subsurface material and storm characteristics. Usually a runoff-producing event is typified by a hydrograph with a near-vertical rise and a sharp fall (Fig. 12.4). This indicates that flash floods in small and medium-sized basins have the ability to significantly modify the alluvial channel. At their peak, flash floods may flow with high velocity; even supercritical flow is possible. A metre-high wall of water may lead the flood wave down previously dry channels (Leopold and Miller, 1956; Schick, 1988). It is, however, difficult to compute the recurrence interval of floods, as the rarity of floods requires a very long record of stream flow for computing flood statistics. Usually such records are not available. If an alluvial fill is present in the bigger main channel that the tributaries converge to, the downstream flood size may be reduced. This may also happen when the floodwater from the channel reaches the sandy floor of a wide basin. In sum, the protracted flows tend to modify the channel but such flows occur rarely.

215



Fig. 12.5

The rock desert

River channel in a dry landscape near Alice Springs, Australia. Photograph: A. Gupta

12.4╇ Arid landforms Landforms in the arid lands can be broadly divided into two classes depending on whether the ground surface is (1) in rock or underlain by a thin veneer of coarse sediment on rock or (2) in sand. In the first case, in spite of the aridity, running water is the chief geomorphic agent; in the second case it is the wind.

12.5╇ The rock desert Rocky slopes cover a substantive part of arid lands. Such slopes may be part of a steep mountain front, from the foot of which other slopes run down to a low mountain-girt basin of deposition. The mountain front and upper slopes are usually bare or under a thin cover of material being transported. This sedimentary cover becomes deeper over the lower slopes, and the low central basin is filled with fine sand, silt, clay and salt of different compositions. This tectonics-related basin and range topography is common in the arid North and Central America, although similar landforms also occur in other places. The contact of the mountain front with long slopes running down to the basin could be sharp, or partly masked by a set of alluvial fans. The nearly bare upper slopes are called pediment, the lower slopes under thicker sedimentary cover, bajada, and the depositional basin, which may or may not carry a lake, is the playa. These are terms from the southwestern United States and Mexico, but the usage has become general. Bajadas are not always present and a thinly veneered rock surface may directly reach the valley or basin floor (Fig. 12.5).

216

The arid tropics

The erosional surface of the pediment may cut across different types of rock and structure. Its surface could be in bare rock, rocks under a thin veneer of coarse clasts, or partially under alluvial fans. These surfaces are dissected by channels of ephemeral streams; and small, scattered residual bedrock hills, known as inselbergs or bornhardts, rise from the general surface. Both of these features are common near the mountains. Bryan (1922) has described these pediments as transportation slopes, over which sediment travels either under sheetflow or along small channels to major rivers or playas. Weathering profiles have also been noticed on rocks of the pediment, indicating their stability. Elsewhere, the country could be a mosaic of rocky ridges from which slopes run down to a depression; but the country remains open and not surrounded by mountains (Fig. 1.2b). Rock slopes from the ridges may end at a riverine plain instead of a playa, but either the river or the playa would operate as the local base level. Alluvial fans, common in the arid landscape at the contact between the steep mountain slope and the plains, are associated with valleys emerging from the mountains. These triangular sedimentary bodies emerge from an apex in the valley and then spread out over the lower ground. They are fan-shaped in plan view, but their longitudinal profiles could be concave or consist of a set of straight segments. The latter style is associated with postfan tectonic movements. The area of the fan is directly related to the area of the basin supplying the sediment. Arid alluvial fans tend to be smaller and steeper than their humid counterparts. The deposited sediment could be from debris flow, hyperconcentrated flow or channel flow. Often a fan displays evidence of all three processes. Fans are classified as debris flow (also called mudflow) fans and fluvial fans. Debris flows are common in arid areas. They result from brief periods of intense rain when accumulated lag sediment on channel floors is periodically moved. Coarser grain sizes are associated with higher fan slopes and variations in slope. However, Bull (1964) has shown that the abundance of sediment may also steepen fan slopes, even for sediment derived from basins underlain by mudstone or shale. As the stream responsible for a fluvial fan emerges from the confined narrow and deep valley in the mountains, it tends to change its form to flow through a wide, shallow, braided channel. It also changes its course frequently. The surface of the fan is therefore marked by deposits of variable processes and age that are related to the channel shifts. The older part of the fan, if left undisturbed for some time, tends to attain a surface patina known as desert varnish.

12.6╇ Running water in arid lands Running water, both as sheet and channelised flow, is the prominent geomorphic process in rock deserts. The water is derived from rain on the mountains or cloudbursts, as described earlier. Given the nature of rainfall in the arid tropics, these ephemeral channels mostly remain dry, being in flood for an extremely brief period of time. A number of floods on rivers draining semi-arid areas have been studied in detail. Examples include floods on the

217

Running water in arid lands

Wadi Watir and Wadi El Arish in Israel (Schick, 1988; Schick and Lekach, 1987) and those on the Finke or Todd River in Central Australia (Bourke, 1994; Bourke and Pickup, 1999). One common observation in all papers dealing with stream channels in arid areas is the episodic widening and filling of desert channels (Wolman and Gerson, 1978; Schick, 1988; Bourke and Pickup, 1999). It follows from earlier discussions that arid and semi-arid rivers need to adjust to variations in rainfall, rapid rises in runoff, flashy flood hydrographs, losses in the transmission of water downchannel, high suspended sediment loads and asynchronous tributary activities (Bourke and Pickup, 1999). All these affect the channel characteristics of the arid rivers as discussed for the Todd by Bourke and Pickup. They conclude that the concept of dynamic equilibrium between discharge, sediment and channel form proposed for rivers in temperate countries does not work for arid rivers which are subject to very high variabilities in discharge. Many arid-zone rivers are non-equilibrium rivers (Graf, 1988). The Todd River flows out of the rocky MacDonnell Ranges in Central Australia through a piedmont zone with Pleistocene fans over rock to an alluvial plain downstream. The river then disappears in the longitudinal dunefield of the northern Simpson Desert. It is part of the large ephemeral stream system that drains into the Lake Eyre Basin (Fig. 12.6). It passes the town of Alice Springs, which has a long-term rainfall record. The rainfall is infrequent and variable. Rain tends to fall between November and March (the southern hemispheric summer) as discrete storms of limited dimensions leading to asynchronous flow in the basin tributaries in this large basin. The riverbed is dry 98 per cent of the time (Fig. 12.7) and when the water arrives, both the rise and fall of the river are swift. The largest discharge measured at Alice Springs was 1290 m3s−1 in 1988, draining a catchment of 450 km2, which gives it a unit discharge of about 2.6 m3s−1km−2. This has been identified as a 50-year flood. In general, all discharges in an arid river like the Todd are episodic and flood-generated. The channel remains dry between floods (Fig. 12.8). Immediately out of the MacDonnell Ranges, the river has a single straight or winding channel, which is wide and shallow, and incised into Pleistocene alluvium, palaeoflood deposits and modern fluvial sediments. The dry channel displays a coarse sandy load that has formed large-scale ripples and tabular bars. Silt and clay occur in localised depressions in the channel. Bourke (1994) and Bourke and Pickup (1999) provide detailed accounts of the morphology and behaviour of the Todd, from which five major observations can be highlighted, summarising the general characteristics of fluvial forms in arid and semi-arid climates. 1. Unlike most streams of the humid tropics and temperate regions, the downstream variations in channel forms do not occur gradually, but as a set of step changes at tributary junctions where abrupt morphological changes are found. 2. The tributary junctions could be discordant (Fig. 12.9), with one channel temporarily barred with the sediment brought down by the other, due to localised rainfall and asynchronous fluvial activities in different parts of the channel net as described by Bourke and Pickup (1999). Schick (1988) has also referred to a study by Lekach on the Wadi Mikeimin, a tributary of the Wadi Watir in southeastern Sinai. Following a localised rainstorm, the Wadi Mikeimin built a fan at their confluence big enough to completely

218



Fig. 12.6

The arid tropics

The Todd river with reference to the Lake Eyre Basin. From Bourke and Pickup, 1999. By permission of Wiley obstruct the channel of the master stream. The obstruction remained for 22 floodless months, until a large flood removed the fan completely. 3. Transmission loss occurs in the river longitudinally, so the river ends in the middle of dunefields. This phenomenon is locally known as the floodout. A floodout is the site where channelised flow ends and floodwaters spill across the adjacent alluvial surface. The river divides into a number of small channels prior to the complete loss of its water. Floodouts in arid Australia have been investigated by Tooth (1999), who identified the marked reduction in channel capacity in the floodout zone and the increased diversion of floods overbank. The main causes of floodouts are either downstream reductions in discharge or the presence of barriers to channelised flow. Downstream reductions of discharge may happen due to the distance from the primary source of moisture, which is the mountain range at the head of the basin, combined with the loss of transmission in rivers flowing on coarse sandy alluvium. The barriers could be aeolian, alluvial, hydrologic or structural in origin (Fig. 12.10). A new alluvium burying older terraces marks the location of a floodout zone. Floodouts can be intermediate, with channels reforming downstream, or terminal where floodwaters dissipate completely.

219

Running water in arid lands



Fig. 12.7

The dry Todd River at Alice Springs. Note the width of the river in flood (as indicated by the length of the bridge) and the size of the cross-section required to transport floodwaters. Photograph: A. Gupta



Fig. 12.8

Annual floods at Alice Springs, 1953–1993. From Bourke, 1994. By permission of IAHS 4. The sediment of the channel displays a chaotic cut-and-fill structure, related to floods of varying sizes. As the channel effectively stays dry between floods coming down the mountains, such structures survive until the arrival of the next significant flood that is at least comparable in size to the one which last redistributed the sediment. Such floods tend to (1) widen the river channel, (2) vertically strip the channel and floodplain sediment, (3) erode new flood channels on the floodplain, (4) form small back channels on the floodplain and (5) mark the floodplain with macroturbulent scours, especially

The arid tropics

220

(a)

Ross

Todd

(b) Ross

Todd



Fig. 12.9

The confluence of the Todd with a major tributary, the Ross. Note the blockage at the confluence created by an asynchronous arrival of floods. From Bourke and Pickup, 1999. By permission of Wiley around trees and bushes. Flood aggradation then builds a channel and floodplain as insets within the flood-widened channel and deposits veneers of sediment overbank. Floodplain formation thus happens primarily in large floods. A section in the alluvium shows a sedimentation sequence repeatedly interrupted by erosional unconformities because of the numerous cut-and-fill activities. Bourke (1994) has described it as chaotic. A cross-section of the floodplain and channel commonly shows a series of steps at different levels on the floodplain.

Running water in arid lands

221

(a)

(b)

(c)

Fig. 12.10

(d)

Example of floodouts. From Tooth, 1999. By permission of Wiley 5. The vegetation, including large trees, tends to occur in channels and on the floodplain close to the banks due to the restricted availability of water. In flood, macroturbulent scours occur round the base of such trees. In the case of the Todd River, it most commonly happens round the large River Red Gums (Eucalyptus camaldulensis), where asymmetric elliptical scours become visible after the flood. The different levels on the floodplain also give rise to irregular flow boundaries (Bourke, 1994; Bourke and Pickup, 1999).

222

Fig. 12.11

The arid tropics

Schematic block diagram of the Todd River channel and floodplain. From Bourke and Pickup, 1999. By permission of Wiley Thus, arid rivers, as demonstrated by Patton et al. (1993), Bourke (1994) and Bourke and Pickup (1999), cumulatively display the effects of floods at varying scale. Low-frequency catastrophic floods are too large for the river channel to accommodate the discharge. The high-magnitude flood therefore erodes the entire channel and floodplain to make room for the water to pass. Subsequently, the channel and floodplain are built up over time by smaller floods depositing within the greatly enlarged channel. Bourke and Pickup proposed channel forms occurring at three scales, which they termed small (within-channel flows), medium (flows which extend across floodplains) and large (catastrophic and rare floods). Figure 12.11 illustrates the channel forms of the Todd River and the comparative size of three valleys excavated by floods at different levels. Bourke and Pickup (1999) indicated that dimensions across the three classes vary by an order of magnitude. In the Todd River, they expect longitudinal bars in these categories to measure 10â•›100 and up to 2000 m, respectively. The estimated return periods for the three categories of these floods are 10, 100 and 1000 years for the Todd. The effects of the smaller floods are more visible in the headwaters and the piedmont reaches because of the transmission loss further downstream.

223

Aeolian geomorphology of sandy areas

The description provided for the Todd can be extended in a generalised fashion to arid regions on rocks or a thin surface veneer on rocks. Rainfall and floods are rare, but running water is still the important geomorphic agent, at least in the arid mountains and the pediment zone. The form of the headwaters in the mountains is structure-controlled and can vary, but the streams widen as they emerge from the mountains to the lower plains. The channel and the floodplain reflect only the work of floods and may demonstrate a nested appearance related to floods of different dimensions and recurrence intervals. The channels remain dry between such events, exhibiting forms on their bed that were created by the last flood. The rivers, however, are likely to disappear downstream in some form of floodout.

12.7╇ Aeolian geomorphology of sandy areas The extent of sandy areas varies between the arid regions of the world. For example, sands cover nearly half of the arid areas of the Sahara, Arabia and Australia. In contrast, sand dunes are localised in the Americas. It is mostly quartz sand, with a small proportion of feldspars and iron or aluminium silicates that have survived weathering in the dry climate. The sources of the sand are riverine plains, beaches, coastal lagoons and playas. Sand from playas also contains evaporites such as gypsum. Large sandy extents carrying dunes of various sizes and forms are known as sand seas or ergs. Small areas with relatively smaller dunes are called dune fields.

12.7.1╇ Aeolian process Sand grains are picked up, carried and deposited by wind. Bagnold (1941) is recognised as the pioneer researcher in aeolian processes and dune formation. He studied sand transport and dune forms both in the field (mostly in North Africa and West Asia) and also in the laboratory. Satellite imagery has recently been added to these techniques as a major tool in dune research. There is a limit to the size of grains that the wind may pick up from the ground and the nature of aeolian transport determines the grain size (Fig. 12.12). Particles smaller than about 60–70 μm are transported by suspension in turbulent wind eddies. Of these, the very small particles (<20 μm) can be carried in the wind for a long time and transported a long distance. Particles, limited to 60–70 μm in size, however, are carried for shorter periods and distances before being deposited as dust grains. Bigger grains, measuring 60–500 μm, move by saltation, the most effective means of transport that is primarily responsible for building dunes. Saltation is a series of downwind jumps. A sand grain is picked up in wind eddies, rises steeply and then comes down in a flatter trajectory, displaying a parabolic path. As it impacts the ground surface the force of the impact causes other grains to jump forward, but they move a shorter distance and at less speed. This process is known as reptation. The original grain may also recoil from the impact and finally come to rest after a series of decreasing hops. The impact of saltating particles also causes grains bigger than

224

Fig. 12.12

The arid tropics

Diagrammatic sketch showing sand transport by the wind. Generalised from Lancaster, 1995 and Kale and Gupta, 2001 500 μm to move or roll along the surface. This is creep. Most of the sand grains in deserts move by saltation. In sandy areas, a cloud of fine sand moves downwind by saltation. The efficiency of this process depends on several factors. Sand grains move when the force of the wind (lift, drag) exceeds the weight of the grain and the grain can be airborne. As the velocity of the wind increases over a surface of sand grains, the grains start to vibrate and at a certain velocity determined by the grain size, they start to leave the surface. The velocity of the wind at which this happens has been termed the fluid threshold. However, the grains which become airborne following an impact with a saltating grain start moving at a velocity lower than the fluid threshold. This second threshold is known as the dynamic or impact threshold (Bagnold, 1941). One saltating grain therefore causes an exponentially increasing number of grains to move. However, as the velocity of the succeeding grains are less, about a tenth of the velocity of the impacting grain, a steady state is reached in seconds with a given number of grains moving in unit time. Wind entrainment of grains is a complex process that is influenced by local conditions, such as the size, shape, sorting and packing of surface grains. The presence of moisture, surface crusts and surface roughness including vegetation all act as local impediments. As a moving grain is commonly deposited on a dune, grains usually travel or come to rest on a sloping surface. A detailed review has been given by Lancaster (1995). Effective wind action only occurs close to the surface, usually within about 1 m of it.

12.7.2╇ Aeolian forms of erosion Wind erodes either by blowing at strength or by bringing fine grains to impinge on standing objects, especially if the surface material is not indurated. Wind erosion is the sum of three sub-processes:€ deflation (removal of material), abrasion (scratching with fine grains) and attrition (breaking down of the grains being carried in the wind by mutual impact). The aeolian erosional forms are almost always smaller in scale than the depositional forms.

225

Aeolian geomorphology of sandy areas

Deflation, which implies carrying away of material, happens where the wind blows across a surface of unconsolidated fines, especially sand and silt. Abrasion causes erosion of individual rock clasts or upright objects. The impact of silt and sand grains being carried by the wind may provide a rock clast with smooth facets or the opposite, eroded pits or lines on its surface. Such rock clasts are known as ventifacts. Ventifacts are generally pebble-sized but may range up to boulders. The impact of sand grains also erodes soft bedrock rising above the land surface, giving it a characteristic ridge-like appearance. Such features are known as yardangs. Yardangs are usually not more than 100 m in length. Bigger yardangs have been reported from the central Sahara and Egyptian Deserts. These are up to 1 km long and locally cut into sandstones. Not much is known about these megayardangs.

12.7.3╇ Aeolian forms of deposition Aeolian forms are usually built by the deposition of sand grains forming small ripples or big dunes. The small-scale depositional form of wind ripples is common in sandy areas, occurring both on flatter surfaces and on dune slopes. They are created by saltation and reptation. The ripples are elongated orthogonal to the wind direction, although local surface slopes may partially change their orientation. Their size depends on the sand grains, with most ripples having a wavelength of 50–200 mm and an amplitude of 5–10 mm. Both measures increase with texture. For ripples in coarse sand or granules, the wavelength and amplitude increase to 0.5–2 m and 10 cm or more, respectively. Wind ripples tend to be asymmetrical with a steeper lee slope that ranges between 20° and 30° (Bagnold, 1941; Sharp, 1963). Sand dunes are the expected depositional forms in arid areas. These dunes are formed from the interaction between sand grains on the surface and the shearing flow of the wind close to the surface. Desert dunes have a dynamic similarity with the subaqueous dunes discussed in Chapter 8. Dunes vary in size and form, which has led to a number of classifications. Several of these also use the relationship between dune form and wind direction as a classifying criterion. It is, however, easier to classify dunes solely on morphology, although the relationship with wind direction does exist. Table 12.1 summarises the morphology and distribution of different types of dune. Four major types (Fig. 12.13) have been recognised. Minor forms also occur due to vegetative or topographic control on wind deposition. A specific dune may exist as a discrete feature, in which case it is a simple dune. If dunes of the same type are juxtaposed, for example, a small crescentic dune climbs up the back of a large crescentic dune, the entire feature is a compound dune. If dunes of more than one type occur together, for example, a crescentic dune is found next to a linear dune, the entire feature is a complex dune. A star dune, which has a peaked form like a pyramid, may have flanks that are linear dunes. Compound and complex dunes are common in large sand seas. According to Lancaster (1995), dunes of sand seas have a hierarchical structure regarding their size and spacing. Wind ripples with a spacing of 0.1–1 m cover about 80 per cent of the land area in all sand seas. All sand seas also exhibit discrete simple dunes or

The arid tropics

226

Table 12.1╇ Morphology of dune types Type and subtype

Common location

Wind regime

Crescentic dunes Barchans

Margin of sand seas and small sand corridors:€Namib Sand Sea, Skeleton Coast dune field; total area coverage small. Covers about 40% of all sand seas. Dominant dune forms in the Thar Desert; Takla Makan and Teneggar Sand Sea; Jafurah; Nafud; eastern and northern Rub’ al Khali; northern Saharan Sand seas; Namib Sand Sea; and North American dune fields. About half of all sand dunes in certain areas. Examples include Simpson-Strzelecki Sand Sea, Rub’ al Khali, Sinai, parts of the eastern and northern Sahara, Akchar Erg (Mauritania), Kalahari Sand Sea and Namib Sand Sea. About 8.5% of all dunes. Grand Erg Oriental, Rub’ al Khali, Gran Desierto, Erg Fachi-Bilma (Niger) and Namib Sand Sea. Found in coastal and semi-arid dune fields, not in large dune areas, the only exception being the Thar Desert.

Wind directions have a narrow range.

Multi-directional or complex wind regimes, especially in the months when the winds do most work. Not well known, probably unidirectional wind regime associated with some vegetation.

Namib coast, semi-arid areas Australia, Kalahari.

Originates from sand trapped by clumps of vegetation.

Crescentic ridges Also called transverse dunes or barchanoid ridges. Simple, complex and compound forms occur.

Linear dunes Simple, compound and complex forms occur.

Star dunes

Parabolic dunes

Other dunes Nebkhas (also called coppice dunes) Lunettes Topography-controlled dunes

Wind directions have a narrow range.

Wind regimes unidirectional or bi-directional, rarely complex.

Occurs downwind from small playas. Occurs upslope or downslope of a topographic barrier to wind.

Note:€Not all sand deposits are in the form of dunes. Zibers or sand sheets are low-relief sand surfaces common in the Namib Sand Sea and Gran Desierto. Sand sheets occur in areas where dunes cannot develop due to factors such as a high water table, coarse sand and vegetation.

227

Fig. 12.13

Aeolian geomorphology of sandy areas

Diagrammatic sketches showing different types of dune compound and complex dunes at a spacing that falls between 50 and 500 m. Wider-spaced (> 500m) complex and compound dunes occur in some sand seas (Lancaster, 1995). In general, the dunes of a region display a common regular pattern, comparable spacing and a size affinity. Dunes are aeolian depositional forms that are built by sand grains. The mineralogical composition of the grains depends on the source. Quartz and feldspar are most common, as the usual sources are weathered igneous rocks and sandstones, but fine rock fragments and volcaniclastic material also occur, especially if a short distance separates the source area from the dunes. The colour of the sand varies. Dunes in certain locations, such as the Simpson-Strzelecki Desert of Australia and the Kalahari in Southern Africa, display a reddish tint, which commonly comes from iron oxides embedded in the minute pits that occur on the surface of the grains. Dunes in other areas lack this reddening effect and, as Lancaster (1995) has described them, they are paler. Sand grains are deposited mainly in three ways. Dune grains avalanche down the lee face of the dunes, leading to dune movement. Some grains remain in suspension for a short period on the lee side in the shadow of the crest where there is less wind movement before falling on the dune. Wind ripples tend to migrate up the windward side of the dune adding sand grains. Each process exhibits an identifiable stratification when a section is cut through the dune. The processes are also reflected in the texture of the sand grains.

228

The arid tropics

Wind ripples are in relatively coarse grains because the material moves on or near the surface by saltation and reptation. In contrast, grains that are deposited from temporary suspension in the lee of the dune crest are finer and better sorted. In general, most dunes are made of very well to moderately sorted fine to medium sands. Finer sand, as expected, occurs on the dune crests and coarser material forms the bases or plinths. Case studies indicate that avalanching deposits form most of the body of crescentic and parabolic dunes, whereas wind ripples are the important deposition process for linear and star dunes (Lancaster, 1995). In some sand seas, areas between dunes cover a significant proportion of the region, and such interdune areas can be bare rock, alluvial material or wind ripple deposits. In sum, the origin and development of different types of dune are related to variations in wind speed and direction and the transport of sand grains. The airflow accelerates on the windward side of the dune, eroding and transporting sand to the crest. The airflow can, however, be varied and complex in lee of the crest, leading to a complex pattern of erosion and deposition of sand grains. In spite of the complexity, dunes tend to be similar in morphology over wide regions, indicating a relationship between dune morphology and regional wind direction and velocity, and the availability of sand, so that a condition in dynamic equilibrium is achieved. Five factors determine dune morphology: • wind regime, especially wind variability and strength • amount of sand that is available • texture and sorting of dune sands; these are expected to have control over the size and spacing of dunes • vegetation, if present • time. Simple dunes have faster transport rates and are controlled by the annual or seasonal pattern of wind speed and direction. In contrast, much larger complex and compound dunes move extremely slowly and are not influenced much by local wind conditions. It is the plentiful supply of sand that affects their morphology and existence. The life spans of simple (10–102 years) and complex or compound dunes (103–105 years) are strikingly different (Lancaster, 1995). The basic dune morphology (Table 12.1) is related to wind direction. Crescentic dunes are formed where wind blows from a single direction; linear dunes are found in regions of bimodal winds; and star dunes originate in complex wind regimes, where the wind may blow from different directions. The minor varieties are related to local environmental parameters, such as vegetation for nebkhas or parabolic dunes, and topographic barrier for dunes that are described as climbing or falling dunes. Sand seas, which cover large areas, thus carry different types of dune (Fig. 12.14) due to spatial variations in the wind regime, wind direction and sand supply (Kar, 1987; Lancaster, 1995). The simplest patterns are those of sand sheets, zibars and low crescentic dunes close to the source area of sand. Zibars are low rolling dunes without slip faces, usually in coarse sand. The dune forms become bigger and more complex with increasing distance, which allows for the accumulation of thicker sand. The dunes thin out and become simpler further downwind from the source area.

229

Fig. 12.14

Aeolian geomorphology of sandy areas

Different dune types in the Namib Sand Sea, Southwest Africa. From Lancaster, 1995. By permission of Taylor and Francis The huge sand seas like the Namib or Australian ones that occur over stable cratons have been formed over a very long period of time, and their extents have fluctuated following climate changes during the Quaternary (Chapter 16). Wind ripples are formed over small areas in a very short time, almost instantaneously. Dunes may form within a range of months to a hundred years and are at least in partial equilibrium with the rate and direction of sand movement governed by surface wind characteristics. Large sand seas evolve in appropriate tectonic settings, over a much longer period, millennia to millions of years (Lancaster, 1995).

230

Fig. 12.15

The arid tropics

Semi-arid landscape near Yulara, Northern Territory, Australia. Photograph A. Gupta

12.8╇ Conclusion Arid and semi-arid tropics extend over a large proportion of the Earth’s surface. More than half of this area is in rock or with a thin sedimentary cover over rock; the rest is in sand. Physical weathering tends to mechanically disintegrate the surface material, except towards the arid coasts where the presence of moisture, commonly as fog, and salt leads to chemical weathering. In spite of the general aridity and the sporadic nature of precipitation, running water as sheetflows or streams is the dominant geomorphic process on rock. The channels and depositional forms of the streams of the arid regions, however, reflect a nested arrangement associated with past flood history. They have been described as nonequilibrium rivers. The sandy deserts are commonly in geologically stable areas, and several of them have existed for a very long time. In spite of stable cores, the distribution and extent of arid areas have fluctuated in time, especially with climate changes during the Quaternary, and relict forms do occur. This is discussed in Chapter 16. Rocky outcrops and ridges occur in sand deserts, but the sand landscape is essentially formed by wind deposition:€a mosaic of wind ripples, dunes of varying morphology and interdune deposits. Figure 12.15 from Australia is a visual summary of a semi-arid landscape. Low hills are flanked by a rocky piedmont replaced by lowland where sand has been deposited to form low dunes under desert vegetation.

Questions 1. Where are the arid tropics? 2. Why is it difficult to generalise statistically on the rainfall in the arid tropics?

231

Questions

3. Describe an arid alluvial fan. 4. Is the role of running water important in the arid tropics? 5. Do the rivers of the arid tropics differ in morphology and behaviour from those in the humid tropics? 6. How do the smaller rivers of the arid tropics end? 7. Explain how the origin and development of different types of sand dune are related to variations in wind speed and direction, and the transport of sand grains.

13

Tropical highlands

Towards the north is the Himalaya, the lord of the mountains, the abode of the gods. Kalidasa, translated from Sanskrit

13.1╇ Importance of highlands Highlands are geomorphologically important regions. A large part of a river’s discharge and nearly all of its sediment may come from mountainous areas. These elevated areas also impact on the regional climate, especially precipitation. Precipitation arrives as snow at high altitudes, even in the tropics, and as rain on the lower windward slopes. Such precipitation can be very high in a tropical storm or when winds converge near the Inter Tropical Convergence Zone (ITCZ). Highlands may therefore supply a large proportion of a river’s discharge. In a seminal paper on big rivers, Potter (1978a) discussed sediment transfer in South America from the active to the passive margin by large east-flowing rivers. From the tectonically active margin to the passive one is a common direction for many large rivers crossing continents. Tectonic activities in the mountains not only create high relief but also increase the supply of sediment by providing fractured and brecciated rocks, steep slopes, earthquakes and volcanic activities (Milliman and Syvitski, 1992). Where high stratovolcanoes are present, erupted material could be the major source of sediment for regional rivers, as in the Andes and the Indonesian islands. On Java, frequent eruptions from the Merapi and other volcanoes provide the material for floodplains and the channel bars of streams that flow south to the Indian Ocean. Volcanic sand even builds the coastal features including large sand dunes (Fig. 13.1). Sediment from highland slopes enters the valleys of small headwater streams (Fig. 13.2) and is ultimately transported out of the mountains by larger rivers. A high proportion of the sediment is deposited at the highland–lowland contact at the foot of the mountains where it builds alluvial fans, which may have huge dimensions. The fans south of the Himalaya have been described as megafans, the name indicative of the volume of sediment coming out of these mountains. A large part of the Andean sediment is similarly deposited in the foothills region east of the tropical Andes. The rest of the sediment is stored in downstream floodplains or conveyed to the coast. It is the Andean sediment that builds the northeastern coastal plains and deltas of South America (Meade, 2007). Smaller fans are found where less sediment is eroded from the highlands, as at the foot of the mostly forested Main Range of the Malay Peninsula. A large river system may collect discharges from all tributaries 232



233

Importance of highlands

Fig. 13.1

Sketch map of part of central Java showing the volcanic complex of Merapi, the south-flowing drainage and the south Java coast. Pyroclastic material erupted from the volcanoes is the principal source of sediment for these streams. Sand of volcanic origin has built the coastal landforms including the well-known black sand dunes of Perangtritis. The height of Merapi changes between maps, as it depends on the latest eruption! draining the entire basin, but it is only the tributaries originating in the mountains that are important contributors of sediment (Fig. 9.1). It is therefore necessary to evaluate the importance of tropical highlands and the processes that operate on their slopes. Three major dynamic processes operate in the mountains:€ glaciation, slope processes and fluvial action. These are reviewed in this chapter.

234



Fig. 13.2

Tropical highlands

Small stream high in meltwater carrying a high sediment load, Lidder Valley, Kashmir Himalaya, very late in April. Photograph: A. Gupta Effective glaciation and related slope activities, however, are currently confined in the tropics to the two massive mountain chains:€the Himalaya and the Andes.

13.2╇ Glaciation in tropical mountains In spite of their location in the low latitudes, the higher slopes of several tropical mountains are perpetually under ice. Examples of such mountains include not only the Himalaya and Andes, but also mountains that are much smaller, such as the high volcanic peaks in East Africa:€ Mount Kenya, Kilimanjaro and Ruwenzori. The present glacier covers on the Kilimanjaro and Ruwenzori are about 5 and 4.5 km2 respectively (Osmaston, 1989a, 1989b). Glaciers occur above 4620 m on Mount Jaya in Papua (Indonesia), where three small permanent ice caps with a total area of about 7.5 km2 still survive although in retreat (Allison et al., 1989). Glaciation was much more extensive on all these mountains during the Pleistocene (Chapter 16). For example, geomorphological and sedimentological evidences indicate glaciation on the Aberdares and Mount Elgon in East Africa, two peaks that are no longer under ice (Mahaney, 1989). Landforms indicative of past glaciations are found even on the top of Mount Kinabalu, a 4131 m high granitic pluton in the Malaysian state of Sabah on the island of Borneo at a latitude just above 6° N.

13.2.1╇ Glaciation in the Himalaya The Himalaya Mountains are a series of parallel and asymmetrical ranges progressively rising in elevation northwards as described in Chapter 2. The high elevation of the northern line of ranges supports glaciation, and the asymmetry in slope and the east–west decrease

235



Fig. 13.3

Glaciation in tropical mountains

Glacier in the Western Himalaya. Photograph: A. Gupta in precipitation control the size of the glaciers. In general, glaciers are smaller on the steep southern slopes and in drier areas such as Ladakh. The glaciated area on the south-facing slopes of the Baspa River valley in Himachal Pradesh is only a quarter of the size of the glaciated north slopes (Vohra, 1981). Given favourable conditions, some of the glaciers reach considerable lengths, e.g. Siachen (72 km), Biafo (62 km), Hispar (61 km), Batura (59 km) and Baltoro (58 km), but most glaciers in the Himalaya are 1–5 km long. The Himalayan glaciers were longer in the Pleistocene, and descended to a lower elevation (Chapter 16). At present, they descend to about 4500 m in eastern Nepal and Sikkim, 4000 m in the central part of the range and about 3700 m in eastern Kashmir further to the west (Fig. 13.3). Glaciers develop best in the northernmost range€– the high Greater Himalaya€– and tend to be concentrated in the vicinity of the great Himalayan peaks. These areas of concentration include, from east to west, Namche Barwa; the area overlooked by Everest, Makalu and Kanchenjunga; the Dhaulagiri–Annapurna–Manasalu region; Nanda Devi; and finally the Nanga Parbat–Kolohai–Nankun stretch (Vohra, 1981).

13.2.2╇ Glaciation in the Andes An asymmetry prevails regarding the presence of permanent ice between the two sides of the Andes. The coastal mountains to the west are arid or semi-arid, and a lack of moisture prevents glaciation there even at high altitudes. On the eastern side, glaciers thrive

236

Tropical highlands

on abundant moisture from the Atlantic and are found as far north as the Caribbean. The Andes, as expected, are heavily glaciated towards the south beyond the tropics, but even the tropical ranges are glaciated because of their high elevation and the moisture availability. The semi-arid ranges that run along the Argentina–Chile and Bolivia–Chile borders currently support glaciers above 6000 m on massive volcanoes, but further to the north, beyond about 20°S, glaciers occur at a lower elevation. Extensive glacial systems are found on the high massive peaks of Cordillera Real, Cordillera Blanca, Cordillera Huayhuash, Cordillera Apolobamba, Sierra Nevada del Cocuy and Sierra Nevada de Santa Marta, but small ice masses may occur on small peaks and narrow ridges rising to 5000 m and above (Clapperton, 1983).

13.2.3╇ Glacial geomorphology of high tropical mountains The high ranges of the Himalaya and Andes have been magnificently eroded by glaciation into a landscape of huge mountain peaks, steep ridges, high vertical rock walls, and wide and deep valleys. Obviously it is the valley glaciation that needs to be studied in these ranges, as ice caps did not develop in recent geological history over any part of the current tropical landmass. There are, however, complications to the general pattern, as local slopes determine the character of glaciers. The Raikot Glacier near Nanga Parbat, western Himalaya descends from about 7000 to 5000 m very steeply in a series of ice cliffs. It then moves on a gentler slope of about 7° to its terminus at 4250 m (Shroder, 1993).

13.3╇ Mechanics of mountain glaciation A glacier is a well-defined body of ice that moves either forwards or backwards, and rests on either bedrock or soft sediment. Some snow exists all year above a certain elevation in the high mountains. This elevation is known as the snow line or firn line. Both low temperature and precipitation are required to accumulate enough snow at this altitude to form glaciers.

13.3.1╇ Formation of glacial ice Freshly fallen snow has two properties. It has a low density of about 0.05 gcm−3 and a hexagonal crystal structure. The density of glacial ice in comparison is 0.85–0.9 gcm−3. The transformation from snow to glacial ice occurs by three changes:€compaction, recrystallisation and the refreezing of the water derived from part melting of the snow. Compaction by overlying layers and partial melting packs snow tighter and changes the hexagonal crystals to near spherical forms. This produces granular ice with a density of at least 0.4€gcm−3, which is known as firn. The transformation from snow to firn may take hours or days in high tropical mountains. The transformation from firn to denser glacial ice takes much longer, decades or more. Pore spaces inside firm are eliminated over time by recrystallisation and the freezing of meltwater. Only a few air bubbles remain as the density of ice approaches 0.9 gcm−3. There may be further changes to the ice, but such changes are minor.

237



Fig. 13.4

Mechanics of mountain glaciation

Mass budget and mechanics of motion of a diagrammatic valley glacier The melting point of ice is pressure-dependent and varies according to depth. Ice therefore melts at slightly different temperatures at varying depths through the glacier and not necessarily at exactly 0°C. The term pressure-melting point is used to indicate this property. Glaciers occur in different parts of the world and they have been divided into morphological and behavioural groups by several glaciologists. As our review is limited to glaciation in tropical mountains we need to be concerned with only two types:€cirque and valley glaciers. A cirque glacier is smaller; it occurs when moving ice in the mountains is confined to an amphitheatre-shaped depression, known as a cirque. A valley glacier extends beyond the cirque, occupying at least part of a valley. Detailed discussions on glaciers of all types are available in various textbooks (Sugden and John, 1976; Bennett and Glasser, 1996; Benn and Evans, 2010).

13.3.2╇ The glacial mass balance A glacier is an open system. Ice is added to it by precipitation and lost from it primarily by melting. A budget for the mass of ice in a glacier explains its behaviour and movements. Ice is added by snowfall and rain that freezes on the glacial surface. Ice is lost in mountains by melting, sublimation and being blown off the glacial surface in strong winds. The addition of ice is known as accumulation and the loss as ablation. Accumulation mostly happens at the upper part of a valley glacier and ablation at its lower section. These two zones are separated by a line (or zone) of equilibrium where the gain and loss of ice is balanced. Thus a valley glacier is thickest at or around the line of equilibrium in the middle, and tapers off at both ends (Fig. 13.4). The glacial mass balance (also known as glacial mass budget) is computed by determining the difference between total annual accumulation and ablation. As expected, accumulation increases during both the winter and the season of precipitation, and ablation rises in summer. If the mass budget is balanced and accumulation equals ablation, the front of the glacier remains almost stationary, although ice moves through its body compensating ablation at its terminus. The advance or retreat of a glacier is directly related to the positive or negative mass balance. The rate of movement, both up and down the valley, is dependent of the relative proportion of budget surplus or deficit with reference to the dimensions of the glacier. A large glacier may not move much if the budget is slightly unbalanced. The

Tropical highlands

238

glacier, however, needs to overcome the friction at its base and the valley sidewalls in order to move.

13.3.3╇ The mechanics of glacial movement We can observe or measure the advance or retreat of a glacier, but our inability to see the ongoing, processes at the base of ice is a major problem in understanding glacial movement. Such processes have been inferred by deductive reasoning from the landforms and sediment left by past glaciers which are no longer in existence, or by inspecting the sedimentary evidence left as a glacier retreats upvalley. Activities at the base of glaciers have also been studied by sending instruments down boreholes sunk into the ice and by observations round the edge of glaciers. Ice, in order to move, needs to overcome (1) basal friction between the ice and rock or sediment at the valley bottom on which it rests and (2) internal friction between the ice crystals and layers within the glacier. Ice resembles a plastic substance in that it can resist a force up to a certain limit, beyond which it deforms. Ice deformation is caused when the ice is stressed (acted on by a force per unit area). There are two types of stress in a glacier:€hydrostatic stress and shear stress. Hydrostatic stress at a point inside the glacier is derived from the weight of the ice on top and is equal in all directions. Shear stress is a combination of the weight of the overlying ice and acceleration due to gravity determined by the slope of the glacier at its surface. A glacier moves according to the resultant between the shear stress and resistance to movement. Movement of a glacier carries two components:€internal deformation (also referred to as creep) and basal sliding. In most glaciers, basal sliding is responsible for about 80 per cent of its movement. Shear stress which causes internal deformation can be expressed as

τ = ρgh sinα,

(13.1)

where τ = shear stress ρ = density of ice g = acceleration due to gravity h = glacier thickness α = slope of the upper surface of the glacier. Shear stress and internal deformation therefore increase with surface slope and ice thickness, and also directly with depth inside the ice mass. Ice deformation is the mutual displacement of ice crystals relative to each other. Similar orientation of most of the crystals in a body of ice helps. The deformation can be summarised by Glen’s Flow Law of 1955:

ε = kτn

(13.2)

where ε = the strain rate k = a constant related to ice temperature, the value of k increasing with temperature τ = shear stress n = an exponent. (Nye, 1957).

Mechanics of mountain glaciation

239



Fig. 13.5

Ice movement Ice is deformed more at depth and at a relatively higher temperature. There are other models for internal deformation but Glen’s Flow Law is a robust and succinct summary of a part of glacial movement. In sum, ice in a glacier can be deformed internally by both shear stress and hydrostatic pressure. Problems arise when ice has to overcome an upvalley slope over an irregular bottom topography or the thickness of ice varies between parts of the glacier (Fig. 13.5). Nye (1952, 1957) suggested compressive and extending flows for overcoming such obstacles. When ice is compressed against an upvalley slope or ice thins as it nears the ablation zone, compressive flows occur with a slowing down of velocity. In contrast, extending flows occur where the bedrock slope steepens or the volume of ice increases. This increases the velocity of the ice. Both these cases of departure from the ambient velocity create a series of slip planes within the ice so that layers of ice can slide downvalley. In difficult cases where deformation cannot be adjusted within the body of ice, blocks of ice may fracture off and travel downslope, especially near the end of a glacier. Fractures produced by compression frequently open up cracks at the surface of the ice, known as crevasses. Glacial movement is also helped by ice slipping over bedrock on a thin layer of water. The water could be from meltwater or precipitation that has made its way to the bottom via fractures and openings in the ice. Water is also derived from ice melting near the base where the pressure-melting point has been reached. Temperature is higher towards the bottom of the glacier, as heat derived from radioactive disintegration of minerals coming up through the crust is trapped by ice. Furthermore, the water acts as a lubricant against basal friction. It also partially offsets the weight of the overlying ice. Uncertainty exists regarding the location of the basal water€– whether it occurs as a continuous film or is found only in cavities in the rock. The problem with irregularities on the rock floor of a glacier could be surmounted by basal water. If the irregularities are small in size, basal ice pressed against them melts and the meltwater flows round the irregularities to freeze on the downstream side, a process called regelation, thus transferring the mass of ice downslope. If the irregularities constitute a large barrier, the ice at the upper end of the obstacle comes under high pressure, deforms and passes round the obstacle. This is called enhanced creep. If glaciers are on soft sediment instead of hard rock, they can also move downslope by pushing and deforming the sediment. A mountain glacier therefore moves by a combination of internal deformation, basal sliding and sediment deformation. Higher up the peaks, in a much colder area, the movement could be entirely dependent on internal deformation.

Tropical highlands

240

(a)



Fig. 13.6

(b)

Idealised velocity distribution in a valley glacier

The average annual surface velocity of glaciers on the lower slopes of the mountains could be 13–200 m (Ritter et al., 1995). The velocity may be much slower on the colder upper slopes. A glacier flows fastest near the line of equilibrium where the ice is thickest. It is slower towards both the colder upper part and the lower section where the ice is thin and melting. Cross-glacier the fastest velocity is achieved in the middle, from where it drops off to the sides where the ice thins and friction becomes significant against the valley walls. Vertically, the fastest velocity is at the surface, but the vertical distribution of velocity is non-uniform (Fig. 13.6). Figure 13.6 is a diagrammatic and simplified presentation. Local variations in velocity always occur as departures from this general model. Variations in velocity also happen over time, for example, the velocity at the ablation zone increases in the summer, presumably because of a higher production of meltwater which assists basal sliding. The most striking short-term change is a glacier surge, which is a sudden, large-scale and quick transfer of ice down a glacier. The velocity in a surge could be as high as 1300 times the average figure. The surging glacier may persist for some time, up to several years, before returning to its average flow rate. During this period, a huge amount of ice and debris will be transferred downslope. This may result in river floods beyond the terminus of a glacier. Surging happens in several ways. It can be caused by: (1)╇the build-up of a large volume of ice in the upper part of the ablation zone and its subsequent collapse (2)╇position changes and interlocking of subglacial channels carrying meltwater enhancing basal sliding (3)╇ the accumulation of volcanic heat underneath glacial ice on volcanoes.

13.4╇ Glacial forms and processes Cirque basins and deep U-shaped valleys are glaciated erosional forms in the mountains. The depositional forms are usually limited towards the end or snout of a glacier. We will first look at the process of glacial erosion and the erosional forms, followed by depositional processes and the forms found lower in the mountains.

241

Glacial forms and processes

13.4.1╇ Glacial erosion Glaciers are active agents of erosion, capable of sculpting out huge basins and deep and wide valleys set in rock. Glaciers erode by abrasion, using rock fragments which become embedded at the base of the ice and scratch the underlying bedrock as the glacier moves. This requires a large supply of rock fragments that may range from sand to jagged boulders. The source of the rock fragments varies. They may be derived from frost-shattered rocks that fall on the ice surface from the valley walls above the glacier, material that drops down crevasses and pieces of rock that originate from the scratching of bedrock by fragments already embedded in the ice. These fragments may be transferred to the glacier surface, move inside the glacier or remain at its base. Glacial meltwater is also a tool, as it enters cavities and cracks in the rock and refreezes due to variations in pressure. This loosens and removes chunks of bedrock, especially if the original rock is highly jointed, fractured, had other lines of weakness or had been weathered before being covered by ice. This is called plucking. A glacier erodes primarily by abrasion and plucking at the rate of several millimetres a year. Faster rates of erosion are expected from glaciers that are thick and flow fast. Cirques and U-shaped valleys in the mountains are created by glacial erosion.

13.4.2╇ Cirques Glaciers start in cirques, which are deep and semi-circular erosional basins with a high headwall, a scooped-out floor and a raised sill at the lower end (Fig. 13.7). Cirques have been compared to an armchair, an armchair with a high back. In size, cirques range from small depressions cut into the hillsides to very large amphitheatres which could be several kilometres long, over a thousand metres wide and with a high headwall of a thousand metres. The slope of the headwall is hardly more than 30° for the lower part, but the upper section can be extremely steep, 70–80°. The sidewalls are commonly curved, but rectangular and triangular cirques have been reported in the literature. When a cirque contains ice, it usually covers the entire floor and stretches to the lower part of the headwall where a deep crevasse, called a bergschrund, occurs near the wall. At the other end, ice spills over the lip of the basin to flow down a valley. A number of cirques are now bare of ice, displaying the full dimensions of the rock basin. Water may collect on the floor of an ice-free cirque to form a lake called a tarn. The beginning of a cirque is still not fully understood but, once started, it expands by erosion. At the beginning, snow may accumulate in a small pre-existing depression to form firn and later ice. Rock fragments are released by plucking, and they move down to the bottom end of the depression to form a ridge called the protalus rampart. All the processes that operate at this stage are collectively known as nivation. A curved base identifies the formation of a nivation cirque. As more ice accumulates and the true glacial ice is formed, its erosive effect transforms a nivation cirque to a cirque glacier with the proper armchair shape. Erosion is carried out by headwall recession due to frost shattering and rock failures, and basal sliding of the ice already inside the cirque. As two neighbouring cirques increase their areas, the intervening ridge is eroded from both sides until it becomes a steep narrow wall known as an arête. Several arêtes converge

Tropical highlands

242



Fig. 13.7

Shape and origin of a cirque to a peak whose sides act as headwalls to several cirques. Such peaks are also known as horns. Nearly all high Himalayan peaks are horns.

13.4.3╇ Glaciated valleys A valley which carries a glacier or used to carry one is a glaciated valley or a glaciated trough. Such valleys display characteristic features. As ice flows out of a cirque it travels down a valley eroded into the mountains by a river. River valleys are steep, deep and narrow and circumvent noses of alternating spurs (called interlocking spurs). Glacial erosion transforms such valleys into wide U-shaped ones (Fig. 13.8). It has been suggested that the transformation takes place in two stages. In the first stage, the V-shaped valley is given a parabolic shape because of differential erosion at different parts of the valley. Valley side slopes, where the velocity of ice (hence shear stress) is higher, are eroded more effectively than the narrow bottom. In the second stage, zones of high and low erosion shift as the valley becomes parabolic in shape, and ultimately the entire valley sidewall comes under erosion giving rise to a U-shaped cross-section (Harbor et al., 1988; Harbor, 1992). Thus the characteristic form of a glaciated valley emerges:€ steep-sided, flat-floored and with all spur-ends truncated. The floors of the major valleys are lowered faster than those of tributaries, the effectiveness of erosion being dependent on the volume of ice. When the ice disappears, the rock floors of the valleys are exposed and the valleys are occupied by streams. The tributaries tend to hang, giving rise to waterfalls. Such valleys are known as hanging valleys. The longitudinal profile of a glaciated valley is not smooth but descends in a series of steps, with long treads at the back of the steps. Such treads may slope upvalley. Plucking continues to erode the bottom of the steps. As glaciers retreat upvalley and the floor is

Glacial forms and processes

243

(a)

(b)



Fig. 13.8

Valley glacier exposed, a string of lakes may occur in the middle of the valley where water is ponded between the steps and the reverse slope of the lower treads. Such lakes are known as paternoster lakes, as they resemble the beads of a rosary to an imaginative observer. Cirques, arêtes, horns and glaciated troughs form the glaciated topography in the Himalaya and Andes. An enormous volume of sediment removed from the mountains during the formation of such features is carried to the lower slopes where ablation is in operation. The processes of glacial erosion, transport and deposition continue as long as the glacier is active. Even after the disappearance of the glacier, a glaciated valley may continue to supply the mountain streams with a high sediment load from the stored material.

13.4.4╇ Glacial deposition The erosive power of glaciers in the mountains generates a large volume of sediment which is carried through the ice and deposited at various suitable locations:€the end of the glacier, at its sides and underneath the ice on top of the bedrock. It is necessary to recognise the difference between the sediment and the forms produced by its deposition. The general term for the glacial sediment (deposit) is drift. Glacial deposits commonly carry boulders. In the early days of glaciology, such boulders were mistakenly identified as deposited not by glaciers but drifting icebergs. The idea has been corrected but the term remains. Glacial drift varies in thickness, being best developed in areas beyond the tropics where ice sheets once prevailed during the Pleistocene. We are dealing specifically with glaciations in the tropical mountains where the drift is thinner and restricted to valley floors and near the terminus or snout of the glacier. The character of the deposit varies depending on whether it has been deposited only by ice, by a combination of ice and meltwater or reorganised completely by meltwater after deposition. As the glacier advances and retreats along a valley, the sediment is considerably

244

Tropical highlands

modified. Two broad classes of drift are recognised:€non-stratified and stratified drift. Nonstratified drift is primarily deposited by ice, whereas in stratified drift the sediment is first deposited by ice and then modified by water. Hence stratified drift exhibits sorting and stratification, differentiating it from non-stratified drift. Non-stratified drift is also known as till or by an old term, boulder clay, which seems to have fallen out of favour but which describes the wide textural range of till. Till is extremely variable in character, but shares a few common properties:€absence of stratification, absence of sorting, striation marks on clasts, grain angularity and a generally bimodal texture. Some rounding of clasts, however, happens if they are transported over a distance embedded in basal ice. In contrast, clasts that have fallen from the valley walls and been carried on top of glacial ice are angular. Stratified drift involves transportation by both glacial and fluvial processes and is recognised as a fluvioglacial deposit. It displays limited sorting, stratification and rounding of clast edges and corners. Stratified drift is deposited by meltwater channels inside the glacial ice, inside localised depressions and openings as between the rocky valley walls and glacial ice, and in a washed-out form in front of the glacier. Both types of drift produce a variety of depositional forms of varying size and internal structure. We only review the forms commonly seen in the mountains:€a full coverage of glacial deposition is available in textbooks on glaciology, which indicate many of the other depositional forms associated with ice sheets at higher latitudes. The depositional forms could be ice-contact features, i.e. formed in contact with the sides and front of the glacier, or proglacial features, which are found in front of the glacier deposited by meltwater issuing from the snout of the glacier. Glaciers tend to move forwards and backwards at the lowest end in response to variations in accumulation and ablation. As a result, ice pushes and thrusts sediment into ridges and mounds, while the zone of proglacial forms also shifts. The zone near the glacial snout is characterised by a complicated assemblage of a range of depositional forms. Perhaps the commonest depositional form in mountain glaciations is the moraine, a ridged glacial deposit that does not follow the local topography (Fig. 13.9). However, ground moraine, which is areally extensive, is not a ridge. As ice disappears, glacial debris settles on the bedrock floor of valleys to create a depositional sheet, which is ground moraine. Ground moraine is pushed into ridges several metres high near the edges of a glacier. If it happens at the side of a glacier, it is called lateral moraine. Where two glaciers join, two adjacent lateral moraines come together to create a ridge in the middle of the combined ice to create a medial moraine. Moraine ridges, however, are best developed at the end of a glacier where several ridges may be found; these are collectively known as the end moraines. The end moraine that marks the furthest advance of the glacier is the terminal moraine, which curves round the edge of the glacier to link up with the lateral moraines (Fig. 13.9). All other end moraines are grouped together as end or recessional moraines, marking locations where the glacier has stopped temporarily while retreating and a moraine ridge has been formed. A subsequent advance may destroy moraine ridges, at least partially. End moraines are also modified by meltwater issuing from the glacier. Ice-contact features are formed at the side or snout of a glacier ending in a terminal zone of thin and slowly disappearing ice. These features will be in stratified drift and formed by

Glacial forms and processes

245

(a)

(b)



Fig. 13.9

Moraines meltwater streams. Deposition by meltwater may happen in channels on top of the glacier, inside or under the ice, in a linear gap between the rocky sidewall and glacial ice, and near a mass of stagnating and melting ice. Once the ice disappears, the sediment resettles and develops structural contortions. The final forms of these ice-contact features are mounds of stratified sand and gravel, collectively known as kames. Kames could be 102 m long and 10€m high. The form of a kame follows the depression in which meltwater had collected and sediment deposited. If deposition occurs in a linear cavity, kames resemble ridges. A deltalike feature is formed when a meltwater stream flows out of the ice depositing sediment. This is the kame delta. When deposition occurs between the rocky sidewall and the ice, the sedimentary form is left with a steep side where the ice used to be. This is the kame terrace. Kames frequently occur on the valley floor with depressions called kettles, constructing the kame and kettle topography. Kettles are formed when a mass of ice is trapped in the ground sediment and melts. Kettles may form small lakes. Long meltwater channels through the body of ice may deposit a linear winding body of sediment which is revealed when the ice melts. This is the esker. Eskers, however, are best developed from ice sheets. Proglacial features are found beyond the terminal moraine, where large quantities of meltwater issue from a glacier carrying a high sediment load. The sediment is deposited both inside the meltwater channels and overbank, building up a plain of washed-out glacial sediment. This plain built by glacial outwash is known as a sandur, an Icelandic term. The channels tend to continue beyond the sandur as the headwaters of a river system. A sandur deposited within a valley bounded by rock walls is known as a valley sandur (valley train in North America). If it terminates in a plain, as happens in the higher latitudes, it is known as a plain sandur (outwash plain in North America). Several levels may be seen in a sandur, related to several advances and retreats of the glacier upvalley. As expected, sandurs display sorting, with coarse grains being deposited near the glacier and finer material downvalley. Morphological changes also occur longitudinally. A sandur carries one or a limited number of channels and a pitted appearance from kettles at its upper end. Downslope, channels become multiple and braided with a wide and shallow cross-section.

246

Tropical highlands

These channels are mobile and a number of abandoned channels mark the surface of the sandur. At the very distal end, the channels become very shallow and, when in flood, the entire sandur could be under a sheet of water. Large meltwater floods modify the morphology of sandurs periodically. The morphology of high tropical mountains is thus modified by glaciations giving rise to a craggy appearance. But the effect of the glaciations extends further in (1) controlling the hydrology of streams originating in the mountains and (2) being an important source of sediment transported by these streams. The valley floors in tropical high mountains are characteristically covered with coarse clastic sediment that not only forms bars inside mountain channels, but also floodplains, terraces and fans overbank. Sediment is also derived from mountain slopes in mass movements. A large volume of sediment arrives in rivers from the glaciated higher slopes.

13.5╇ Slopes and valley floors in high mountains Below the high glaciated mountains, slope failures are a major supplier of sediment to streams. The large mountains are tectonically unstable due to their association with convergent or strike-slip plate boundaries. Slopes fail periodically due to such tectonic movements and also following high rainfall events. If these are volcanic mountains, then pyroclastic eruptions and lahars bring down vast amounts of material along river valleys. Brunsden et al. (1981) described the geomorphological character of the lower slopes of the Himalaya in Nepal. The terrain is a mosaic of deep trunk valleys, long and steep (20–50°) side slopes and a dense network of steep V-shaped tributary valleys. The major river valleys include channels laden with sediment, fragments of successive river terraces, tectonically elevated and variably tilted fans, and boulder fields backed by major scars on the hillslopes. The slopes also display a network of gulleys, especially if deforestation and agricultural practices are common. Although mass movement is the major denudational process, considerable sediment is brought down to the major rivers by surface and subsurface water erosion. High-magnitude floods in the trunk valleys tend to move most of the sediment at intervals. This general description holds true for most of the high mountains in the tropics.

13.5.1╇ Mass movements in the high mountains Shroder and Bishop (1998) attributed the Himalaya with being the ideal laboratory for the study of high-magnitude slope failures at both low and high frequencies. They indicated a general relationship between altitude and types of failure in the Himalaya. Avalanches tend to be common between 9000 and 4000 m, rockfalls and slides between 7000 and 2000€m, debris flows between 4000 and 2000 m, and mudflows and earthflows between 3000 and 1300 m. In general, the failed material is finer and wetter downslope. Such failures contribute an enormous amount of sediment. Earthquakes in mountains also contribute episodic instances of large quantities of sediment to streams. Goswami (1985) recorded repeated

247

Rivers in the tropical mountains

additions of extra sediment into the Brahmaputra River from tectonic slope failures in the eastern Himalaya. The massive slides and flows may block rivers in the mountains, giving rise later to damburst floods. A commonly used example of such an event is the damming of the Indus in the winter of 1840–41 by the earthquake-caused collapse of a section of the Nanga Parbat in the river. The lake formed upstream of the barrier was 64 km long and over 300 m deep. The dam was breached the next June by the Indus and a tremendous flood swept down the river. Rising to 30 m in height, this flood washed away part of a Sikh army camping on the floodplain at Attock, 400 km downstream of the dam. Apart from episodic seismic disturbances and the long-term general degradation of steep slopes leading to collapse, mass movements that arise out of high-intensity rainfall in large storms erode mountain slopes and bring an enormous amount of sediment to the valleys. Starkel (1972) described one incident of high-magnitude rainfall from a Bay of Bengal tropical cyclone that, during 2 to 5 October 1968, deposited 700–1130 mm of water in the eastern Himalaya, near Darjeeling. The valley flats of the Tista and Great Rangit rivers were blanketed with very coarse material including a high proportion of boulders. Such sediment is deposited not only inside the channel or on top of the valleyflat surface, but also accumulates at breaks in slopes to create alluvial fans that may grow to huge dimensions. The material contributed by lahars can be very large, e.g. following the 1985 eruption of Mount Huascaran in the Colombian Andes or periodically from Mount Merapi in central Java (see Chapter 14). The huge amount of material brought down the mountain slopes is stored in the valleys as fans, terraces and floodplains and also within channels as bars of various sizes (Fig. 8.6). As the rivers emerge from the mountains, they deposit a considerable amount of this sediment at the break in slopes at the highland–lowland contact to build fans (Bruijnzeel and Bremmer, 1989). Sediment is carried further downstream and stored temporarily in channels and on floodplains in mixed bedrock and alluvial valleys.

13.6╇ Rivers in the tropical mountains From the first-order headwaters to the major rivers that emerge from the mountains, the general function of mountain streams is to convey water and coarse sediment over a steep gradient to the plains. According to Douglas and Guyot (2005), tectonically active regions provide very high sediment yield, of the order of 104 tkm−2yr−1. The nature of the sediment depends on the regional geology, which includes active subduction zones, collision belts, volcanic arcs, granitic plutons, rift valleys and uplifted sedimentary rocks (Scatena and Gupta, in press). Many of the rivers carry runoff derived not only from glacier meltwater in late spring and early summer, but also from a very large amount of annual precipitation that may reach thousands of millimetres. The effect of this large rainfall is even more pronounced in the seasonal tropics, where about 80 per cent of the annual rainfall may arrive in 4–5 months (Gupta, 1988). The steepness of the channel gradient and the coarseness of the bed and suspended loads may leave their imprints on the channel forms in the mountains by enhancing the seasonality in form and function of river channels.

248

Fig. 13.10

Tropical highlands

The Swift River in the Blue Mountains, northeastern Jamaica displaying the characteristic features (see text) of a montane river. Photograph: A. Gupta Episodic rainfall, as from tropical storms, may be very high. As expected, such rainfall gives rise to large floods capable of transporting huge volumes of sediment downvalley. Such sediment is derived from previously stored material in the valley and new material that has moved en masse down the steep mountain slopes as a result of the storm. The effect of these rainfall events could be catastrophic as the big storms may shed rain over an entire river basin within a steep mountainous area. The steep slopes, saturated soils and intense rainfall common in the mountains lead to high-discharge events, large floods, sediment transportation and channel modification. The floods are powerful enough not to allow long-term log jams and accumulation of coarse wood debris (CWD) in the channels, as happens in the temperate montane streams. In tropical streams, such wood debris that arrives following landslides and large floods is usually removed in months or within a couple of years. Bare solitary tree trunks deposited on a channel bar in a flood are more common. Scatena has observed that although several CWD dams produced by hurricane defoliation and uprooting may last for five years in the Luquillo Mountains of eastern Puerto Rico, the majority of these barriers are broken up and redistributed within less than six months. CWD dams have not been observed in streams higher than second-order (Scatena and Gupta, in press).Wohl et al. (2009) concluded that in the Upper Rio Chagres Basin of Panama, large wooden debris produced by floods and landslides lasts for two years or less. The fluvial system seems to alternate between short periods of stored wood in the valleys and long periods without. Wohl and Merritt (2005) suggested that montane streams have at least a gradient thresÂ� hold of 0.002 m m−1. The gradients of tropical montane streams are commonly well above this threshold. Their longitudinal profiles tend to be punctuated by waterfalls and alternate reaches of step and gentle gradients determined by bedrock (Hancock et al., 1998). The rivers carry the usual features of montane streams:€bedrock channels, boulder bars, boulderlined channels, step-pools and pool-riffle sequences (Fig. 13.10). A veneer of alluvium and boulders cover the bedrock, and channels are in many places best described as mixed

249

Sediment from tropical mountains

bedrock-alluvial channels (Whipple, 2004). A thick cover of alluvium is rare. In places, channels may be full of boulders. Such boulders could be either exhumed corestones or they could be transported to the channel by mass movements, especially debris flows, in large storms (Ahmad et al., 1993). Terry (1999) described the passage of sediment in an oceanic island with high volcanic relief that is subjected to high rainfall events. This account can be taken as representative of sediment derived from similar island environments. Burbank et al. (1996) commented on the poorly understood competition between bedrock uplift and erosion in tectonically active mountains. They refer to the extreme high annual incision rate of the Indus River (2–12 mm) in its middle gorge near Nanga Parbat, Pakistan. An equilibrium is maintained in this tectonically active region between bedrock uplift and river erosion, with slopes adjusting with failures to the fast downcutting of rivers. Gabet et al. (2008) studied the relationship between high rainfall and rapid erosion in the much wetter Nepal Himalaya using a network of ten river monitoring stations. They conclude that the average erosion rates increase with discharge and precipitation. They also suggest that long-term rates in parts of the Himalaya tend to reflect the effect of past glaciations and not the present precipitation pattern. Interestingly, working in Andes catchment, Aalto et al. (2006) concluded that erosion rates are directly related to the average basin slope and relief rather than runoff. Gabet et al (2008) suggested that short-term rates peak on local conditions, such as high precipitation during the present or intense glaciation of the past, but such variations average out over time to provide a regional uniform rate for the Nepal Himalaya. These studies probably indicate that a large volume of sediment may emerge from high tropical mountains with elevated rates localised over time and space. The mountains need to be studied in greater detail.

13.7╇ Sediment from tropical mountains As Figure 13.10 illustrates, a huge amount of material comes off the slopes in high mountains and is stored as fans, terraces, floodplains and channel bars within the river valleys. Glaciers and slope failures produce most of the sediment which is redistributed on the valley floor by mountain streams. Volcanic material is added to the assemblage where volcanoes are present, as in the Andes and the islands of Indonesia and the Philippines (Chapter 14). A very large proportion of material coming off tropical mountains is stored on the lower slopes to build alluvial fans, especially if the contact of the mountains with the plains is sharp. This tends to happen in tectonically active areas and case studies on alluvial fans have been reported from tropical South America, Central America, northern India, Papua New Guinea, northeastern Australia and possibly other locations. A series of large alluvial fans, including several especially dynamic ones, extend along the entire length of the Himalaya (Fig. 9.6) and the Andes. These are known as megafans and are differentiated from the common alluvial fans (Leier et al., 2005) by their large size (103 km2 or greater), low gradient (commonly < 0.1°), typical sedimentary texture (boulders at the apex and predominantly silt and mud at toes) and depositional process (mostly by wandering channels

250

Tropical highlands

of a single river). Horton and DeCelles (2001) expect megafans to drain a catchment area of about 104–105 km2 and to have a depositional area of 103–105 km2. Leier et al. (2005) stated that although alluvial fans may occur anywhere at breaks in the slope, fluvial megafans are found within 15–35° of latitude in both hemispheres, indicating an association with seasonal discharge, large sediment and a mountain front facing a large depositional basin that allows space for the megafans to form. The modern fluvial megafans, of the Rio Grande, Rio Parapeti and Rio Pilcomayo, emerge from the central Andes in Bolivia to debouch onto the low-relief Chaco Plain. These measure 5800–22 600 km2 (Horton and DeCelles, 2001). Alluvial fans of limited size and catchment areas occur between two neighbouring megafans. The present active channel for these fans has a straight pattern. Sedimentation in a channel occurs as midchannel and side bars of sand, whereas overbank deposition is characterised by crevassesplays and fine sediment. Abandoned divergent channels, suggesting frequent channel avulsions, are found overbank and the main channel of a megafan remains stable only for a brief period. Other Andes megafans occur towards the south in Argentina. The mobility of the main channels may be related to large-scale river capture or lateral shifting in large floods (Baker, 1978; Horton and DeCelles, 2001). Megafans are important stores of sediment off tectonically active areas in the seasonal tropics. Their construction is probably due to short-term deposition in one location followed by avulsion, which shifts the centre of deposition to a new location. Smaller and commoner alluvial fans develop in the tropics where the necessary attributes for megafan formation are absent. Fine-textured sediment, low frequency or the absence of tectonic movements, and a non-seasonal river discharge lead to the formation of a number of small steep fans, for example, below the Main range of the Malay Peninsula, at the base of sandstone escarpments in tropical Australia and at the foot of Indonesian volcanoes. The best-known of the Himalayan megafans (Fig. 9.6) is that of the Kosi River. A number of studies have described the properties of the Kosi Fan (Gole and Chitale, 1966; Singh et€al., 1993; Wells and Dorr, 1987). The fan measures 154 × 147 km and longitudinally slopes 0.89–0.06 mkm−1. The cross-fan slopes range from 0.1 to 0.25 mkm−1. The area is affected by the Indian monsoon system and the regional annual rainfall averages 1300–1800 mm. Over 80 per cent of this arrives between June and October. Individual large storms and considerable variability in the annual total are both superimposed on this general pattern of rainfall. Large volumes of clastic sediment are delivered by the Kosi from the Himalaya. Singh et al. (1993) have recognised four distinct zonal distributions of facies in the deposits of the Kosi megafan:€zone 1 (gravelly, sandy, braided channels), zone 2 (sand, braided channels), zone 3 (fine sand and mud, straight channel) and zone 4 (fine sand and mud, meandering channel). Zone 4 has by far the longest reach. The deposits consist of multistoried sand sheets interbedded with overbank muds. Individual sand sheets are commonly 8–10 m thick and multistoried bodies measure 16–20 m. Sand is replaced upstream by gravel. The Kosi has shifted 113 km westward in 228 years in episodic jumps of up to 19 km a year. The reason for such shifting is not clear, although a number of explanations have been put forward. Tectonism is a common cause of channel shifting in the Himalaya. Large floods due to various factors have also caused river channels to shift. Shifting in floods is

251

Conclusion

feasible, as during the wet monsoon the Kosi turns into a 15–30 km wide river. The concept of a 15 km wide river avulsing is difficult to comprehend. Unlike a number of rivers, however, the Kosi shows a low statistical correlation between channel shifts and dates of earthquakes or high-magnitude floods. Wells and Dorr (1987) suggested that the channel of the Kosi shifts in annual floods in a random fashion, enabled by the flat gradient of the fan. In 2008, following a major flood, the channel of the Kosi surprisingly shifted towards the east. The Tista River that flows from the Himalaya to the Brahmaputra has built an even bigger megafan of 18 000 km2 that changes its slope from 0.19° at the apex to 0.01° at its distal end. The slope across the fan is gentle as expected, about 0.01°. The river, like other Himalayan rivers, demonstrates a pronounced seasonal pattern of discharge with an average annual discharge of 609 m3s−1 and the highest average monthly discharge above 2000 m3s−1. The fan displays several lobes, a radial framework of current channels and a network of palaeochannels (Chakraborty and Ghosh, 2009). The main rivers of the Himalayan megafans flow into the Ganga as a well-integrated fluvial system. Flowing into a major river, however, is not a global characteristic. The Andean megafans tend to end in swamps and depressions. Megafans and large-scale swampy depressions are also related in the Okovango, Southern Africa (Stanistreet and McCarthy, 1993). In such conditions, huge amounts of sediment remain stored in continental interiors and do not reach the sea. Horton and DeCelles (2001) mentioned the deposition of a huge amount of clastic sediment that is derived from the Central Andes between 18–22° S latitudes on the Chaco Plain, primarily forming megafans. Sediment that bypasses megafans finally reaches large rivers that collect the drainage from high mountains. Sediment grains ultimately reach the sea, but in large rivers it may take any length of time up to thousands of years to do so. Meanwhile, sediment is stored in channel bars, floodplains and terraces along the valley. Several cycles of storage and transfer could be required before an individual grain reaches the coast, where it contributes to delta building or is deposited on a coastal shelf (Meade, 2007). Very little sediment is derived from cratonic highlands in hard metamorphic rocks, although tributaries draining such highlands may contribute a major part of a trunk river’s discharge (Chapter 9). Isolated volcanoes may produce a considerable amount of sediment for local rivers, mainly from pyroclastic flows and lahars (Chapter 14). Given the near-sea location of subduction-zone volcanoes, a large part of volcanic sediment reaches the coast quickly and tends to build up coastal features such as beaches and dunes. Volcanic material draining east from the Andes is an exception. Such sediment travels a long distance along the Amazon, like sediment of any origin.

13.8╇ Conclusion Tropical mountains are the source of a large volume of water and almost the entire bed load and suspended load of rivers that emerge from such high areas. The high sediment load from the mountains is due to a combination of factors:

252

Tropical highlands

• glaciations • extensive and frequent slope failures • high and intense rainfall • tectonics • regional lithology and structure • volcanic eruptions. The sediment travels episodically. A large part of it probably stays stored for decades and centuries in the foothills, forming fans and, further down the valley, forming floodplains. In the case of large rivers, e.g. the Amazon or Ganga, sediment could remain stored for thousands of years. Meade (2007) described the passage of the post-Pleistocene sediment of the Orinoco and the Amazon from the eastern slopes of the northern Andes Mountains to the Atlantic Ocean. Much of the Andes-derived sediment is stored for 102–103 years or longer in the wide alluvial plains of the Andean foreland basin or the huge floodplains along the mainstem of the Amazon and the Orinoco and a limited number of its large tributaries. Meade attributes about 90 per cent of the sediment of the Amazon to the Andes. He is of the opinion that the rest comes from the Andean foreland basins. In that sense, all of the Amazon’s sediment is derived from a high tectonically active tropical mountain range. The source of the sediment transported by rivers also depends on the geology of the highlands. The suspended load of the Nile is relatively modest, given its size. Most of the load (72 per cent) arrives via the Blue Nile during the wet monsoon from the steep, volcanic Ethiopian Highlands. Relatively little is derived from the Central African Highlands where the White Nile rises, and much of this load is then lost as the river traverses the flat swamp of the Sudd (Woodward et al., 2007). Relative contributions of sediment from different parts of the drainage basin can be identified from chemical analysis of the isotopes of Sr, Nb and Os in the sediment. It has been shown that almost all the sediment of the Brahmaputra River comes from the Himalaya and 45 per cent from the part of the eastern Himalaya known as the Eastern Syntaxis, where the upper Brahmaputra makes a U-turn in a deep gorge around the Namche Barwa Peak in a region of heavy monsoon rainfall and tectonic instability (Singh, 2007b).

Questions 1. Using Figure 13.1, describe the transfer of sediment from the volcanic slope of Merapi to the sand dunes of Parangtritis in Java. 2. What are the depositional features associated with mountain glaciation in the tropics? Do these act as significant sources of sediment to the regional rivers? 3. Read the description of the lower Himalayan slopes in Nepal by Brunsden et al. (1981). Explain the origin of the morphological features described.

253

Questions

4. Shroder and Bishop (1998) described the Himalaya as the ideal laboratory for the study of high-magnitude slope failures at both low and high frequencies. Why? 5. Describe the general character of tropical montane streams. 6. What are the form and function of megafans? 7. Tropical mountains act as the source for a large volume of water and almost the entire bed and suspended loads of large rivers that start in the mountains. Why?

14

Volcanic landforms

The houses, people, traffic seemed Thin fading dreams by day, Chimborazo, Cotopaxi They had stolen my soul away! W. J. Turner

14.1╇ Introduction The material discussed in Chapter 2 leads us to conclude that volcanoes occur in specific locations. Such locations include convergent plate margins (for example, volcanoes of the central Andes), islands that form higher parts of mid-oceanic ridges (Ascension Island in the Atlantic) and volcanic islands formed over hot spots in the middle of the oceans (Mauritius and Réunion) or continents (Tibesti, Chad). Volcanoes also occur in rift valleys. Mounts Kenya and Kilimanjaro of the East African Rift are probably the best-known of the rift valley volcanoes in the tropics. A number of less well-known ones, such as Nyamuragira and Nyiragongo of the Virunga Mountains, Central Africa also form prominent features on the regional landscape. Flood basalt erupted on a continental scale in the geologic past over parts of the tropics such as the Deccan Plateau of India and the Drakensberg of South Africa, giving these areas their characteristic landscape (Fig. 14.1). In all these areas, the landscape is primarily determined by the lithology and relief created by volcanism. The pattern persists even where geomorphic processes have considerably altered the original rocks (Fig. 5.2B). Volcanic landforms are therefore an integral part of the tropical environment. Almost 70 years ago, Cotton (1944) wrote a textbook on geomorphology indicating the importance of volcanoes in the title of his book. To an extent, the book focused on New Zealand, where volcanoes are common in certain locations such as the TaupoVolcanic Zone. An account of tropical landforms also requires a discussion of the volcanic landscape for the same reason. We discuss the volcanic landforms of convergent plate margins and flood basalts.

14.2╇ Types of volcano and the related landscape

254

A volcanic landscape is one of high relief but with a variety of forms. The type of volcanic form indicates its association with plate tectonics and lithology. Four broad types of

255



Fig. 14.1

Types of volcano and the related landscape

Distribution of volcanic landforms in the tropics volcanic landform can be identified following the nature of the eruption involved:€monogenetic volcanoes, polygenetic volcanoes, shield volcanoes and calderas. A monogenetic volcano is built by a single eruption; a polygenetic volcano is created by more than one eruptive event; and a shield volcano is a large but gently sloping volcano, almost exclusively basaltic, and commonly originating from a hot spot (Francis, 1993). Calderas are formed by catastrophic events as described below. The following account is to a large extent, based on Francis (1993) and Verstappen (2005).

14.2.1╇ Monogenetic volcanoes Monogenetic volcanoes tend to build relatively small single cones of scoria or tuff and small surface craters. Scoria refers to relatively light volcanic rocks with a fragmented, cindery vesicular texture. Scoria cones are usually made of basalts, and are commonly about 800 m in diameter and 200–300 m in height. They have relatively large craters for the diameter of their edifice. Scoria cones are not always symmetrical, as they extend along fissures or the ambient wind direction during eruption. If the magma comes into contact with water during the eruption, shallow explosions blow out surface craters instead of building cones. Such depressions later become lakes called maars. These circular depressions are surrounded by low rims of ejected material. Tuff rings are similar features but they are built by the accumulation of eruptive material in a circle above ground; they are not craters. Hot, water-saturated material builds tuff cones, which resemble scoria cones morphologically but are steeper and smaller. In contrast, scoria cones result from dry explosions. Once formed, monogenetic volcanoes are subjected only to erosion, whereas polygenetic volcanoes may also experience several new explosions over time, rebuilding the volcano and changing its appearance. Erosion is not the sole remodelling process.

14.2.2╇ Polygenetic volcanoes Polygenetic volcanoes are much bigger and vary in their morphology and structure. Francis (1993) listed simple cones, composite cones, compound cones and volcanic complexes. Simple cones are scoria cones but with a history of multiple eruptions. They appear symmetrical with a single summit conduit and a small crater. Several well-known volcanoes

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Volcanic landforms

fall into this category:€Mount Mayon, Philippines (2400 m) and El Misti, Peru (5822 m). Volcanoes with a central vent have simple structures but eruptions may take place from both the central vent that reaches the summit and lateral vents that feed eruptions on volcanic slopes; the latter phenomenon is known as the flank, parasitic or satellite eruption. This happens in large volcanoes. Feeder eruptions may take place anywhere, but they are very common in areas of crustal extensions or rifting and locations over hot spots. Composite cones generally reflect volcanic activity at the same site, and hence the edifices have a radial symmetry, but multiple eruptions give them a complex history of volcanism. Compound volcanoes do not have individual cones that dominate, but their massifs cover a large area which will include several cones, domes and craters. Volcanic complexes are even more complicated. They incorporate an assemblage of several lava flows, volcanic domes and pyroclastic rocks that are connected over space, time and origin, showing vent migration over time. The Nevado Chuchuan Volcanic Complex near El Misti in Peru is a good example. The form of stratovolcanoes (volcanoes built by stratum over stratum of volcanic material) may have complex forms due to lava outflows in mid-slope, flank eruption craters and different slope processes operating over the middle or lower slopes. Faulting also results in the alignment of eruption openings and thereby the shape of the volcano. Calderas are conspicuous features in a volcanic landscape (Fig. 14.2). They result from violent eruptions that are accompanied by the collapse of the roof of the magma chamber or blowout of a huge amount of material from the upper part of the volcano. Calderas can be 2–20 km in diameter, with a circular or oblong appearance. The Toba caldera in north Sumatra is an oblong caldera, elongated along a strike-slip fault.

14.2.3╇ Shield volcanoes These are huge convex volcanoes (40 000–80 000 km3 in volume) with gentle (<10°) upper slopes near the summit. Shield volcanoes are made of a large number of individual basaltic lava flows and owe their origin to a hot spot underneath. The Hawaiian volcanoes are good examples, especially Mauna Loa and Mauna Kea that rise more than 4000 m above the surface and a total of 9000 m from the floor of the Pacific. Built on a thin oceanic crust, such massive edifices tend to sag into the aesthenosphere (Francis, 1993). Huge landslides may modify the shield volcanoes morphologically, as is happening on the southeastern flank of Kilauea, Hawaii. The Galapagos Islands off Ecuador also provide good examples of shield volcanoes.

14.2.4╇ A comparison The dimensions and slopes of volcanoes depend on the composition of the construction material and the eruptive stage. Volcanoes constructed of oceanic basalt tend to be large and with gentle slopes. Andesitic and dacitic volcanoes of the subduction zones and island arcs tend to be steeper and shaped by explosions due to the viscosity of their silicic and alkaline magma. The appearance of volcanic cones is also dependent on the erosion that follows the building of the cone, especially from rain falling on fresh volcanic material (Fig. 14.3).

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Fig. 14.2

Lava and pyroclastic deposits

The crater of the stratovolcano of Raung in east Java is about 2 km across. The volcano is high (3332 m) and erupts regularly; it has already been active several times in the twenty-first century. Flow and explosive deposits and screes of volcanic material mark the crater rim and the caldera floor. A second crater with a deep volcanic vent rises asymmetrically from the floor of the caldera. IKONOS satellite image© Centre for Remote Imaging, Sensing and Processing, University of Singapore (2009), reproduced with permission. See also colour plate section V-shaped gullies, separated by ridges, cut into slopes displaying a radial pattern away from the summit, give a pattern termed parasol-ribbing. A radiating drainage network with channels tens of metres deep develops rapidly, especially in tropical conditions with frequent high-intensity rainfall as discussed in Chapter 3. This network is often very effective in transferring volcanic material from the edifice to the surrounding area. A set of volcanic landforms therefore not only incorporates individual volcanoes providing high relief but also material erupted, eroded and transferred away from them, modifying the surrounding landscape. The material is transferred by a number of processes:€volcanic flows, mass movements and running water in channels down the slopes. If the coast is near, as in Indonesia, the Philippines and the volcanic islands of the Caribbean, volcanic sand becomes an important source of material for coastal landforms (Fig. 13.1). We need to examine the processes that transfer volcanic material downslope and the nature of such material.

14.3╇ Lava and pyroclastic deposits The morphological features of a volcanic surface depend on the nature of the lava and pyroclastic deposits involved. Basaltic lava flows are usually grouped into two classes:€aa and pahoehoe, both names derived from Hawaii. Aa lavas are more viscous, with a surface

258



Fig. 14.3

Volcanic landforms

Erosion on fresh volcanic deposits covering the slopes of Anak Krakatau (the child of Krakatau) growing at the same location where the blown-apart volcano of Krakatau existed until 1883, between Sumatra and Java, Indonesia. Photograph: A. Gupta of irregular, sharp-edged blocks. Individual lava flows have two parts:€an upper part which cools fast into rubbles and a lower part with a massive structure that cools slowly over a scoria-rich base. Vesicles are seen in the lower part due to the escape of gases from the lava. Pahoehoes are low in viscosity and flow fast to build a smooth, shiny, rope-like surface. Andesitic lavas are more viscous than basaltic lavas. They travel a short distance from the erupting vent, giving rise to large smooth-sided blocks on the surface. Dacites and rhyolites form stickier lavas and are relatively rare, especially rhyolites. The surface expressions of such lavas are often plugs and domes that rise vertically above the ground surface. Pyroclastic deposits are products of volcanic explosions. The explosive act in a volcano includes three stages. First, the magma is fragmented by coalescence and the bursting of bubbles arising through it; second, the erupting mass is blasted through the vent to the crater; and third, the erupted material rises as a column on top of the crater or spills over its sides. The eruptions could be sustained jets (Plinian eruptions), short blasts separated by near-instantaneous or longer intervals (Strombolian eruptions) or discrete blasts at intervals (Vulcanian eruptions). The erupted material rises in a column high above the crater that could reach tens of kilometres and enters the stratosphere before spreading radially to an umbrella-shape that could be symmetrical or extended asymmetrically downwind. The rate of eruption, and hence the eruption column itself, is controlled by the volatile content of the magma, its exit velocity and the radius of the vent. The cloud consists of tephra. Tephra is a general term for pyroclastic material that falls to the surface from an eruption cloud. A rude reminder of the solid composition of this ash cloud is occasionally provided when a jet aircraft flies through it and loses engine power. Volcanic eruptions also disrupt air traffic schedule from time to time as after such experiences aircrafts normally try to avoid eruption clouds. Most of the deposit falls to the ground as a grey blanket of finegrained material that varies in thickness, decaying exponentially away from the vent, and

259



Fig. 14.4

Lava and pyroclastic deposits

Eruption column over Pinatubo, 12 June 1991. The convective rise and the umbrella region at the top are visible. Photograph by Dave Harlow, US Geological Survey in many instances covering a wide area, but part of it may transform into flows by contact with water along channels. Tephra is divided into three size terms. If it is less than 2 mm it is described as ashes, between 2 and 64 mm it is known as lapilli (little stones) and material coarser than 64 mm is called blocks. Compositionally, tephra includes material from both the rising magma and rocks from the walls of the vent, which are not necessarily volcanic. There are several types of pyroclastic deposit.

14.3.1╇ Ash fall deposits Major air-fall deposits result from Plinian eruptions. Plinian eruptions are huge eruptions where high sustained exit velocities from the vent lead to strong convective columns that rise into the stratosphere. Plinian tephra usually consists of pumice clasts of dacitic to rhyolitic composition, and Plinian eruptions generally occur at converging plate margins where silicic material is available from the continental crust. Exceptions, however, do occur. Pinatubo Volcano in Luzon, Philippines (Fig. 14.4) erupted in June 1991 with ash columns that rose for about 40 km (Newhall and Punongbayan, 1996) with the paroxysm peaking on 15 and 16 June. The eruption, however, lowered its summit elevation from 1745 m to near 1400 m, and left its top with a 2 km wide steep-sided caldera. The ash fall buried more than 4000 km2 of area with tephra. The volume of the tephra has been estimated to be between 8.4 and 10.4 km3, most of it loose sand-textured andesite–dacite fragments. The tephra fall and pyroclastic flows filled a number of radial valleys, overtopped the divides and created a new topography. The loose material is mobilised after rainfall to create dangerous mudflows or lahars. The volcano has not erupted like this for several hundred years (Francis, 1993; Newhall and Punongbayan, 1996; Nossin, 2005).

260

Volcanic landforms

14.3.2╇ Pyroclastic flow deposits Pyroclastic flows are mixtures of rock fragments and gases that descend volcanic slopes hugging the ground during explosive eruptions. They may originate from several causes:€partial or complete collapse of an eruption column, collapse of an active lava dome or flow and spilling of ash from an active crater. Pyroclastic flows travel at the high Â�velocity of 10–65€ms−1 and reach a temperature between 100 and 800°C (Blong, 1984; Cas and Wright, 1987; Francis, 1993). Pyroclastic material ranges from vesiculated low-density pumice to dense clasts of lava or dome fragments. The deposits left by pumice flows are known as ignimbrite and that from the denser non-pumice type is called block and ash deposits, a term which describes their textural characteristics. Correct terminology is important here to avoid confusion between flows (process) and deposits (sediment). Pyroclastic flows of very low density are called surges. A reduction in the volatile content leads to instability in the erupting column. So does an increase of the radius of the vent beyond a certain size. In such cases, pyroclastic material begins to fall from the column (Francis, 1993). This may imply that the eruption column is overloaded with too much pyroclastic material and convection cannot be sustained. Francis (1993) attributed the mobility of a pyroclastic flow to two causes:€(1) transformation of potential to kinetic energy due to fall from a considerable height in the eruption column and (2) fluidisation of the material. A huge quantity of very fine material in the flow is fluidised to act as a lubricant to transfer the large clasts. As fluidisation takes place, large pumice clasts, which are of low density, float towards the top, whereas the fine-grained material concentrates towards the bottom. This segregation structure is seen in deposits of ignimbrite. Ignimbrite can therefore travel over a large distance in a process similar to that of rock avalanches (see Chapter 6). The material descends to the surface of the ground and spreads round as flows of pyroclastic material because of its high speed and large discharge. This is pumiceous ignimbrite which is being deposited. It consists of bits of pumice, dust and gaseous matter.

Ignimbrite After deposition, ignimbrites can undergo welding and be anything from loose sandy ash to a solid rock with the appearance of glass, termed vitraphyre. They can cover a wide range of areas and a large-scale ignimbrite deposit may give rise to a distinctive topography. They often give rise to near flat-topped plateaus dissected by channels. Ignimbrites are gravity-driven flows and therefore tend to concentrate as flat-topped deposits in valleys and lowland. Its distribution, unlike that of ash fall, is guided by the topography. Only in exceptional cases does ignimbrite climb over topographic barriers. Even then the thickness of the deposit varies considerably between the topographic highs and the valleys. Sumatra has striking ignimbrite deposits from past pyroclastic eruptions. Canyon-like river valleys, for example, have developed near Bukittingi in West Sumatra, between it and the neighbouring village of Kota Gadang, which is well known for silver artefacts. Rivers have eroded ignimbrite to create valleys with flat floors and cliff-like walls. The river itself is

261

Lava and pyroclastic deposits

entrenched within the flat valley floor. Landforms in ignimbrite are also prominent in north Sumatra, where there have been several huge explosions from the Toba Volcano. The last eruption about 75 000 years ago released 2000 km3 of ash which reached India. This created Lake Toba, which is the subsiding caldera of the old volcano with the subsequent rise of the island of Samosir in its middle. This is an example of a resurgent caldera. Ignimbrites from the catastrophic 1815 eruption of Tambora formed vertical coastal cliffs on Sumbawa, Indonesia.

Pyroclastic surge deposits Pyroclastic surges have been described by Francis (1993). They consist of low-density dilute material. As a result, pyroclastic surges travel with high velocity, are turbulent and less constrained by topography. Fast surges that emerge radially from the base of the erupting column are known as base or ground surges and are probably the best-studied. In general, they form a wedge-shaped body that thins from the volcano with cross-bedded dune structures in sand that change with distance into structureless material. This is then replaced by planar deposits of fine grains at a further distance. Surges are extremely violent and turbulent, and erosive features are common in the deposited sand. The destructive effects of the pyroclastic surges have been studied in detail for the 1965 one from Taal in Luzon, the Philippines. Taal is a volcano south of Manila on an island within a lake. There is another small lake inside its caldera. It has an explosive history, but the region is highly populated with paddy fields and fishing in the large lake. The eruption took place on 28 September 1965 when for 48 hours powerful surges swept out from the base of the erupting column. It killed about 200 people, damaged trees on the mainland and left thick mud around the trunks.

Block and ash flow deposits Denser block and ash flows are deposited by pyroclastic flows from a hot avalanche, often described as nuées ardentes. Pyroclastic flows on the Merapi in 2006 are described later in the chapter. Block and ash deposits, as the term describes, include magmatic clasts in an ash matrix. They flow under gravity and are therefore usually confined in valleys radiating out from volcanoes. The material originates from viscous andesitic and dacitic lavas of convergent-plate volcanoes. The clasts can be as big as boulders, metres in diameter. Such boulders are often seen on top of the block and ash deposits. The front of a pyroclastic flow is turbulent and highly fluidised. This leads to finer particles being elutriated up into the ash cloud. The cloud then descends down the valley with a proclastic flow emerging from its base. This gives rise to thin and fine-grained ash-fall deposits which are not preserved as well as the denser flows. Pyroclastic deposits therefore cover the slopes of a subduction-related volcano, with thicker accumulation in the radiating valleys. This material with the addition of water gives rise to volcanic debris flows called lahars, a term from Bahasa Indonesia. The material in a dry state may also avalanche down the slopes of a volcano.

262

Volcanic landforms

14.4╇ Volcaniclastic flows:€debris avalanches and flows A continuum exists between dry debris avalanches and wet mudflows called lahars. Debris avalanches can occur in a range of sizes; almost an entire side of a volcano failed at Mount St. Helens in the May 1980 eruption. Francis (1993) listed three types of collapse event that lead to debris avalanches:€magmatic eruptions that blow out the side of a volcano, nonmagmatic explosions due to phreatic steam escapes, and earthquakes and any other event which may cause the volcanic slope to fail in a cold state. The areal size of the large debris avalanche deposits range from 1 to 20 km3 (Francis, 1993). Past failures are indicated by distinctive crescent-shaped scars on the mountainside, although later eruptions of lava may repair the damage. The debris avalanche deposits, however, give rise to a characteristic hummocky terrain at the base of the volcano. The oft-quoted example of such terrain is the 250 km2 on the lower slopes of the Galunggung Volcano in west Java, popularly known as the ten thousand hills of Tasikmalaya. The 1982 eruption of Galunggung is notorious for blocking the engines of aircrafts that flew through the eruption column. About 500 km2 of the surface of the Atacama Desert is covered by the deposit of an avalanche that occurred on the Socompa Volcano, north Chile. Kilometre-size blocks, known as torevas, were detached and slid downslope. In general the blocks slid undisturbed. Part of the mass has been estimated to travel at velocities of 100–200 kmhr−1 and reach a distance of 35 km, but the original stratigraphic features were maintained within the deposit (Francis, 1993). In general, the finest grains in a debris avalanche occur at its base with boulders towards the top. Most of the material tends to travel as a plug without much disturbance. Their long travel tracks have been explained by fluidisation and the conversion of potential energy to kinetic energy, as discussed earlier. The presence of any hydrothermally spread surges, ignimbrites or water acts as a lubricant. Lahars are volcanic mudflows that can be described as slurries that carry boulder-sized material and mud (Fig. 6.6). They originate from water coming into contact with tephra that has accumulated on volcanic slopes and especially with the material in the channels cut into the slopes of a volcano. Hence, material from pyroclastic flows deposited on upper slopes may build lahars on lower ones. The arrival of water could be from rainstorms, the spilling of water from crater lakes, snowmelt on the high volcanoes and pyroclastic flows reaching a stream. Lahars are common on tropical volcanoes during the wet season. The heavy monsoon rainfall over the volcanic slopes of Indonesia and the Philippines and many other places in the humid tropics periodically give rise to destructive lahars which may travel at a very fast pace. The speed of lahars has been listed up to 90 kmph, but is usually much less at 4–8€ms−1. A tropical cyclone in 1991 that caused heavy rain to fall on the newly erupted ashes on the slopes of Pinatubo gave rise to catastrophic lahars over a wide area that destroyed life and property. Unlike Pinatubo, several volcanoes remain active in Southeast Asia and on these volcanoes, such as Merapi, lahars are a perpetual hazard. Mount Semuru, at 3676 m the highest volcano in east Java, is a persistently active volcano (Thouret et al., 2007). Eruptive activities are common with short-lived eruption columns appearing several times a day, frequent explosive eruptions and an unstable summit

263



Fig. 14.5

Volcaniclastic flows: debris avalanches and flows

Paths of destructive lahars originating from Navado del Ruiz eruption of 1985. After Francis, 1993 and Voight, 1996

cone. Small lahars travelling down channels are regular events, due to the daily supply of pyroclastic debris over the summit cone and the high runoff during the rainy season. Large pyroclastic flows tend to occur once in every five years. At least five large lahars, each exceeding 5 × 106 m3, have occurred since 1884. This is a permanent threat to a number of settlements at the foot of the volcano, including Lumajang, a town of 85 000 people. The threat of lahars is especially high towards the southern and southeastern valleys, which annually experience tens of rain-triggered lahars for a distance of 20 km from Mount Semuru. The annual rainfall is 3700 mm, and 500 mm of rain has fallen in 48-hour storms. The rain obviously falls on a plentiful supply of loose volcanic material, including fine-grained ash which leads to hyperconcentrated flows. Thouret et al (2007) determined the annual sediment yield in the local Curah Lengkong river valley to be 2.7 × 105 m3km-2 with an annual denudation rate of 4 × 105t km−2. This is a very high rate and so is any rate approaching it. We may conclude that the slopes of pyroclastic volcanoes that are subjected to high rainfall in the tropics contribute a huge amount of material to the lowlands and ultimately to the coast (Fig. 13.1). Volcanic mudflows are also created, very destructive ones at times, when water and mud escape from crater lakes due to volcanic eruptions. Perhaps the best-studied example comes from the Kelud Volcano on Java which has a deep crater lake. Lahars originate from the lake periodically, unless the level of the lake is lowered. A system of tunnels was engineered to do so, but later volcanic eruptions damaged the arrangement so the lower slopes of the Kelud remain vulnerable to a mixed flow of mud and hot, acidic, sulphurous water. Eruptions on high volcanoes with a snow cover on their upper slopes also give rise to lahars. A catastrophic one occurred in November 1985 following a small eruption on Nevado del Ruiz in Colombia (Fig. 14.5). The incident illustrates not only the origin of this type of volcanic mudflow but also the need to have an effective hazard prevention system which would have prevented the loss of life and property in the valleys. As described below, the disaster killed 23 000 people and destroyed settlements, including the town of

264

Volcanic landforms

Armero, communication links, and crops and livestock. In terms of people killed from a volcanic hazard it ranks 4th (Voight, 1996) surpassed only by Tambora (1815), Krakatau (1883) and Mount Pelée (1902). A small eruption on the 5200 m high Nevado del Ruiz generated a number of pyroclastic flows and surges that scoured and melted ice and snow on the volcano. The meltwater and pyroclastic debris descended the slopes and accumulated in narrow channels to form lahars that reached the settlements within two hours of the eruption. West of the volcano, lahars originated in the headwaters of the rivers Molinos and Nereidas and swept down destructively along the River Chinchiná (Fig. 14.5). Lahars coming down both headwaters reached the River Chinchiná with a flow rate of 13 000 m3s−1 and travelled for 70 km more to reach the River Cauca. Towards the east, lahars came down the River Lagunillas to destroy Armero and the River Guali to pass Mariquita. The eastern lahars were bigger, about 20 000 m3s−1 on the River Guali and an average of 25 000–30 000 along the River Azufrado, which was the main supplier of lahars to the River Lagunillas. The peak value attained was 47 000 m3s−1, one-fifth of Amazon’s discharge (Francis, 1993; Voight, 1996). The peak flow velocity in most of the lahar-scoured channels ranged between 5 and 15 ms−1. Flow depths, except in the River Nereidas, exceeded 10 m. The depth and velocity of the lahars produced a high boundary shear stress and eroded the channel walls to transfer sediment and interstitial water. This in turn increased their flow volumes and velocities. The event showed that a small eruption can produce lahars of huge dimensions and destructive capability. It also illustrated the failure of disaster management in spite of the expectedness of the event. As Voight described it, the ‘disaster happened because of cumulative human error€– by misjudgment, indecision and bureaucratic shortsightedness’ (Voight, 1996:€764). The inhabitants of Armero could have been saved if an alarm had been raised in time. Lahars are thus extremely dangerous, and effectively modify the lower slopes of a volcano, especially the channels. They can travel long distances; lengths of over 100 km have been reported. In sum, material from inside the Earth’s crust that erupts through craters, shapes the volcano and the lower grounds at its foot and, in suitable locations, a high sediment load reaches the sea via the regional streams. The chain of eruption, ash falls, pyroclastic flows and lahars can be seen clearly in high-resolution satellite images. Box 14.1 is an example on the Merapi. Volcanoes have smooth slopes on ashes and fragmentary material. Verstappen (2005) mentioned that although slopes on the stratovolcanoes of Southeast Asia appear to be concave, overall they can be divided into three sections. The top part is straight and essentially formed by the transfer downslope of ashes and coarser material under gravity. These slopes may reach 34°. The middle section is straight with about 8–12° of slope. Most of the volcanic slope may belong to this section which is modelled by deposits from lahars. Valleys are filled and overflows may happen, but this phase of deposition is followed by rapid erosion after an eruption. The lower slope of a volcano is modified by sediment deposited in small radiating ravines in rainstorms. The approximately 2° slope of this section merges with an alluvial plain below.

265



Box 14.1

Volcaniclastic flows: debris avalanches and flows

The 2006 pyroclastic deposits of the Merapi volcano Merapi is a particularly active volcano of southcentral Java from which pyroclastic flows and lahars threaten a populated region that includes the city of Yogyakarta, located close to the southern flank of the Merapi. A number of eruptions and pyroclastic flows continued on the Merapi from May 2006. Four pyroclastic flows due to dome collapse, occurred on 14 June, the largest among them reached 7.78 km. Two of these block and ash flows were especially destructive, burying the village of Kaliadem under 7 m of deposit and killing two people. On 16 June, two IKONOS satellite images acquired by the Centre for Remote Imaging, Sensing and Processing (CRISP), National University of Singapore captured an active block and ash flow. This was first satellite imagery of a transient pyroclastic flow (Thouret et al., 2010).The images also provided a synoptic view of the fresh block and ash deposits of 14 June. It was possible to reconstruct the events from the summit of the volcano to its lower slopes. Figure 14.6, dated 16 June 2006, shows a billowing ash cloud hugging the ground at the front of the flow and rising further towards its back. It was supported by elutriated fine ash and hot gases from the block and ash flow which is seen emerging from below the cloud. One of the IKONOS images demonstrates the passage of the block and ash flow and surge deposits on the lower slopes (Fig. 14.7). Several flow lobes and tongues were seen confined within the box-shaped gorge of the Gendol River. These ash cloud surge deposits and knocked-down trees appear on both slopes where the gorge narrows. A fan-shaped overbank deposit crossed an interfluve to invade a minor valley. Massive overbank deposits reached the right bank of the Gendol River to destroy Kaliadem. Part of the flow avulsed into a tributary valley. The image showed evidences of avulsion into neighbouring valleys due to obstacles formed by previous flow deposits and channel banks plus anthropogenic barriers constructed in the channel to control the

Fig. 14.6 DEM of Merapi draped by the IKONOS image of 16 June 2006, showing the southern slope of the volcano and the Gendol River Basin. A pyroclastic flow is emerging from below the billowing ash cloud. The river channel is marked by deposition of earlier pyroclastic flows. From Thouret et al., 2010. By permission of Elsevier. See also colour plate section

266

Volcanic landforms

Fig. 14.7 Lower slopes of the DEM of the Gendol Valley showing details of a volcanic flow. Segment 1: upper section of the valley; Segment 2:€widening of the valley and a fan of volcaniclastic material; Segment 3:€a wider valley entrenched in a volcaniclastic fan. A:€the moving pyroclastic flow from upstream; B:€a ridge built by a previous flow; C:€an area stripped by an ash-cloud surge; D:€spilling and re-routing of flow at a bend; E:€avulsion of re-routed flow to a neighbouring channel; F:€lobes of flow reaching overbank which buried the village of Kaliadem (KA). From Thouret et al., 2010. With permission from Elsevier. See also colour plate section

passage of flows. The volume of deposited material in the channel suggests the possibility of another lahar after the next rainfall event.

14.5╇ Landscape on flood basalts At certain periods of the Earth’s history, continental flood basalts (CFBs) emerged through innumerable fissures fed by dyke swarms to cover extensive areas. The basaltic lava flowed out in sheets and piled up to form highlands. Weathering and deep dissection give such areas a characteristic stepped topography over time (Fig. 14.8). The upper part of each lava flow is scoracious and softer than the lower parts. The lower part of a flow therefore gives rise to a vertical step, whereas the upper flow is eroded to create a flatter tread-like surface. Cumulatively, this creates a staircase-like appearance, described as a ‘trap’ in old Swedish (Francis, 1993). The term is in common use. For example, the landscape of CFBs in Deccan, India is known as the Deccan traps. The Deccan traps are 65 million years old and they are estimated to have had an original extent of 1.5 × 106 km2, of which 0.5 × 106 km2 survives. The average thickness of the Deccan traps is more than 1 km. Other areas in the tropics where large stepped plateau-like features on CFBs occur are Paraná, southeastern Brazil; an area centred on the Walvis Bay in Southwestern Africa; the Ethiopian and East African Highlands; and the Drakensberg of South Africa (Fig. 14.1).

Landscape on flood basalts

267



Fig. 14.8

Plateau on flood basalt, Deccan, India



Fig. 14.9

Schematic representation of trap rivers, showing an inset river channel within a wide open valley and side slopes in flood basalt. From Deodhar and Kale, 1999. By permission of Wiley Figure 14.9 is a schematic representation of trap rivers from India. Most of the Deccan trap is in a low-rainfall area. The rain shadow of the escarpment marks the edge of the CFB plateau, which rises steeply from the narrow coastal plain of the Arabian Sea. At the top of the escarpment, known as the Western Ghats, the rainfall is from the southwestern monsoon and exceeds 5000 mm annually. Towards the east in the rain shadow of the Western Ghats, the figure drops to 500 mm within a distance of 200 km. As expected, the seasonal pattern persists. The trap rivers tend to have low-gradient wide valleys. Within such valleys, the major rivers have channels incised into bedrock or alluvium. The tributaries, in contrast, have steep upper courses with a change of gradient as they reach the valley flat of the trunk streams. The channels are stable and, apart from seasonal high discharges, they may, as in the Deccan, carry high flows of low-frequency floods. The landforms of stepped plateaus in the CFB areas, as displayed in Figure 14.9, are associated with deep weathering

268

Volcanic landforms

and soil formation. Wide trunk valleys with the river channels meandering in and out of bedrock canyons and incised upper courses of tributary streams are common. The channels carry a fine suspended load of silt and clay and a coarse bed load, mainly of pebbles and cobbles, derived from the jointed bedrock. The coarse material is generally contributed by the tributaries draining the high side slopes of the valley and also by the trunk streams flowing over jointed basalt in locally steep sections. The bar material is therefore coarse and bed armouring is common. Floodplain building does not occur in many rivers and an accumulation of old sediment may form terraces along with gravel fans at tributary mouths, both of which are eroded in low-frequency, high-magnitude floods. The main channels are box-shaped (see Chapter 7). Gorges in bedrock and deep channels in alluvium deposited over basalt are typical, but are not seen everywhere. In places, rivers may be very wide with multiple channels and rocky islands where it flows over the exposed bedrock. Such features have been described for the Mekong River where it crosses an exposure of Mesozoic basalt at 4000 Islands near the Lao PDR–Cambodia border. The maximum river width measured for the Mekong across islands and sub-channels is 15 km (Gupta and Liew, 2007). Deodhar and Kale (1999) provide a list of erosional features such as potholes, grooves, inner channels, scablands and boulder berms. There is almost no sediment on the channel floor at bedrock reaches because of secondary eddies, flow separation and the formation of vortices during large floods (Kale and Hire, 2004).

14.6╇ Conclusion Volcanic activities do not only form a distinctive landscape and provide a huge and widespread volume of sediment; they also influence the local population. Surrounding areas may be devastated periodically and the deposited pyroclastic ash fall and flow material is eroded by rills and gulleys across the slopes. Volcanism, however, has its beneficial aspect. It provides fertile black soils with a high water-holding capacity which are extensively cultivated. This may lead to areas close to volcanoes being densely populated, thus exposing a large number of people to volcanic hazards from time to time. Farming and high-density rural settlements are also encouraged by the huge amount of groundwater that is stored in many places in the large stratovolcanoes. The groundwater appears in a series of springs around the edifice where the slope changes. Faults and old lava flows may complicate the picture (Verstappen, 2005). Volcanoes may dominate a single island such as Montserrat or Lombok, where coralline limestone surrounds the volcano to provide the island with its final shape.

Questions 1. Where in the tropics does volcanism primarily determine the landscape? 2. Explain how the dimensions and slopes of volcanoes depend on the composition of the construction material. What else affects their appearance?

269

Questions

3. Compare the modifications brought to an earlier landscape by ash fall and pyroclastic flow deposits. 4. Why are lahars important as a regional geomorphic process? 5. How hazardous are lahars? 6. Describe the passage of volcanic material down the slope of the Merapi in June 2006, using Figures 14.6 and 14.7. 7. Describe the landscape formed on continental flood basalts. Where in the tropics do such landscapes occur?

15

Tropical karst

Where Alph, the sacred river, ran Through caverns measureless to man Samuel Taylor Coleridge

15.1╇ Introduction to karst The word karst denotes a set of special landforms that develop mostly on carbonate rocks due to their high solubility. The rocks, however, do not dissolve uniformly, but mainly along lines of structural weakness such as joints or bedding planes. This happens both at the surface and inside the body of the rock. Water enters the subsurface through openings on the exposed face of the rock and flows through subterranean passages and caves, probably emerging some distance away in springs. As the drainage progressively passes underground, a bare dry stony landscape evolves; this is karst. The word comes from a high barren plateau in western Slovenia where such features are well displayed. The original word kras was later modified to its present germanicised form, karst, which is used to denote similar landscapes anywhere. The publication of Jovan Cvijíc’s book Das Karstphaenomen in 1893 accelerated the study of karst geomorphology. A terminological complexity exists in karst geomorphology. Almost all karstic features are known under several names, each derived from a regional usage. This account carries only the common international terms in order to avoid this difficulty. A complete list is usually found in books specialising in karst geomorphology, such as Ford and Williams (2007). Such volumes should also be consulted for greater details. Complications also arise because of the use of the word ‘karst’ as a suffix to describe landforms that are not really karst but strongly resemble it. Pseudokarsts are landforms which resemble karsts but are not produced by solution or other usual processes by which a karst landscape is formed. Thermokarsts are areas with topographic depressions caused by the thawing of ground ice. Palaeokarsts are old karsts that are now covered with younger sediment and are no longer affected by solution activities. Some geomorphologists do not restrict the word to limestone features but extend it to solution forms such as furrows on granite or basalt (Ford and Williams, 1989). This chapter covers only the true karst on carbonate rocks:€ its forms and operating processes. Karst topography is best developed on limestone which consists of calcite (CaCO3). Dolomitic limestone, Ca,Mg(CO3)2, and limestone metamorphosed to marble may also display karst features but less distinctly. Such rocks are not easily soluble. 270

271

karst in the tropics: the geographical distribution

Hollows, often with open shafts at their lowest point, pockmark the surface of a karst area and lead to subterranean passages and subsurface caves in limestone. These features are essentially formed by solution or structural collapse aided by solution. In certain cases, small steep hills rise above the hollowed surface. Special features such as erosional notches develop where the sea reaches the limestone. A sufficient amount of water is required to dissolve part of the rock. Karsts are not associated with an arid climate. It has been suggested that certain karstic features are more common in the humid tropics or a modified version of the general form occurs in the humid tropics. Such forms have been well described but their connection with the humid tropics has not been clearly explained. It is generally suggested that the process of karstification operates at a high rate in the hot humid climate where organic acids also are easily available at ground level (Jennings, 1985). In this chapter, the processes associated with the formation of a karst landscape are discussed, with special reference to the features common in the humid tropics.

15.2╇ Karst in the tropics:€the geographical distribution According to Ford and Williams (1989), surface exposures of carbonate rocks cover about 12 per cent of the dry ice-free land. Their estimation of karst areas, however, is a little less, between 7 and 10 per cent. A number of environmental conditions are essential for the development of a good karst landscape, but these do not always coincide with carbonate areas (Thornbury, 1960). These are: 1. The presence of a soluble rock, preferably limestone, at or very near the surface. 2. Such a rock should be dense, highly jointed and preferably thinly bedded, encouraging the circulation of water along marked paths. 3. The area should have at least a modest amount of rainfall. Sweeting (1972) suggested a minimum of 250–300 mm annually for limestone karst development. Karstic features found in arid areas are therefore relict and indicate a wetter climate at the time of their formation. 4. The presence of vegetation which increases the acidity of the circulating water by adding biogenic CO2 and organic acids. 5. The presence of entrenched trunk valleys below the limestone with karst, allowing groundwater to descend through the body of limestone and emerge as surface streams at a lower elevation. A continuous passage of water through the rock is extremely helpful. The requirements for the good development of a karstic landscape therefore explain why (1) karst is not everywhere associated with surface exposures of carbonates and (2) karst is expected to be well developed in the humid tropics. Jennings (1985) has described the tropical variety as the botanical hothouse karst. In the tropics, good karst landforms occur in south China and countries of Southeast Asia such as north Viet Nam, Lao PDR, Thailand, east Myanmar, north Sumatra, north and central Java, central and south Sulawesi, islands of eastern Indonesia, Papua, New Guinea,

272

Tropical karst

Sabah and Sarawak. Karst landscapes have also been reported in India and Southern Africa and Madagascar. They are well developed in Central and South America and the Caribbean islands, including northern Yucatan, Jamaica, northern Puerto Rico, western Cuba and Barbados. Karst probably occurs in more areas in the tropics, but some locations have not always been well reported. Karstic features are also seen in the evaporites (gypsum and salt) in the arid tropics.

15.3╇ Karst hydrology Water in karst can be divided into three types, based on location. 1. Atmospheric water from direct precipitation over a karstic area enters the subsurface. This is known as autochthonous or autogenic water. Water also arrives from outside the karstic area by either surface flow or runoff from rivers, subsequently passing into the subsurface only through specific openings in the rock surface. Such flows from external noncarbonate rocks to karstic areas are allochthonous or allogenic water. Generally both types of water are present in the bedrock (Fig. 15.1). 2. Water that enters karst accumulates in the vadose zone, the zone of aeration above the groundwater table. This water passes through the body of the rock either by percolation or, more effectively, along openings provided by joints, fractures, etc. In the vadose zone above the water table, usually only a fraction of the holes are filled with water; the rest are filled with air. The proportion of holes filled with water depends indirectly on the length of time since the arrival of water originating from precipitation or runoff. The entry of water through the pores can be complicated. The passage of water may be impeded by air bubbles or locally modified by perched aquifers above a localised impermeable layer of shale or chert. 3. Water that has reached the groundwater table and augmented it. It is controversial whether the pores in the rock below the groundwater table are all filled with water or water is stored mostly along various passages and openings and thereby occurs at different depths in the rock. According to the second view, water is stored in an interconnected network of passages. These passages do not occur at the same level, and no typical water table as such exists. Presumably, the pattern of water storage in the subsurface depends on the physical property of the local rock. Pores in some rocks may be too small and not interconnected enough to store water. Williams (1983) recognised a special zone in the weathered upper part of the bedrock below the soil:€the subcutaneous or epikarstic zone. This zone is periodically saturated, especially after storms when the water table rises. Cavities and passages are efficiently enlarged at this level but not at the level immediately underneath. A perched water table therefore occurs episodically. This then slopes towards larger passages that originated from joints, fractures, etc. This lateral component of flow results in enlarged openings at the surface through which rain or surface water subsequently enters the bedrock. The nature of the passage depends on the size of the openings. If the openings are small then the flow is assumed to be laminar. Flow velocity rises in large openings, leading to

Dissolution of karst rocks

273



Fig. 15.1

Both types of karstic water in bedrock fluctuating eddies and turbulent mixing. Flow through large openings is therefore more capable of both reacting with and mechanically eroding bedrock. Ford and Williams (1989) have discussed the hydrology of a karst landscape in detail.

15.4╇ Dissolution of karst rocks The general reaction between a limestone and water can be written as

H2O ⇔ H++ OH− CaCO3 ⇔ Ca2++ CO32− 2+ Ca + CO32− + H+ ⇔ Ca2+HCO3−

(15.1) (15.2) (15.3)

The equations are unbalanced and the reaction moves forward at rates proportional to the concentration of reactants (Ford and Williams, 1989). The reaction is accelerated by the presence of CO2 in the soil, which may occur in the pore spaces that are not filled with water. Green plants and soil fauna and flora respire quantities of CO2. The microfauna and flora in soil are densest in and near the rooting zone, from which CO2 diffuses out to the soil base and ground surface. The production of CO2 increases with temperature and it is likely that this is a factor behind well-developed karst in the tropics. The amount of CO2 in the soil, however, varies tremendously. Ford and Williams (1989) listed a number of factors that influence the reaction between the rock and the passing solution. These factors are various properties of the soil (types, texture, horizon, depth, drainage and exposure), vegetation cover, soil flora and fauna, and seasonality in warming and wetting.

274



Fig. 15.2

Tropical karst

The result of mixing of two saturated solutions (A and B) in a CO2 or H2S system. It produces an undersaturated solution (C) resulting in subsequent dissolution along CD. Shown for both calcite and dolomite. From Palmer, 1991 Soil CO2 is a heavy gas and drains into fissures and caves below. The gas is further increased in volume by the decomposition of vegetal matter and animal droppings that wash into caves. Subsurface drainage, especially streams in flood flowing through underground caves, picks up the increased amount of CO2 and becomes more corrosion-prone. Mechanical mixing of karst waters from two different sources increases the aggressiveness even if both types of water are saturated. The mixing leads to water being unsaturated and more carbonate is dissolved from the bedrock, especially if the concentration of CaCO3 in any of the waters is below 250 mgl−1 (Fig. 15.2). The solution of limestone thus depends on various factors. It has been suggested that among these factors precipitation acts as the principal control. Ford and Williams (1989) referred to a study by Maire in very wet Papua New Guinea (annual rainfall 5700–12 000€mm) where the calculated annual solution denudation rate is very high, 270–760 mm in 1000 years. Karst is also denuded by mechanical erosion as turbulent water carrying sediment rushes through subsurface openings and passages. The amount of available water, its CaCO3 content, the percentage of limestone in the basin and the time that is available all control the solution denudation rate.

15.5╇ Karst landforms The development of karst landforms in soluble rocks depends on the ambient hydrological and geochemical conditions. Such landforms are developed best in a humid climate where enough water is available to circulate through the bedrock. Temperature is also an important factor, as it influences the water budget, rate of chemical reactions and biochemical processes. The lithological and structural properties of the bedrock determine the solute pathways, the rock strength of the support roofs of underground passages and the susceptibility of the ambient bedrock to solution and mechanical erosion (Ford and Williams,

275

Karst landforms

1989). Although similar landforms are found in karst worldwide, landforms in the countries of the humid tropics tend to display certain characteristics. We first introduce the general case of karst landforms and, in the next section, review the special case of tropical karst. The dimensions of karst features range from millimetre-scale dissolution pits and groves on limestone surfaces (also called limestone pavements) to hills several hundred metres high and underground caves that stretch for kilometres.

15.5.1╇ Limestone pavements Small-scale solution features, millimetres to centimetres in scale, with a circular or linear geometry, evolve on flat or nearly horizontal exposed limestone. It is thought that such small features are also associated with algal or bacterial activities, and certainly bacteria, fungi, lichens and mosses, once established in the indentations, can grow. Larger circular features are known as solution pits and linear features as karrens. Karrens elongate along joints and veins in rock, or as rills along lines of dissolution. Such channels may widen and deepen downslope. These surface indentations vary in appearance and size, and large areas with such features (karrenfeld) are found either on bare rock or under soil and vegetation. Large fissures of this type are known as lapiés and, if exposed, they are difficult to cross, as they may measure up to metres in scale. Lapiés and other indentations, however, may occur under a cover of soil and weathered material, which is coloured red from the release of iron impurities from the carbonate bedrock as it dissolves. The red regolith is known as terra rossa. The term pinnacle karst is used to denote an area where these depressions are separated by large carbonate spikes that are sharpened towards the top. The term ‘stone forest’ is aptly used to describe such karst in China. Pinnacles more than 45 m in height have been reported in Papua New Guinea (Williams, 1971) and Sarawak (Osmaston and Sweeting, 1982). Karren forms may also develop on inclined surfaces that are bare.

15.5.2╇ Dolines or sinkholes Dolines, small openings through which surface drainage passes underground, are probably the most common feature of the karst landscape. The term, in common use, is derived from Slovenian. These depressions are known as sinkholes in North America; this is another term widely in use. The openings are circular or near-circular in shape, with surface diameters ranging from several metres to a kilometre and depths from several to hundreds of metres. Their side slopes vary between gently sloping to near vertical, and they can resemble a saucer, a funnel with a shaft at its bottom, or even a vertical hole. Dolines may occur as single individuals or as a densely packed group giving the landscape an egg-box appearance. Their distribution pattern may also be skewed by lines of structural weakness. Dolines are classified according to their origin. The common types are solution doline, collapse doline, subsidence doline and suffusion doline. Other less common types have been listed (Ford and Williams, 2007). In general, a single process dominates, although more than one is usually involved in the formation of dolines (Fig. 15.3). Solution dolines are formed by the mass transfer of carbonate material in solution from the surface to the subsurface, creating a surficial depression. This requires the establishment

Tropical karst

276



Fig. 15.3

Types of doline of a drainage system below the doline to connect the surface recharge to an underground passage. Usually while limestone is either exposed or under a thin cover of alluvial material at the centre of the doline, its flanks could still be under an overlying impermeable rock. Once the doline is formed, a positive feedback comes into operation because of the centripetal concentration of flowing water and associated corrosion inside the depression. The biogenic production of CO2 in soils overlying dolines helps. The connection to the subsurface passage usually improves with time, thereby accelerating drawdown and the passage of water underground, speeding up the formation of dolines. Collapse dolines have smaller openings and steeper sides, although both characteristics may change over time. They are formed by the collapse of the surface layer exposing a shaft that connects with underground passages. The collapse may happen in three ways: 1.╇ final collapse of part of a cavern roof earlier weakened by solution 2.╇ undermining of the roof from below as the cavern expands 3.╇collapse of a rock layer due to removal of buoyant support by the lowering of the water table. This makes part of the rock unstable. In all cases, the collapse of a rock layer provides a connection pathway for the surface drainage towards the flow in the underground passage. In some cases, the shaft is quite deep and, after rainfall and recharge, a temporary underground waterfall to the cavern may operate. Subsidence dolines are formed when the ground surface sags down forming a coneshaped depression without a significant faulting of the rocks. No rupture in the rock happens as in the case of collapse dolines, but the sagged depression concentrates surface water and forms and behaves like a doline. Water in suffusion dolines seeps through thick unconsolidated regolith or detritus overlying a karstic rock. This creates a depression by the removal of fines in solution or downwashing, a process known as suffusion. Water then percolates through the sagged surficial material to develop a karren vertically, and finds a connection with the cave drainage. Other terms have been used by researchers to identify dolines of different origin, but these four are the ones normally used. Where dolines are closely spaced and all the space is taken up by depressions with shared divides, the landscape is known as polygonal karst. Thousands of dolines may occur inside one square kilometre, although their frequency is often in single figures within this area. It has been suggested that dolines vary in morphology in different climatic zones, but the concept is still in a hypothetical stage. White

277

Karst landforms

et€ al. (1984) showed that the average depth of depressions is high in the tropical karst regions of Puerto Rico and the Dominican Republic when compared to temperate karst in various regions of the United States (White et al., 1984). It could be due to higher rainfall in the humid tropics and a positive water balance. The density of dolines, however, varies widely.

15.5.3╇ Uvalas and poljes Dolines are closed depressions of limited size, but closed depressions bigger than dolines may also occur as part of a karst landscape. Uvalas are formed by merging dolines and, as expected, are characterised by an uneven floor with multiple topographic lows. They can be of irregular shape, and with sizes ranging between 5 and 1000 m in diameter and 1 and 200 m in depth. Poljes are bigger depressions with steep sides and usually a flat floor, several kilometres in length. A number of them are likely to be structurally controlled, and some are elongated along lines of structural weakness. They are surrounded by impermeable rocks that overlie the soluble carbonates, and poljes have been described as the inliers of a fluvial landscape (Ford and Williams, 1989). External streams flow across poljes, frequently disappearing at a doline on the polje floor. The dolines are usually not efficient enough to cope with flood conditions and a shallow short-term lake may episodically cover the floor of the polje. This enables the depression to expand by lateral corrosion. The floor of poljes often acts as a window on the local water table, which fluctuates leaving the polje with either a dry flat floor or under an aggressive sheet of water.

15.5.4╇ Karst valleys River valleys that cross the karst have distinctive features. An allochthonous river with a large discharge may succeed in crossing a karst landscape without losing all its water underground. Such rivers usually flow in gorges that are created due to the combination of the solution of carbonates by the chemically aggressive river discharge flowing across a carbonate floor and abrasion by the clastic sediment it carries. Perennial allogenic rivers function as local base levels for the section of the karst they flow through. Karst gorges also originate from the collapse of roofs of underground caves, making part of the subterranean river visible. Smaller rivers may not succeed in crossing the karst. Such rivers tend to lose water through the dolines along its valley and ultimately no water is left, so the valley below the last active doline becomes dry. This lower valley only carries water in large floods which exceed the capacity of the upstream dolines to capture water. Such valleys are known as semiblind valleys. Over time, a difference in elevation builds up in the form of a rise that separates the upper valley, which carries water and is thus capable of eroding and lowering its level, from the lower valley, which, in the absence of erosion, remains near the original elevation. The lower part of the valley may remain dry even when floodwaters arrive. It is then a blind valley.

278

Tropical karst

Pocket valleys are the opposite of blind valleys. Pocket valleys occur where a stream descends through the soluble rock and emerges at its contact with an impermeable layer underneath. This results in a resurging spring and a U-shaped valley with steep walls and a receding headwall at the spring head. The size of the valley is determined by the volume of water discharged. Pocket valleys are also known as steepheads.

15.5.5╇ Cone and tower karsts The sharply edged standard dolines, the commonest feature of the karst landscape of the temperate latitudes, are often replaced in the humid tropics by star-shaped depressions surrounded by cone-shaped hills. This feature is known as cockpit karst, a name derived from Jamaica where such depressions were used in the past for holding battles between fighting cockerels with an uninterrupted view for the spectators. Like dolines, they usually occur in concentrations. Cockpit karst has been reported from outside the tropics, but it is more frequent in tropical countries such as Jamaica, Indonesia (especially Java), the Philippines, Viet Nam and southern China. Several explanations have been forwarded to explain the morphological difference between cockpit karst and standard dolines. A rapid enlargement of dolines to create cockpit karst has been attributed to the increased acidity in surface and subsurface water from rapidly decomposing vegetal matter in the humid tropics. It has also been suggested that cockpit karst develops where dolines expand faster than they are deepened. Dolines can only deepen to the water table, which in the humid tropics is usually shallow, but they can continue to enlarge laterally. Several neighbouring cockpits may ultimately overlap and merge. The low hills marking the boundary of a cockpit are converted over time into cone-shaped rises by the expansion of cockpits from various sides. This kind of landscape is known as the conekarst. Cones also display remarkable symmetrical shapes. Kangning (1992) measured the morphometry of about 750 cones in Guizhou, southern China. The cones were generally symmetrical, and their slopes measured between 45 and 47°, regardless of their structure. A tower karst is a steeper variety of karst risers (Fig. 15.4). In general, towers are steeper and higher than cones, and have a bigger height–diameter ratio. They are considered to be common tropical karst forms although, like cones, temperate examples have been reported. These residual hills come in a range of shapes, from towers to almost conical or hemispheric. The common dominant feature is their sharp rise from the surrounding plains, although some may have pedestals. Towers are commonly grouped, although they do occur in isolation. It is better not to expect a typical appearance as tower karst exists across a range of forms, any of which may carry the suggestion of a tower. Wilford and Wall (1965) grouped different types of tower karst in Sarawak, where all the towers are residuals on a planed surface: • residual hills that emerge from a planed carbonate surface covered with alluvium • residual hills emerging from limestone inliers in a surface which is a plane mostly across non-carbonate rocks

279



Fig. 15.4

Karst landforms

Diagrammatic section showing tower karst, caves and surface drainage. After Jennings, 1985 • residual hills emerging through a surface cover of thick alluvium • isolated towers that rise from steeply sloping bases of varying lithology. Karst towers have been recognised as time-transgressive forms, developing slowly, and being older near their tops and younger at their bases (Ford and Williams, 1989). They have been studied in Puerto Rico, Cuba, Malaysia, Viet Nam, south China and other places inside the humid tropics. Towers as high as several hundreds of metres above the local plains have been reported (Wilford and Wall, 1965). There is a general inclination to attribute their shape to structural control, the rock disintegrating along joints and fracture zones. Tower karsts are known under different regional names:€pepinos, haystacks and mogotes. The term mogote is common in Spanish-speaking countries. Asymmetrical mogotes in Puerto Rico have evoked considerable research interest. Such an asymmetry with steeper windward eastern slopes has been attributed to case-hardening of limestone by repeated cycles of solution and reprecipitation of CaCO3. Water runs down the eastern slopes due to frequent rainfalls in the prevalent trade winds, followed by drying of the same slopes by the same winds between rain events. This leads to precipitation on the slopes of amorphous calcium carbonate, which is harder and difficult to dissolve. Hence the steeper slopes on the windward side (Monroe, 1976). This explanation, however, has been questioned and the asymmetry attributed to other factors, such as basal erosion. Only a fraction of towers are asymmetrical in Puerto Rico and they have not been reported to be asymmetrical elsewhere. A number of hypotheses have been offered to explain the origin of towers and why they are found in certain areas instead of cones. These suggestions involve the depth of regional water table, surface fluvial processes, intense corrosion along zones of structural weakness, river erosion, etc. It is possible that towers may owe their origin to a variety of processes (Williams, 1987).

15.5.6╇ Caves Morphology in karst is studied both on the surface and underground. Karst rocks, especially limestone, carry a number of underground openings, called caves, through which water in the subsurface circulates. The standard definition of a cave is a solutional opening underground, large enough for a human being to enter. According to Ford and Williams

280



Fig. 15.5

Tropical karst

Diagrammatic sketch of a cave in limestone. The upper part (left side) carries vadose water via dolines and a higher entrance. This part stays dry periodically. The lower part (right side) carries phreatic water, mainly from groundwater rising through a fissure. The water from the cave emerges at a spring to feed a surface stream (1989), a karst cave is a solutional opening, bigger than 5–15 mm in diameter or width. This is effectively the smallest aperture that allows turbulent flow to form. Larger caves are also called caverns. They are common in wet climates, and allogenic bigger rivers tend to flow through larger caves. However, large caves do occur in arid regions, possibly as relicts from an earlier wetter climate. Caves may occur as isolated openings in rock but they usually have input and output points connected by conduits, at least 5–15 mm in diameter. The entire set-up behaves as an integrated system. Caves occur at different depths through the bedrock. A long integrated cave system may continue at various depths (Fig. 15.5). The upper part of a cave may be in the vadose zone and dry, but its lower part may be in the phreatic zone and filled with water. As the water table inside the rock fluctuates, the proportion between dry and wet parts of a cave changes. The form of the passage also tends to be different in the two zones. Vadose caves tend to have passages resembling canyons, with steep walls and vertical shafts. Phreatic caves are tube-shaped and more equidimensional. The form of a cave passage depends primarily on the solution in vadose (free fall) conditions, phreatic (pressure flow) environment, or on conditions that alternate between the two during floods. Caves therefore have entrances via dolines and vertical shafts (input points); passages; larger openings called rooms or chambers; and spring outlets (output points). They tend to form along paths of greatest discharge of groundwater and solutional aggressiveness (Palmer, 1991). Caves may end abruptly with rockslides or collapses, the filling of the passage by silt and clay, and the narrowing of the passage due to a change in rock characteristics which allows water to permeate but does not allow people to pass. Such instances are known as termination. Cave chambers can be huge. The biggest known chamber (Sarawak Chamber, Good Luck Cave) occurs in the Mulu National Park of Sarawak, East Malaysia. It measures 20 million m3. The Belize Chamber in the Tun Kul Cave in Belize is about a million m3. Many large chambers, 100 000–500 000 m3 in volume, have been reported (Ford and Williams, 1989). Palmer (1991) stated that caves usually need a time-span of 104–105 years to reach a traversable size. The pattern of linked underground passages in a cave depends on the mode of groundwater recharge. For example, discharges that pass underground via sinkholes tend to form what is called a branching cave, where subterranean tributaries join to create a higher-order system. Other cave forms, often more complicated

281

Karst in the tropics

and with passages at different levels, are created by factors such as steep gradients, greater undersaturation and lines of geological weakness (Palmer, 1991). Palmer is of the opinion that climate affects caves only indirectly via changing patterns of rate and character of recharge. Tropical caves formed at a higher presence of CO2 may enlarge quickly but their morphology remains comparable to caves formed in other climatic regions. The underground passages and chambers are formed along lines of structural weakness, and the geological structure of the bedrock may influence the shape of the cave. The enlargement of the openings is carried out both by corrosion by circulating water and abrasion by sediment that arrives in high flow from the surface via input points such as dolines with vertical shafts. Small sculptured markings, formed by corrosion and abrasion, also occur on the walls, floors and ceilings of such openings. Caves display forms made with chemically precipitated carbonate deposits; the general term for such forms is speleothem. The deposited material is called travertine. Such precipitation comes from circulating waters because of the mixing of water and air in caves, each with a different CO2 concentration. CO2 is then diffused to the atmosphere of the cave because the partial pressure of the gas is much higher in water. This leads to precipitation of CaCO3. Water in the cave may also evaporate to some extent, causing precipitation of CaCO3. In either case, CaCO3 is precipitated from the ceiling to form elongated near-vertical bodies of lime stretching from the roof, and known as stalactites. Alternatively, lime-rich water may be evaporated from the floor leading to the precipitation of CaCO3, which slowly builds upwards to form incipient pillars called stalagmites. Stalagmites are also formed on the floor by water dripping from the roof. In some cases stalactites and stalagmites coming from opposite directions meet to build pillars. The passages of floodwater entering from input points and flowing through caves are geomorphologically very efficient. Floodwaters tend to enlarge caves and openings both by abrasion from clastic load carried in suspension and along the bed of the underground streams, and by corrosion. Corrosion may occur high up the walls and even at roof levels. The hydrologic effect of caves varies. A number simply operate as an underground draining system. Others may transfer water across topographic divides as underground water flows follow hydraulic gradients, not surface slopes. Such underground transfers may lead to stream piracy or cutoffs across rock meanders.

15.6╇ Karst in the tropics Research on karst landscapes located in the tropics really started in the twentieth century (Lehmann, 1936; Sweeting, 1958). The morphological difference between the tropical forms and the type of landscape common in the standard karst regions, such as the high plateau on the eastern Adriatic coast, northern England and the southern Appalachian Plateau, was noted. Cockpits, cone karsts and tower karsts were accepted as characteristic features of the tropical karst. Such forms have been described in the literature but a complete account of their origin is still elusive. The morphological characteristics of the tropical karst have been attributed in a general way to high temperature and precipitation, dense vegetation

282

Tropical karst

and the high production of organic acids. The karst of Southeast Asia, for example, was described by Jennings (1985) as the ‘botanical hothouse extreme’. Viles (1984) discussed karstic landforms produced largely by direct biological processes. Boring plant filaments from colonies of blue-green algae, red algae and diatoms have been observed as etching agents below the colonies. These activities are found on bare rocks in warmer climates and particularly on seacoasts, as described from the small Caribbean island of Hell (Folk et al., 1973). Gillieson (2005) recognised the karst in Southeast Asia as most diverse, and listed a number of factors€– varied geology, uplift history, eustatic change, past and present climate€– to account for this diversity. In Southeast Asia, karst occurs from the sea level to nearly 4000 m, comprising large plateaus with dolines, cone karsts, tower karsts and lowlying swamps. All karst in the humid tropics probably displays a diversity of similar forms, some of which are less common in the temperate regions. A standardised terminology for large-scale karst morphology of the type that is common in the tropics has been borrowed from karst in southern China. The type occurs near Guilin, Guanxi Province, China, where karst is formed on pure, thick limestone. This limestone has been uplifted and subsequently subjected to a warm and humid monsoon climate over a long period. Three principal karst styles have been recognised, the description of which has been taken from Gillieson (2005). 1. Fenglin (tower karst):€a number of limestone towers that rise above a plain, commonly under alluvium. 2. Fengcong (cone karst):€ a limestone massif with a large number of closed polygonal depressions, sharing a common rock base and usually at a higher elevation than the surrounding area. The shapes of the conical hills separating the depressions are determined by local structure and rock type. 3. Gufeng (isolated towers):€single towers rising out of an alluvial plain formed by extreme dissection of fenglin by meandering rivers. The rivers isolate blocks of limestone and they are subsequently eroded by a combination of undercutting, cave development at the foot of the blocks/hills, and material spalling off cliffs. Such karst areas of high relief occur throughout the humid tropics, modified by local environmental factors. For example, the karst near Ipoh in the Kinta Valley of Malaysia shows types of both the fenglin and fengcong style of karst. On the eastern side of the valley, the rugged towers resemble fenglin; on the westen side they are more like fengcong with gentler relief. Gillieson (2005) has ascribed the difference to greater metamorphism of the rocks on the eastern side, as they are next to a granitic upland. Some tropical examples of the high relief karst of towers and cones are better known than others. The Chocolate Hills of Bohol Island of the Philippines are a good example. This is a landscape of smooth, cone-shaped, isolated residuals of limestone. The cones have been locally metamorphosed and reflect the gentle folding on a regional scale. Several of these present the classic cone karst (fengcong) or modified forms of high-relief karst. The Gunung Sewu area of southern Java is also a good example of both. Here karst and volcanic deposits have been juxtaposed, with thick volcanic ash filling the depressions and cones and cockpits being aligned with faults and major lineaments. Highly concentrated

283

Karst in the tropics

organic acids passing through the bedrock and across the surface of the limestone mark the tropical towers and limestone walls with solution rills and ripples. Swamp slots up to several metres deep, filled with fine-grained alluvium, also occur at or near the base of towers and limestone walls, indicating the location of past or present water levels. These are at the elevations where the alluvial plains met the limestone towers. Another characteristic landform of the humid tropics karst is pinnacles, a spectacular landform covering a small area. The pinnacles are sharp-edged vertical blades of limestone, up to 50 m tall and closely packed. They may also penetrate deep into the subsurface. Thick organic soils, root mats and deep fissures separate individual pinnacles from each other. Their origin has been attributed to structural control and intense rainfall. Pinnacles are best developed in massive limestone with well-developed horizontal and vertical joints (Osmaston, 1980; Waltham, 1995). Several hectares may contain hundreds of pinnacles. A classic location occurs on Gunnung Api, Sarawak, at an elevation of 1100 m, with pinnacles crowding a shallow bowl-shaped depression about 2 ha in area. Certain landforms such as poljes and dolines may occur in the tropics without morphological modification. Poljes are common on tropical limestone plateaus in Malaysia, Viet Nam and New Guinea. They tend to be underlain by alluvium or an impermeable sedimentary rock such as shale or sandstone. Spring-fed streams may traverse such poljes. Dolines may occur with the standard morphology or modified as star-shaped depressions. The star-shaped depressions are enclosed by convex hillslopes, and steep gullies flow to a sink in the centre of the depression. The most spectacular example of a tropical karst arguably comes from the Gunung Mulu National Park in the Malaysian state of Sarawak, on the island of Borneo. This involves a 1500 m thick Tertiary limestone that has given rise to a spectacular line of hills deeply dissected by river gorges into four separate units, as shown in Figure 15.6. The limestone is pure, massively bedded, and dips steeply to the northwest while faults and joints run northeast. The upper surfaces of the hills are very rugged with structure-guided pinnacle karst, deep gullies and cliffs. Locally, limestone hills display the standard forms:€dolines, valleys and caves. Several of these features are impressive, e.g. the 1 km diameter collapsed doline towards the south known as the Garden of Eden. The rivers shown in Figure 15.6 have cut deeply incised gorges overlooked by 300 m cliffs and high-level caves with alluvial fans at the contact between the limestone hills and the gorge floors. The presence of past alluvial fans and the wide vertical range of caves in the hills indicate a history of regional uplift. The caves of Gunung Mulu (Mulu Hills) are famous, with about a hundred known so far and a total mapped length of 295 km (Brook et al., 1982). The caves include large passages that follow strikes, excellent examples of speleothems, and evidence of past floods. The assemblage includes Deer Cave, which with a 120–150 m diameter is the largest natural cave passage in the world; and the fault-guided Sarawak Chamber in Gua Nasib Bagus (Good Luck Cave), the largest chamber reported inside a karst cave. It is 600 m long, 415€m wide and 80 m high. A number of other very large caves and passages occur in the Mulu Hills. Flow through the caves is still active. Seasonal rivers with outside sources rise in floods and continue to aggressively develop the openings to these caves. Karst features near sea level have been repeatedly submerged and exposed due to sealevel changes during the Quaternary. This has given rise to magnificent tropical limestone

284



Fig. 15.6

Tropical karst

Limestone massifs in the Gunung Mulu National Park, Sarawak, Malaysia. From Gillieson, 2005. By permission of Oxford University Press coasts as in the Ha Long Bay of north Viet Nam, where a large number of towers emerge above water as spectacular steep-sided islands separated by narrow channels.

15.7╇ Tropical karst as an environment Karst geomorphology in the humid tropics is fascinating because of the rich development of the standard karst landforms, the prevalence of high-relief karst and the scale of the underground openings. Tropical karst regions are also associated with resource development and ecological niches. As a number of tropical countries with karstic areas are also



285

Tropical karst as an environment

Fig. 15.7

IKONOS image of coast near Phang Nga, southern Thailand. Limestone towers are partially inundated by a Holocene rise in sea level and surrounded by mangroves. Boats provide scale. Seawater frequently invades sinkholes in karst giving rise to a spectacular landscape. IKONOS satellite image © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2007), reproduced with permission. See also colour plate section developing countries, such associations are important and require in-depth studies of tropical karst geomorphology. For example, limestone in the Malay Peninsula, as in the Kinta Valley discussed earlier in the chapter, occurs next to granitic bodies. Hydrothermal mineralisation has given rise to deposits of tin, tungsten and other minor ores, part of which occurs inside the limestone, like the vein deposits of tin-bearing cassiterite in the walls and ceilings of caves of the Kinta Valley. The change in sea level associated with glacial advances and retreats during the Pleistocene affected karst topography near sea level. During the last advance of ice, limestone next to the sea was exposed subaerially, leading to the development of karst topography. The subsequent Holocene rise of sea level inundated part of the topography, creating a coast with cliffs, caves and fascinating forms such as blowholes which developed from limestone shafts. This has resulted in the development of tourism in many areas where the coastal features are essentially partially submerged karst (Fig. 15.7). Good examples are the innumerable islands and narrow passages of Ha Long Bay of north Viet Nam that developed from a rising sea level inundating the lower part of a tower karst landscape. Similar features are also seen in the Malaysian coastal resort of Langkawi and part of the coast of southern Thailand. This connection between a partially submerged karst landscape and tourism is common in many parts of the world.

286

Tropical karst

In the tropics, karst areas are often recognised as maintaining a range of biodiversity not found elsewhere. The range of hydrology, temperature and light all lead to the development of specialised flora and fauna on the exposed surface of the karst and in the dark passages underground. Research in tropical karst still has the lure of the dark unknown. The karst environment in the tropics also provides the backdrop for agriculture (polje floors and limestone pavements), specialised forestry (in high-relief karst areas) and water management. A basic knowledge of the geomorphology of karsts in the tropics is therefore extremely useful.

Questions 1. What are the essential environmental conditions for the development of a good karst landscape? Are such conditions common in the humid tropics? 2. Where would you expect good karst topography in the tropics? 3. What characteristics of karst are traditionally attributed to the tropical climate? 4. Describe cockpits, cones and tower karsts. 5. Examine Figure 15.7. What can you say about the origin of the landscape from this IKONOS satellite image? Why are such areas becoming economically important?

16

Quaternary in the tropics

This is like déjà vu all over again. attributed to Yogi Berra, New York Yankees

16.1╇ Introduction The present landforms, geomorphic processes and sedimentary deposits in the tropics have been influenced by events which took place in the Neogene. Events in the Tertiary included the building of two huge mountain chains, the Himalaya and Andes, which in turn modified the tropical climate. For example, the origin of the Himalayan Ranges and the uplift of the Tibetan Plateau either created the South Asian Monsoon system, or strengthened it immensely from its previous state. This gave rise to a strong seasonality and storminess in regional rainfall that translated into seasonal river discharges and high-magnitude floods in the rivers of South, Southeast and East Asia. Prell and Kutzbach (1992) associated the rising of the eastern Himalaya with an east–west decrease in precipitation along the mountain chain, enhanced aridity in Rajasthan and western Sahara, and cooler conditions over northern Asia and Europe. The very high mean elevation of the Tibetan Plateau (4000–5000€m) prevented the influx of the monsoon from the south, leading to aridity and a low level of glaciation north of the Himalayan chain (Gasse and Derbyshire, 1996; An et al., 2001). The world subsequently went through a repeated pattern of glacial and interglacial times in the Quaternary. Although explanations of Quaternary ice ages are primarily based on variations of the Earth’s orbital geometry that changed the seasonal distribution of solar radiation at intervals of tens of thousands of years (Imbrie and Imbrie, 1979), events that occurred in the Tertiary€– such as the closure of the Panama Isthmus, the expansion of ice over Antarctica, the building up of ice sheets in temperate latitudes, the expansion of tropical grasslands replacing forests and the development of high plateaus€– also helped to create favourable conditions for a cold period (Ruddiman and Kutzbach, 1991; Raymo and Ruddiman, 1992; Quade et al., 1995; Derbyshire, 1996; Williams et al., 1998). For example, the closing of the Panama Isthmus prevented warm water from the Pacific entering the Atlantic to keep the North Atlantic warm. In the Quaternary, the temperate and polar areas of the world underwent tremendous changes due to alternating glacial and interglacial periods. In the tropics, the Quaternary was also a time of repeated change associated with cold and warm periods, although the effect of direct glaciation, unlike in the higher latitudes, was limited to high mountains. 287

288

Quaternary in the tropics

Such mountains included the Himalaya and Andes, the high volcanoes as in East Africa and isolated elevated peaks like Mount Kinabalu in Sabah, East Malaysia. On the other hand, repeated climate and sea-level changes have left strong impressions in the tropics. The Quaternary in the tropics is not as well studied as in the higher latitudes. Certain summary observations, however, can be made. The landforms and geomorphic processes in the tropics were mainly affected by three factors: • enhanced glaciation in the mountains • repeated climate change • repeated changes in sea level. These three changes operated simultaneously and often in an interrelated fashion. Large rivers tend to originate in high mountains, flow through a long valley and finally reach the sea beyond a delta or a coastal plain. The upper basin of such a river is affected by glaciations and climate change, the middle part in the valley by climate change and the lower part mostly by changes in sea level. Thus, different but simultaneous changes affected different parts of the rivers, and the river as a system had to adjust to all the changes. Such events may have happened multiple times. The evidences of the three major changes are incorporated in current tropical landforms, geomorphic processes, soils and subsurface alluvium. Events in the Quaternary explain the presence of relict alluvium in tropical river valleys, palaeochannels of major rivers, and the origins of deltas and coastal plains. These changes are discussed separately in the following pages, followed by two case studies which illustrate the simultaneous modifications of different parts in a geomorphic system. The two examples come from the Ganga River system and coastal Southeast Asia.

16.2╇ History and structure of the Quaternary The Pleistocene epoch of the Quaternary period covers approximately the last 2 million years and is marked by a number of glacial advances and retreats. The cooling started in the late Pliocene, but the Pleistocene is the period of repeated advances of glaciers separated by warmer times. From the middle of the nineteenth century, land-based glacial sediment and relict landforms contributed immensely to our knowledge about glaciations in the Pleistocene. But the sequence and dates were only worked out much later in the twentieth century, when cores from the ocean floor could be examined. Scientists examine cores from the deep oceans for various attributes. D. B. Ericson plotted the fluctuations in a planktonic foram (Globorotalia menardii) in sea-floor cores from the Caribbean. Its abundance varied at different depths; the deeper the core sample, the older was the material. As the foram is temperature-dependent, the fluctuation indicated a time-based pattern of alternating cold and warm intervals. When C. Emiliani plotted oxygen isotopes from the same cores, he noted a similar pattern of variation. Skeletons of marine organisms carry two isotopes of oxygen, O16 and O18, the same as seawater. O16 evaporates at a higher rate from the oceans than the other isotope, and when glaciers

289



Fig. 16.1

History and structure of the Quaternary

Estimated sea-surface temperature from the core V23–82, Northwest Atlantic. After Bradley, 1985 expand on land a progressively greater amount of O16 is locked up in the ice, and the oceans become richer in O18. As the climate becomes warm, the ice melts and the ratio of the isotopes in the oceans changes back in favour of O16. This type of change in the ratio of these isotopes in the core samples indicated the repeated advance and retreat of glaciers, of cold glacial and warm interglacial times. Further investigations standardised the series of changes, which is known as the marine oxygen isotope series (MIS). Dates for specific changes, numbered sequentially from the top, have also been established (Shackleton and Opdyke, 1973; Imbrie and Imbrie, 1979). Twenty-six isotope stages have been identified for the last one million years, and the number rises to 65 when these are traced to the beginning of the Pleistocene. It has also been established that the change from a warm to a cold climate was slow and gradual, but the other type of change, from cold to warm, was very swift, a characteristic that has been described as a sawtooth pattern. Figure 16.1 illustrates the last five MIS changes. The Holocene (the present interglacial) is MIS 1, the previous full-interglacial is stage 5e (128€000–111 000 years ago). In between, there are several glacial stages and warmer periods, which are not real interglacials. In the tropics, the repeated advance and retreat of glaciers only happened in the higher parts of tropical mountains. In comparison, the effect of the other two changes, climate and sea level, affected the entire tropics, although not with the same intensity everywhere. The landforms and sediments in the tropics are partly relicts from the Pleistocene. The modifications were not as direct as that of ice sheets moving across land surfaces in the higher latitudes, but more subtle, such as changing the discharge pattern and amount of sediment load in rivers. This in turn affected alluvial valleys and deltas. Our knowledge about such modifications is derived mostly from the last glacial advance which lasted approximately

290

Quaternary in the tropics

35 000–10 000 years BP, peaking roughly around 18 000. The previous climatic changes appear to have followed similar patterns.

16.3╇ Quaternary glaciation in the tropics Glaciers tend to advance during the cold stages and retreat during the warm ones. The snow line therefore moved up and down tropical mountains repeatedly, with the same happening for the upper limit of vegetation. Besides the low temperature, advancing glaciers require a considerable supply of moisture. This is illustrated by the limited movement of glaciers in many parts of the tropics around 18 000 years BP, identified as the Last Glacial Maximum (LGM). It was a very cold period, but the rather dry climate prevented glaciers from expanding in places such as the Himalaya. A current illustration is the location of the present glaciers on the high volcanic massifs of the Andes:€Cayambe, Antisana, Cotopaxi and Chimborazo. They are all on the wet windward eastern side (Clapperton et al., 1997). Van der Hammen et al. (1981) mentioned glaciers in the Colombian Andes reaching their greater extent earlier than the rest of the world, as precipitation was high in this location before the LGM but not later when the climate remained cold but very dry. In contrast, Himalayan glaciers advanced briefly during the Early Holocene, when the climate was not the coldest but considerable moisture was present. Since the Middle and Late Holocene, glaciers on tropical mountains have been in retreat. The Andes in South America went through periods of accelerated glaciation during the Pleistocene (Clapperton, 1983). The largest development of ice was in the south beyond the tropics, in Chile and Argentina. Enlarged cirques and valley glaciers, however, also developed in the tropical north, on the higher peaks. A number of the high peaks of the Andes rise to more than 5000 m and currently carry glaciers. The Andean snowline of the glacial period has been identified as 400–1000 m lower than the present. A much larger area was then under ice. A number of plateaus, higher than 4000 m, occur in the Andean Cordillera, plateaus which are ice-free at present. These supported small ice caps and glacier systems during the colder times in the Pleistocene. Batholithic massifs and folded ranges that rose to higher elevations developed large mountain ice caps. Huge troughs, up to 16 km long and 500–800 m deep were eroded into such massifs, in places exploiting zones of structural weakness. It has been suggested that during this period, lakes occupied a number of intermontane basins or tectonic troughs of the Andes. Such pluvial lakes have been attributed to meltwater from neighbouring glacial systems and higher precipitation. Some of these basins no longer carry a lake, whereas others, like Lake Titicaca, became smaller. Erosional benches and terraces round Titicaca indicate a past bigger area. The warmer Holocene probably started around 11 000 years BP in the Andes, taking a couple of thousand years to peak. This period of hypsithermal conditions was followed by a worsening of the climate, and the last 5000 years have been colder and marked by minor glacial advances. This pattern has changed from the beginning of the eighteenth century, resulting in a marked retreat

291

Quaternary glaciation in the tropics

of Andean ice (Clapperton, 1983). Clapperton has referred to the occurrence of massive arcs of both lateral and terminal moraines extending downvalley beyond the moraines of the Holocene epoch. Stadials commonly occur behind the older terminal moraines. During glacial times, ice accumulated not only on the Andes or Himalaya, which are very high mountains, but also on highlands as in Australia and Papua New Guinea. In the latter case, snow lines were lowered by about 1000 m. This led to the development of small cirques and valley glaciers, and extended periglacial activities, which now are identified as relict features (Williams et al., 1998). In Southeast Asia, the current snowline is at 4650€m. Glaciers occur above 4620 m on Mount Jaya in Papua (Indonesia) which rises to 4884 m. Currently, three very small glaciers, covering a total of about 7.5 km2 are found in these mountains (Allison et al., 1989). Even at this near-equatorial location, glaciation occurred during the last cold period, and at the LGM 18 000 years ago, ice covered 1600 km2 in Papua New Guinea and a small area of 5 km2 on the summit of the 4101 m high granitic pluton of Mount Kinabalu in Sabah (Hope, 2005). Hope (2005) suggested that the snowline was lowered by about 1000 m. Snow could have been common on volcanic peaks above 3000 m in the islands of Sumatra, Java, Lombok and Sulawesi, but these peaks do not seem to be high enough to support glaciers. Similarly, enhanced glaciations happened only on high mountains in Africa, which usually are volcanic peaks and not a continuous range. Knowledge of glaciations in Africa is fragmentary and the best account available is for the mountains of East Africa. Some of the high peaks, Mounts Kenya, Kilimanjaro and Ruwenzori, currently carry glaciers, but the geomorphological and sedimentological evidence indicate a more extensive glaciation that also occurred on the Aberdares and Mount Elgon, which are no longer under ice (Mahaney, 1989). Hastenrath (1984) estimated that a maximum possible total of 800 km2 was glaciated on these mountains. The present extent of glacial ice is about 10 km2. It is likely that multiple phases of glaciations occurred on the East African Mountains, separated by warm interglacial periods. Evidence of past glaciation has also been reported from the mountains of Ethiopia:€ over 600 km2 in the Bale Mountains, 160€km2 on Mount Badda and 10 km2 in the Simen (Messerli et al., 1980). These high areas no longer carry glaciers. Glaciation has been suggested on mountains in the Sahara, in the Drakensberg and the mountains of the Western Cape, but such suggestions are still at a hypothetical stage. Extensions of glaciers to lower elevations left their mark on the tropical highlands by eroding large U-shaped valleys that are currently drained by river systems. A number of valleys also carry a large amount of glaciogenic sediment, which now provides the rivers with considerable amounts of clastic load. It is also possible that the retreat of glaciers has produced pressure-release joints on the valley walls, leading to enhanced slope failures. A cold climate has left its mark on the mountain landscape and the sediment of headwater streams in parts of the tropics. The glacial–interglacial change happened a number of times, but in most cases, only the last glacial advance near the LGM and the sequence of events after that can be clearly inferred.

292

Quaternary in the tropics

16.4╇ Climate change Information regarding climate change in the Late Pleistocene and Holocene is derived indirectly. For example, it comes from ice and sea-floor cores, the reconstruction of vegetation history using pollens and the palaeohydrology of arid areas such as old lakes and river sediment. Collectively these are known as climate proxies. Computer simulations of palaeoclimates also provide valuable insights. In general, fluctuations in the climate of the tropics match the isotope stages. It has been suggested (though it is still not a definite conclusion and probably did not happen everywhere in the tropics) that glacial periods were associated with colder, drier, more seasonal climates, whereas interglacials were warmer and wetter times. A number of case studies from different parts of the tropics indicate a cooling of several degrees C on land, and even a cooling of the ocean waters. Based on terrestrial pollens, Van der Kaars and Dam (1995) inferred a 4–7°C drop in temperature in Java. Similar figures have been reported for the Amazon Basin, e.g. Colinvaux et al. (1996) indicated a glacial cooling of 5–6°C from pollen information. The tropical oceans were found to be cooler by approximately the same amount near Barbados (Guilderson et al., 1994), in the tropical Atlantic (Thompson et al., 1995) and the Southwest Pacific (Beck et al., 1997). Evidences for climate change are common in rain forests (Kellman and Tackaberry, 1997) and deserts (Williams et al., 1998), two completely different ecosystems. Of the three large blocks of rain forest, Indo-Malesian, African and South American, the IndoMalesian forest of South and Southeast Asia was the least affected, presumably because of its occurrence on peninsulas and islands where a wet maritime influence persisted to some extent. Controversy exists about the exact extent of the Amazonian rain forest during the glacial periods. It is likely that it retreated to a smaller area during glacial times and the relatively drier edges were probably replaced by grassland. Rain forests over West and Central African lowlands were probably the most affected. Vegetation change at such a scale would undoubtedly impact the regional river discharge, erosion and sediment transfer. Williams et al. (1998) described deserts as remarkable repositories of palaeoclimatic information. Such information comes from relict landforms that are the result of processes that have stopped operating or have changed. Evidence such as sediment from past lakes, the location of old and defunct river channels, fossil dunes and fossil fauna provide information about past climates in arid areas. As discussed in Chapter 12, landforms in deserts may be very old but that age does not necessarily reflect the surficial sediment (Williams et al., 1998). Climate change is reflected in such sediments and depositional forms created by them, e.g. sand dunes, palaeosols, and river and lake deposits. Collectively, these young forms and sediment indicate periods of wetter and drier phases in the Pleistocene. An onset of aridity from about 2.5 million years has been identified in several parts of the tropics. Large inland lakes started to dry out in the Sahara, Ethiopia, the Arabian Peninsula, India, Australia and southern Africa. Glacial periods in tropical deserts were arid times. Examination of equatorial Atlantic deep-sea cores indicates an association between dust from the Sahara, times of glacial maxima and low sea surface temperature (SST). Data from the Gulf of Aden and the Red Sea match this pattern. Williams, et al. (1998) included

293

Climate change

a summary discussion of this characteristic, using data from several sources, and recognising dry periods during glacial maxima and wet ones during interglacials, although a caution against oversimplification is included. It has been established, however, that at the LGM, 18 000 years ago, tropical areas were more arid and cold, trade winds blew more strongly, active desert dunes were much more widespread than their present distribution, and desert dust and loess were carried out and deposited on both land and sea. Rivers in the tropics, especially those that originate in high mountains and flow into the sea or an inland lake, also went through a series of adjustments during the Quaternary. Their base level changed repeatedly with the rise and fall of the sea level (as discussed in section 16.5); their discharge and flood patterns fluctuated due to climate change; and the amount and nature of their sediment also altered, matching a changing precipitation pattern and vegetation cover over their drainage basins. Evidence for this, at various scales, is derived from the subsurface alluvium in the river valleys, preserved palaeochannels, evidence of former large floods in slackwater deposits (Box. 16.1) and deltas. Deltas include both the present onshore one and, in many cases, an older submerged extension. The story varies for each river, and a case study on the Ganga is discussed in section 16.7. The size and volume of lakes also altered between the glacials and interglacials. This has been studied for many desert lakes, and evidences for lakes caused by meltwater and natural dams in glaciated mountain valleys are also available. In Australia, such former lakes are referred to as megalakes. Megalakes covered larger areas of a structural basin and the extra water was available due to a wetter climate (higher rainfall, reduced evaporation, increased cloud cover) or changes in runoff following a different pattern of vegetation cover in catchment basins (Bowler, 1981). Climate change in the tropics, however, is a complex story, further complicated by the feedback from changes in sea level. A higher sea level during a warm climate increased maritime characteristics over land and probably increased precipitation. A lower level in a glacial stage brought in continentality and a drier climate. Changes in climate continued in the Holocene. The Early Holocene was wetter in a number of locations including the Sahara and monsoon Asia. Changes in climate also influenced regional palaeohydrology, as seen in the alluvium and morphology of a number of rivers including the Amazon, Nile, Narmada and Murray Darling. Pollen studies from peat swamps (commonly on tropical mountain slopes) also indicate similar climate fluctuations in the tropics. Our knowledge regarding climate change is best for the last pair (Fig. 16.1) of glacial and interglacial times. Duplessy (1982) showed a weak southwest monsoon and a strong northeast one using oxygen isotopes of the planktonic formaminifera deposited on the floors of the Arabian Sea, northern Indian Ocean and the Bay of Bengal near the LGM, about 18 000 years ago. A number of rivers that drain to the Bay of Bengal€– Ganga, Brahmaputra and Salween€– carried reduced discharges during this period of weak southwest monsoon. In contrast, the southwest monsoon at the onset of the Holocene and the Early Holocene was a warmer and wetter time. Goodbred and Kuehl (2000a) estimated that between 11 000 and 7000 years BP, the sediment released by the Ganga–Brahmaputra system was double that of the present amount and this increased rate was sustained for 4000 years. Given that the present annual sediment discharges of these rivers total about a billion tonnes, the Early

294



Box 16.1

Quaternary in the tropics

Slackwater deposits Slackwater deposits tend to be fine-grained sand and silt that accumulated rapidly from suspension during large floods. These are deposited in areas where the velocity of a river is reduced, and are preserved in locations which are protected from erosion in subsequent floods (Baker et al., 1983). As slackwater deposits frequently contain datable material, usually for radiocarbon dating, and as several slackwater deposits may occur in a vertical sequence, the flood history of a river can be extended for a long period of time. It is thus possible to estimate the frequencies of past catastrophic floods. River gorges in both rock and indurated alluvium are potential environments for accumulating and preserving slackwater deposits. The commonest location is the mouth of tributaries where fine sediment is deposited by the backflooding of tributary channels from the main river. As the water backfloods, its velocity rapidly decreases and suspended sediment is deposited (Fig. 16.2). High-angled river confluences are particularly good for accumulating slackwater deposits. Other suitable areas include the lee of bedrock protrusions in the channel, shallow scallops in bedrock wall (rock shelters), abrupt width changes in channels and the inside of tributary meanders. Thin horizontal laminations are frequently seen in slackwater deposits because of their fine texture and process of deposition. Slackwater deposits can also be made of structureless fine sand or silt. Cross-bedded units with foresets indicating an upvalley movement of water from the main river are occasionally seen. Individual sedimentation units could be capped by fine-grained organic material. Other evidences, such as buried palaeosols, mudcracks and rapid changes in sedimentation pattern, can be used to separate the slackwater deposits of individual floods in a vertical sequence. Slackwater deposits have been used to reconstruct past large floods and compute their recurrence intervals. The flood history of a river can thus be extended beyond the recorded period. They are extremely useful for determining the flood climate of the Holocene, going back thousands of years.



Fig. 16.2 Diagrammatic sketch of accumulation of slackwater deposit (S) in the Katherine Gorge, Northern Territory, Australia.

From Baker et al., 1983. By permission of Wiley

295

Sea-level change

Holocene must have been a time of strong monsoons and accelerated erosion. A number of studies on lake sediments from the Thar Desert of western India similarly indicate a wetter Early Holocene when the lake levels were higher. This was followed by drier conditions and less moisture availability and ultimately the present-day arid conditions (Wasson et€al., 1984). Large-scale climate changes between the last glacial advance and the Early Holocene have also been reported from Australia. Kershaw (1978) reviewed the case of increased rainfall and a rapid spread of rain forest in northeastern Queensland. Enzel et al. (1999) demonstrated changes in climate in northwestern India during the Holocene from a careful analysis of sediments from the Lunkaransar dry lake in Rajasthan. The shallow lake levels fluctuated frequently in the Early Holocene and then abruptly rose to a higher level around 6300 C14 years BP. It dried out completely around 4800 C14 years BP along with a period of intense dune destabilisation. We should recognise that climate changed repeatedly even in the Holocene in the tropics. Wasson (1995) concluded that a peak in available moisture occurred in many regions between approximately 6000 and 8000 years BP. He listed monsoon Australia, India, Arabia, China, Japan, Taiwan and Thailand. The departures from this widespread pattern occurred in southern India and eastern China.

16.5╇ Sea-level change Repeated glacial advance and retreat during the Pleistocene were associated with the rise and fall of the sea level. The fall in sea level during the last glaciation, MIS 2, and the rise afterwards are the best-known. Similar instances have occurred with other isotope stages, but the history of such changes is less well documented. The falling sea level tends to be slow; a rate of 1–5 m per thousand years has been suggested by Williams et al. (1998). The rise is faster; the average rate is 5–10 m in a thousand years (Williams et al., 1998). It is generally accepted that the drop in sea level during MIS 2 was about 120 m, but the exact amount has been different in places. Areally, such a drop is most effective where shallow seas existed on cratons:€ exposing large areas of the sea floor, linking offshore islands to the mainland and providing land bridges. This happened globally, and a good tropical example comes from Southeast Asia where the retreat of the sea for hundreds of kilometres across the shallow Sunda Shelf linked mainland Asia with islands of western Indonesia and those in the South China Sea (Biswas, 1973; Batchelor, 1979). Similarly, the exposed Sahul Shelf connected northern Australia to Papua New Guinea (Fig. 16.3). Such changes not only affected coastal landforms, the distribution of mangroves and coral reefs, but on a bigger scale, they also affected regional climate, ocean currents, river systems, and the diffusion of flora and fauna. These changes can be reconstructed from species distribution and relict sediment. The story is more complicated along active coasts and coasts of volcanic islands, where corrections need to be made for post-change tectonic movements. It is a curious fact that clearer evidences for former sea levels tend to be recorded in such unstable regions, presumably because the regional uplift preserves a former sea-level marker by raising its elevation. Such records come from Papua New



296

Quaternary in the tropics

Fig. 16.3

Approximate extent of emergence of land around Southeast Asia (exposing part of the Sunda Shelf) and Australia (exposing part of the Sahul Shelf) during the last glacial maximum. The land connection permitted distant migration of flora and fauna from the mainland of Southeast Asia. Later, a rising sea level separated part of the land mass into islands, isolating such flora and fauna from the mainland. Note the major river systems on the Sunda Shelf. These drainage networks were connected with headwaters which now form detached river systems on land after the rising sea level inundated lower trunk valleys and alluvial sediments. The exposure of the Sahul Shelf connected Australia with Papua New Guinea and interrupted ocean currents and climatic parameters. The present pattern of ocean currents and climate started with the drowning of the exposed land by a rising sea level Guinea, Timor Leste (Chappell and Veeh, 1978), Sumba (Bard et al., 1996), Barbados (Bender et al., 1979), Haiti (Dodge et al., 1983) and other places. The most striking example of exposed coral communities due to a falling sea level during the Pleistocene comes from the Great Barrier Reef of Australia. The fall of the sea level left this line of reef as a ridge, approaching 160 m above the LGM sea level. The exposed limestone would have undergone active dissolution, giving rise to the development of caves and pinnacles. A subsequent rise in sea level inundated the reef, but its present appearance

297

Sea-level change

is thus partly due to geomorphic processes in operation during the former fall in sea level (Carter and Johnson, 1986). The distance from the sea could be responsible for reduced precipitation and an increase in continentality on the landmasses. Lower sea surface temperatures (SSTs), as mentioned earlier, may lead to less storminess, as warmer waters are required for a tropical storm to develop. Climate change also followed the diversion of ocean currents following the emergence of the former sea floor as a land barrier. The linking of northern Australia with New Guinea across the Sahul Shelf diverted the warm South Equatorial Current and isolated North Australia from an important supply of moisture from incoming maritime air (Fig. 16.3). As a result, aridity was enhanced in northern Australia (Williams et al., 1998). Afterwards, in the Early Holocene, summer moisture returned with a rising sea level, and rain forests replaced eucalyptus forests and woodlands in northeastern Australia (Kershaw, 1978). In some cases, a falling sea level exposed large land areas. Rivers extended their channels across this area to meet the distant sea. In the process they incised their lower courses due to the fall in base level and also linked up with other rivers to form bigger systems. Land bridges also allowed the migration of flora and fauna from the mainland to the present-day islands. The subsequent Holocene rise in sea level drowned the exposed continental shelf, disrupting the lower parts of the extended river systems, and isolating the migrated flora and fauna on islands. This happened in many places but with local variations. The case for Southeast Asia is discussed as an example in section 16.7. Current deltas at the mouths of rivers flowing to the sea were formed when the sea level started to rise at the beginning of the Holocene. Stanley and Warne (1994) constructed a model of delta-building using data from boreholes sunk into delta material. Radiocarbon dates from the base of deltaic deposits suggest that delta-building started between 8500 and 6500 years ago. The rise of sea level slowed down after an impressive start during which the exposed sea floor was inundated and the lower parts of former rivers were submerged. The slowing provided accommodation space at the mouth of the rivers, sediment was deposited and deltas started to build out. Thus, although a particular river may be old, its delta is Early Holocene, although the exact dates may differ between deltas (see Chapter 11). From LGM onwards (about 18 000 years BP), the geomorphology of the tropics needed to adjust simultaneously to both climate and sea-level changes. Climate change determined the supply of water and sediment to rivers either directly by controlling temperature and precipitation patterns, or indirectly by influencing the frequency of slope failures and vegetation cover. At the same time, the rise in sea level affected the lower course of rivers, when water and sediment, adjusted to the environment of the upper basin, arrived. Usually this is a complicated story which needs to be sorted out regionally. Figure 16.4 summarises the changing environment of South and Southeast Asia, using data from various sources including climate proxies (Kale et al., 2003). Two regional examples are summarised in sections 16.6 and 16.7. They illustrate the multiplicity of factors changing simultaneously in different parts of a geomorphic system. The first is a summary of the changing form and behaviour of the Ganga River from MIS 3 (58–24 000 years) to the present, based on a source-to-sink study of the river by Goodbred

298



Fig. 16.4

Quaternary in the tropics

Changes in sea level and climate over Southeast Asia and South Asia. The two changes operated simultaneously. Generalised from Kale et al., 2003 (2003). The second is an account of the sea-level change over the shallow Sunda Shelf in Southeast Asia (Gupta et al., 1987).

16.6╇ The Ganga River system:€Quaternary adjustments River systems as large as the Ganga integrate signals from a series of changes over a huge area. The headwaters of a large river system produce most of the sediment that the river transports. This is due to high precipitation on the windward side of the mountains and tectonic movements that provide uplift, steep slopes and fractured rocks. Low-gradient, mixed bedrock–alluvial valleys below the mountains mainly function as conveyor belts transporting sediment (Fig. 8.1). Such sediment is deposited near the sea in the form of a prism that extends to the highstand level of the sea. Incised valleys may extend under water from the high to low stands of the shoreline, marking the distance the sea retreated during glacial times (Castelltort and Van Den Driessche, 2003; Blum, 2007). Large rivers tend to be polyzonal, indicating that not all of their tributary basins respond in the same way to climate and other changes. The rivers may also traverse rock-cut and alluvial sections alternately, requiring an adjustment at each transition. During the Pleistocene, such rivers

299

Quaternary changes around the Sunda Shelf

had to adjust near-simultaneously to enhanced glaciation in the mountains, climate change over the basin and a falling sea level in glacial times. The direction of change reversed during the interglacials. The Ganga–Brahmaputra system is described in Chapter 9. The Ganga rises in the high Himalaya from the Gangotri glacier, flows eastward though the Himalayan foreland basin to meet with the Brahmaputra and Meghna, and with them builds the biggest delta in the world before emptying to the Bay of Bengal. The Himalayan tributaries of the Ganga flow through steep valleys, many of which have been glaciated and a number of which still carry glaciers. As it flows east, the Ganga collects tributaries from the north, the larger ones of which originate in the Himalaya. These rivers and the Ganga itself have deposited megafans at the contact between the Himalaya Mountains and the plains. The southern tributaries of the Ganga drain the Indian Peninsular craton and have no history of glaciation. The system is very seasonal as far as water and sediment discharges are concerned, and almost all the work of these rivers is carried out during the 4–5 months of the Indian southwest monsoon. The river was affected almost simultaneously by glaciations, climate change and sealevel change in different parts of the system; and the changes were different at each marine isotope stage. Goodbred (2003) summarised such changes from the high mountains to the floor of the Bay of Bengal (Table 16.1). The length of the river varied, and the extended river of glacial times left its mark on a submarine canyon, the Swatch of No Ground, and a very large submarine fan extending into the Bay of Bengal (Fig. 16.5). The summary table illustrates that temperature and precipitation in combination affected glaciers in mountains and erosion/deposition in valleys, the length of trunk streams fluctuated, and, over time, the centre of deposition shifted between deep-sea fans and deltas. The growth of the Bengal Delta is discussed in detail in Chapter 11.

16.7╇ Quaternary changes around the Sunda Shelf Sea-level cores from the middle of the South China Sea have revealed plant debris, oxidised claystone, mottled sediment and laterite wash, indicating a former marine regression that exposed a large area of the current sea floor (Biswas, 1973). Some controversy exists regarding how much of this shallow sea over the cratonic Sunda Shelf was lowered, but a figure of 120 m is generally accepted. Figure 16.3 indicates the extent of the retreat of the sea and the exposure of land. It is likely that the retreat of the sea to such an extent would bring a less maritime climate to the present coastal areas of Southeast Asia, which would thus have a drier and seasonal climate. Verstappen (1975) proposed a rainfall gradient with elevation, with reduced and seasonal rainfall in the lower areas. He postulated a tropical forest and grassland vegetation replacing part of the present rain forest (Verstappen, 1980). The exposed sea floor would have functioned as a land bridge, allowing the migration of flora and fauna. The existing river systems that now flow into the South China Sea would have extended across the exposed area and merged to form larger systems. At least three such systems have been identified (Fig. 16.3).

Slightly cooler and drier than present

Cold and dry

Warm and very wet, stronger monsoon

More relaxed monsoon, less rainfall

MIS 2/ LGM (24–18 ka)

Hypsithermal (18–7 ka)

MidHolocene (7 ka to now

Climate of the region

MIS 3 (58–24 ka)

Time

Low

Very high

High

Very high

High

High sediment transfer

Sediment

Low

Moderate

River discharge

Retreating glaciers, sediment removal and aggradation

Brief glacial advance. Erosion of valley fill

Limited glaciation and valley aggradation

Marked glacial advance and valley aggradation

Himalayan valleys

Possible aggradation

Probably inactive River linked with the Bengal Fan via canyon Aggradation shifting to delta

Less aggradation

Incision and possibly local aggradation

Inactive. Possible incision

Very rapid aggradation

Aggradation but slower

Aggrading megafans, upper Ganga plains and southern tributaries Largely inactive fans, minor incision, alluvium in southern tributaries Widespread rapid erosion, nearly to bedrock in valleys Low aggradation

Deep-sea fan

Delta and coastal shelf

Megafans and alluvial plains

Table 16.1╇ Source-to-sink response of the Ganga River system since MIS 3 (from Goodbred, 2003)

301



Fig. 16.5

Quaternary changes around the Sunda Shelf

Diagrammatic sketch showing the relative position of the catchment of the Ganga, the Ganga–Brahmaputra Delta, the Swatch of No Ground and the submarine Bengal Fan. Generalised from Goodbred, 2003 Molengraff (1921) recognised a North Sunda River flowing north between the presentday Malay Peninsula and the island of Borneo, collecting tributaries from both these regions. Emmel and Curray (1982) suggested that a group of rivers came down the western side of the Malay Peninsula, met rivers coming off east Sumatra and the joint system flowed northwestwards. They traced morphological changes along this system, identifying a braided pattern in the upstream area which was replaced downstream by a meandering channel that ended in a delta. A third major river flowed east, collecting left-bank tributaries from south Borneo and right-bank tributaries from northeastern Sumatra and northern Java. Afterwards, the rise in sea level inundated the exposed sea floor and truncated the river systems. The higher parts of the land appeared as peninsulas and islands leading to the

Quaternary in the tropics

302

(a)

(b)



Fig. 16.6

Geomorphological processes operating in the Blue Nile Basin during (a) the Last Glacial Maximum (global scale) and (b) the early Holocene. From Woodward et al., 2007. By permission of Wiley isolation of flora and fauna on various islands with no further communication between them. The coarse alluvium deposited by the braided headwaters of the past river systems can be found in the subsurface in many parts of Southeast Asia. For example, it is known as the Old Alluvium in Singapore (Gupta et al., 1987) and as the Simpang Formation in West Malaysia. Similar deposits have been found in Sarawak, Thailand and southern Viet Nam. It has been traced under the waters of the Malacca Strait, as a sedimentary cover lying over a granitic regolith (Aleva et al., 1973; Batchelor, 1979). A complete picture of what exactly happened, however, requires more investigation.

16.8╇ Conclusion The tropics, like the rest of the world, underwent repeated changes during the last two million years of the Quaternary. These changes included enhanced glaciations in the high

303

Questions

mountains, and changes in climate and sea level. The geomorphological effects of such changes were enormous (Fig. 16.6). The explanation of current landforms and sediments requires a careful investigation of such changes. It is also likely that the present climatic conditions are part of this sequence of changes, and geomorphology in the tropics has altered repeatedly throughout the Quaternary. A knowledge of what happened in the tropics during the Pleistocene is therefore essential for understanding the current landforms and processes (Nott and Hayne, 2001; Thomas et al., 2007) and, as discussed in Chapter 19, the knowledge of Pleistocene and early Holocene conditions in the tropics should help us in forecasting the possible geomorphic effects of climate change.

Questions 1. What kind of changes happened repeatedly in the tropics during the Quaternary? 2. List the mountainous areas where glacial ice accumulated during the colder phases in the Pleistocene. 3. How do we deduce the nature of climate change during Late Pleistocene and Early Holocene? 4. What is the evidence for a warmer Early Holocene? 5. It has been said that repeated glacial advances and retreats during the Pleistocene were associated with the fall and rise of the sea level. What effect did it have on the river systems in the tropics, starting from MIS 2? 6. It is generally accepted that during MIS 2, the drop in sea level was about 120 m. What kind of tropics did exist during this time? 7. Slackwater deposits are considered as important indicators of palaeohydrology. Why?

Part III

Anthropogenic changes

17

Anthropogenic alteration of geomorphic processes in the tropics Man has too long forgotten that the earth was given to him for usufruct alone, not for consumption, still less for profligate waste. George Perkins Marsh

17.1╇ The beginning Anthropogenic transformation of the physical environment changes processes and landforms. Large-scale alteration of the landscape started about 10 000 years BP, with the beginning of agriculture which led to the destruction and replacement of natural vegetation. Additional demands for construction, urbanisation and shipbuilding made further inroads into the forests. As natural vegetation was depleted, many areas were subjected to a changing hydrology and geomorphology. Rapp et al. (1972) provide a striking example from Tanzania (Fig. 17.1). Their study area was a pediment sloping at 3.5° from an inselberg to a floodplain. A grass-covered, thin, red, sandy-loam soil rested over the bedrock. The grass cover effectively prevented soil and water loss. A number of experimental plots with a variety of land use were established in mid-slope, and water and sediment loss measured from such plots. As Figure 17.1 shows, no sediment was lost from a plot under ungrazed thicket and surface runoff was equivalent to only 0.4 per cent of rainfall. When thickets were replaced with grass, the water lost by runoff increased to only 1.9 per cent of rainfall, but there was still no perceptible erosion. Losses of both water and sediment increased dramatically when millet was planted in rows or the land was left as bare fallow. This pattern of increased runoff and sediment loss happens with removal of vegetation everywhere; the numbers depend on the local environment. In 1864, George Perkins Marsh wrote Man and Nature, the title of which was changed subsequently to The Earth as Modified by Human Action:€ a New Edition of Man and Nature. He felt compelled to write this book in order to point out how much erosion and sedimentation have increased due to anthropogenic actions. Similar landscape changes have occurred in the tropical countries for a very long time, but it became considerably more large-scale and intensive in the twentieth century. Given the nature of tropical rainfall, once natural vegetation that protects the thick soil and weathered material in the humid tropics is destroyed, accelerated erosion and sediment transfer happen quickly (Fig. 17.1). The increased runoff and erosion marks the slopes with gullies and scars of mass movement, and the streams need to adjust to a new regime of water and sediment supply. In 307

Anthropogenic alteration of geomorphic processes

308



Fig. 17.1 Soil erosion with different land use covers from mid-slope at Mpwapwa, Tanzania. From Rapp et al., 1972.

By permission Wiley/Blackwell

many places, such excessive sediment even reaches the coast. The transformation of the landform and process is particularly effective where the rainfall is intense, slopes are steep and the region is tectonically unstable. Destruction of vegetation and alteration of the environment is mainly due to: • deforestation • changes in land use and rural migration • river impoundments • urbanisation.

The beginning

309

Table 17.1╇ A summary set of sediment measurements from Southeast Asia Type of land use

Sediment yield (tkm−2yr−1)

Forest Large river basins (mixed land use) Urbanisation Shifting cultivation Agriculture Logging, early stage

100–102 102–104, mostly 103 103 103 from plot measurements, 102 for entire basins 102–103, depending on conservation measures used 103, maximum recorded 17 000

Note Range of sediment yield compiled from a large number of published measurements in Southeast Asia. From Gupta, 2005a

The effect on the physical environment from the first three is reviewed in this chapter and the effect of urbanisation is discussed in the next one. Observations on anthropogenic modification of landforms and processes in the tropics have increased since the 1960s. When vegetation is removed, soil and regolith are de-structured, natural slopes are altered, and erosion rates and sediment yields reach figures which are accelerated and no longer natural. A combination of anthropogenic alteration and fragile landforms give rise to very high local sediment yields. The resulting changes may affect not only proximal slopes and rivers, but also distal deltas, beaches, mangroves and coral reefs. The changes have become large enough to be mapped from satellite images:€both the areas that are being eroded and also the resulting sediment plumes on the coast (Gupta and Krishnan, 1994; Gupta, 1996; Gupta, 2005a). The degree of change varies with location, and it is not always possible to make a quantitative measurement with reasonable accuracy and robustness. Erosion and sediment loss have been estimated for various locations using diverse methods, such as the calculation of ground lowering in small erosion plots (Box 17.1), sediment-sampling in streams and the rate of extension of a delta into the sea. Variations in local conditions and measurement techniques make it difficult to rely on a single figure to represent the effect of a particular type of anthropogenic change, such as deforestation or urbanisation. It is safer to use a range of figures than an average of all the measurements available. It can, however, be stated that the present rates of erosion and sedimentation are distinctly higher than earlier ones. This is illustrated in Table 17.1, which summarises sediment measurements in Southeast Asia. In spite of the generalisation involved, it is clear that sediment yield increases considerably once the vegetation is removed and the soil is exposed to erosion. The increased loss happens cumulatively over time and also with episodic jumps, as illustrated by the landslide devastation of logged slopes and lower valleys in southern Thailand in November 1988 after about 1000 mm of rain in five days. A similar incident of soil erosion, huge landslides and flooded rivers happened in the same year following deforestation in the Bengkulu Province, Sumatra. Therefore, the ongoing geomorphological processes and evolving landforms in many parts of the world should not be considered entirely natural. The current rates of erosion, sediment yield and building of sedimentary bodies such as bars in the river or beaches on the coast may differ from past ones. Destruction of vegetation and other anthropogenic modifications of the landscapes tend to degrade the environment and, at times, lower the

310



Box 17.1

Anthropogenic alteration of geomorphic processes

Measuring erosion rates and sediment loss Various techniques exist for measuring erosion rates from a geomorphic unit such as a slope. They range from very simple to rather complicated, and practices vary in efficiency and robustness. Sediment loss by erosion can be measured in two ways. It can be determined as a measure of ground lowering in the area undergoing erosion or the rate at which sediment is deposited at the lower end of the system. The measure of erosion or sediment accumulation, depending on which technique is used, is generally expressed in volume or weight of sediment per unit area, such as m3km−2 or tkm−2. Low sediment loss can also be expressed in kilograms and small areas are often calculated in hectares, e.g. kgha−1. As the sediment loss (also called sediment yield) is usually expressed as a rate, the total weight of sediment eroded in a year is expressed as tkm−2yr−1. Several common techniques are listed below, but there are others. For details of application in the field please consult a field manual. It is always good to have rates from several techniques on the same area so that an idea about reliability can be formed. 1. Erosion plots. Probably the easiest technique. An area on a sloping surface is marked off and long pins are inserted into the ground with washers below the pinhead lying directly on top of the ground surface. Over time, material is eroded from the surface and the washer drops down the pin. Thus the amount of erosion at various points of the marked area can be measured from which an overall rate can be computed. The problem is in extrapolating the data from such a small area to a region. 2. Experimental plots. A small area, usually at the scale of about 10–20 m2 is selected (Fig. 17.2). A perimeter trench is cut and partly filled with clay on top of which a metal sheet is planted which should extrude well above the ground surface. The plot is thus insulated from subsurface flow and rainsplash. The plot should be of trapezoidal or similar shape, so that water and sediment move towards the narrow end where storage tanks can catch sediment-laden water after a spell of natural or artificial rainfall. Both rates of runoff and

Fig. 17.2â•…Experimental plot for measuring runoff and sediment loss looking downslope. Sediment-bearing water is collected in the two containers in the background. The surface is currently bare. It is possible to control the amount of groundcover for experimental purpose. Experimental plot designed by Ausafur Rahman. Photograph: A. Gupta

311

Deforestation, land use changes and rural migration

sediment loss thus can be determined. The small-scale of the plot allows the plot land use to be changed (bare, grass, shrubs, etc.), which is an advantage, but it too has the problem of extrapolation to a large area. 3. Surveying. Periodic surveys can be carried out over an area and with reference to monumented points. This indicates which area is being eroded and where the sediment is going. This has been done with a wide range of instruments, from simple hand levels to total stations, presumably depending on the resolution required. 4. Suspended sediment in rivers. Suspended load in a stream can be measured in various ways, by using a vertically integrated sampler or a set of sample bottles fixed to a base and collecting at various depths of water. If the river does not carry much solution or huge amounts of bed load, the figure can be used to represent the load of the river, as bed load is less than 10 per cent in most rivers, below the precision of suspended load instruments. The amount is representative of a large area, the entire upstream drainage basin, but it provides a general figure and as such would not indicate the variation in erosion between the different parts of the basin. 5. Boreholes. Sedimentation rates can be determined by sinking boreholes in floodplains and deltas. Often it is possible to date sediment from carbon samples collected at various depths. This provides a rate of sedimentation which reflects upstream erosion, but this technique is applicable only for a long time-period, such as a delta built up since Early Holocene.

quality of life. A geomorphologist may use this knowledge to ameliorate environmental degradation associated with anthropogenic alteration of the landscape. This is one of the fascinations of tropical geomorphology.

17.2╇ Deforestation, land use changes and rural migration 17.2.1╇ Causes of forest destruction By the mid-twentieth century, forests of the world had been reduced to between one-half and one-third of the original extent (Eckholm, 1976). A substantial part of this was the tropical rain forest. Although other types of tropical vegetation were also destroyed to meet various demands, it is the accelerated destruction of the rain forest that attracted many researchers, including geomorphologists. Tropical light hardwood trees of the rain forest are in high demand for industrial purposes. These trees are tall, frequently festooned with creepers, and when they are felled they tend to bring down smaller trees with them. A very large area therefore is cleared when the tropical rain forest is logged. Mechanical extraction of such trees requires logging roads into the forest, which leads to further forest clearing. A limited amount of firewood is collected from the rain forest, but firewood collection is far more destructive for drier forests in the tropics. The collection of firewood is particularly disastrous near desert fringes or on mountain slopes. In both cases depletion of the vegetation cover significantly increases erosion.

312

Anthropogenic alteration of geomorphic processes

Rain forests are destroyed not only for their resources but also for their land. Large-scale land clearance was carried out in the second half of the twentieth century to settle migrant farmers. This has happened both under government planning and from the spontaneous migration of poor and landless people. A spread of pastures has destroyed forest in Central and South America. A variety of other demands and developmental policies such as mining and highway construction are also destroying tropical rain forest in various countries. The causes of forest destruction vary in importance from forest to forest. Rain forests have been destroyed for logging and plantation agriculture in Southeast Asia, unsustainable shifting cultivation in Central Africa, settlement and ranching in South and Central America, and firewood collection in drier areas and on the slopes of the tropical mountains. In the ten years between 1980 and 1990, it has been estimated that tropical forests have been lost at an annual rate of 174 000 km2 (FAO 1993, as reported in World Resources Institute, 1994). Another estimation by Trexler and Haugen (1994) put this �figure at 170€000 km2. The estimates are very close. The highest annual deforestation rates occur in two regions:€(1) continental Southeast Asia and (2) Central America and Mexico. It is the easily accessible lowland forest, commonly covering a rolling landscape, which was destroyed most between 1980 and 1990. The fastest rate of loss, however, was from hill and montane forests.

17.2.2╇ Effects of forest destruction Forest clearance on the slopes of the humid tropics gives rise to slope failures and gully formation, which periodically contribute large slugs of sediment to a river at the bottom of the slope. If deforestation continues over a wide area, such sediment would accumulate in a number of river channels, changing their flow pattern and physical character. In time, the eroded material from deforested slopes would reach the coast as plumes of sediment affecting coastal beaches, mangroves and coral reefs. A longitudinal research project, measuring water and sediment changes that followed alterations in land use started at the Danum Valley, Sabah, Malaysia in October 1987 (Douglas et al., 1992). The data were collected from hilly catchments between 0.5 and 10€ km2 in area and generally on easily eroded lithology. An unlogged catchment was compared with one in a logging concession, where road construction started in June 1988 and logging in January 1989. After storm rainfalls, both the discharge and sediment load increased in the stream which drained the logged basin. Sediment in the logged catchment increased to four times that of the natural catchment, after a logging road was built across the upper part of the catchment; it increased to five times after logging within 37 m of the road; and increased to a peak of 18 times in the five months after logging of the rest of the catchment. The amount of sediment lost was reduced a year after logging stopped. The combined pattern of accelerated supply of sediment that reaches a peak after logging and then tails off and logging roads operating both as water conduits and as a second sediment source due to gully formation can be taken to represent the general sequence of events that follow active deforestation. The logging tracks, cleared of vegetation and stamped hard by heavy machinery, often function as run-off channels following intense tropical rainstorms. Thus both the flood potential and sediment load of local streams

313



Fig. 17.3

Deforestation, land use changes and rural migration

Land use as main control in sediment production. Kenya. From Dunne, 1979. With permission from Elsevier increase. Sediment accumulates in the channels and floodplains of forest streams and slugs of sediment later travel down-channel following episodic rainfalls. Sediment from small coastal streams may reach the sea in a few years. Coastal plumes of sediment in Southeast Asia are now clearly visible even in low-resolution satellite imagery (Gupta and Krishnan, 1994). It is the same for other tropical regions. The cover of the book presents an excellent illustration. Post-logging sediment yield can be very high in the rain forest (103tkm−2yr−1) for the first one or two years and then it exponentially tails off as secondary vegetation returns to the bare slopes. The highest figures on record from deforestation are more than 7000 tonnes from Sabah, Malaysia (Greer et al., 1994) and 17 000 t from Java, Indonesia (Ruslan and Menan in Lal, 1987). It is difficult to choose a representative figure, but a number of more than a thousand tonnes is probably a reasonable estimate. Lai et al. (1996), working in the forests of Malaysia, suggested that the sediment yield from a logged catchment could be as much as 50 times of the yield from a natural basin.

17.2.3╇ Land use changes and sediment yield The effect of changes in land use on sediment yield was clearly shown by analysing sediment derived from 61 drainage basins in southern Kenya (Dunne, 1979). Although climate and topographic variables are important factors, it is the land use that primarily controls erosion and sediment yield in these basins (Fig. 17.3). A very large portion of the sediment is eroded during the wet period in the ambient seasonal climate and then transferred along rural roads. The roads behave as conduits for water and sediment transfer, and in consequence they become deeply eroded. This resembles the water and sediment transfer following deforestation. Sediment yield primarily depends on vegetation cover and basin steepness. Within a single type of land use, however, sediment yield indicates a direct

314

Anthropogenic alteration of geomorphic processes

relationship with runoff. The annual erosion rates in forested basins were 20–30 tkm−2. For farmed basins this figure varies enormously. The extreme values given by Dunne (1979) are about 10 tkm−2yr−1 for a flat basin on permeable lavas and low rainfall, and more than 4000 in wet and steep cultivated catchments. In sum, lack of vegetation cover, steep slopes and considerable rainfall is the worst combination for erosion and sediment transfer in this semi-arid landscape, and it is the removal of vegetation or insensitive farming that primarily increases sediment yield. It is generally accepted that sediment loss increases under agriculture, especially on hillslopes. Reliable measurements, such as carried out by Dunne and his colleagues in southern Kenya, however, are limited. For example, Douglas et al (1993) stated that the only valid soil erosion data for Southeast Asia are measurements by Hatch (1981) on 40 m2 plots in Sarawak. It is, however, difficult to extend such data to a regional scale. Weera and Kittipong (1987) computed figures for large areas in Thailand. Their work suggests that sediment yield of up to several thousand tkm−2yr−1 may be expected from river basins when the fields are open and no proper conservation measures are adopted. According to Murtedza and Ti (1993), such figures may be reduced to hundreds of tonnes when the entire river basin is considered, part of which could still be under forest. Agriculture therefore increases erosion and soil loss, but the numbers may vary widely depending on local conditions (Table 17.1). Moving and resettling a large number of people, as has happened in Malaysia, Indonesia and Brazil, also increases erosion and sediment load. Shifting agriculture is a farming system common among the local settlers of the rain forest who usually live in small communities. In shifting agriculture, a small tract of land is cleared for farming. Such areas are abandoned after several years as the fertility of the land falls, and new tracts are cleared for the next batch of farming. A return to the originally cleared area should not happen for years until the soil recovers its fertility. Such a rotation with a reasonable number of years also allows the soil to repair the damage from erosion. Shifting agriculture has thus been described as a rotation of fields rather than a rotation of crops. Used within its constraints, it is a sustainable system. As the area of the rain forest decreases, it becomes difficult for the forest population to properly rotate the fields, and a quick return to old areas or longer occupations leads to accelerated erosion of the surface material.

17.3╇ Temporal and seasonal patterns of sediment transport Erosion of the regolith and associated sediment yield from a disturbed site, however, does not continue at a uniform rate but changes over time (Douglas et al., 1992). Data on the impact of deforestation from their research sites in Sabah show that sediment yield increased to 18 times the pre-forest disturbance rate, but started to tail off in 2–3 years. Secondary vegetation grows rapidly in the humid tropics, and after several years the yield figures dropped to numbers representative of undisturbed areas (Greer et al., 1996). This is a common pattern and unless the disturbance to the natural vegetation continues, secondary vegetation covers up the bare ground and the erosion rate and sediment yield decrease.

315

Spatial transfer of sediment

The adjustment of slopes and rivers to changing geomorphic processes thus continues over time. Temporal and seasonal changes in erosion and transfer of sediment were described for a 225 km2 area of northern Lao PDR that is drained by the Mekong River (Gupta and Chen, 2002). This is an area of steep-sided ridges and valleys with the ridge crests rising to over 1000 m, where the Mekong flows at an elevation of 250 m within 5 km of the ridge crests. The annual rainfall is around 1700 mm. The rain comes mostly from the southwestern monsoon, with 85–90 per cent falling between May and October. The steep slopes are generally under forest (primary or degraded) or shifting cultivation. Plots are cleared for shifting cultivation at the beginning of the dry season in December to January and the biomass on the ground is burnt later (April to May). The rural population is low (tens per km2) but the cultivated slopes are very steep. Areas under vegetation or cleared were mapped repeatedly from multispectral SPOT satellite images, as such measures changed every month. Multiplying the values of these areas with a measure of average sediment yield from the literature (Table 17.1), allowed estimation of the sediment removed from the slopes on a monthly basis. This is only a semi-quantitative estimate, but it does indicate (Table 17.2) that (1) a seasonal pattern of erosion and sediment transfer and storage occurs each year, and (2) the amount of sediment loss and transfer is directly related to rainfall. This exercise indicates that erosion and sediment yield vary between years. For example, significantly less sediment would be released in a dry El Niño year. Working in the Philippines, White (1995) attributed erosion, short-distance sediment transfer and its storage to frequent thundershowers; and efficient long-distance transfer of sediment to tropical cyclones, which occur at much longer intervals. Sediment production from farming on steep slopes would therefore produce a fluctuating, but not necessarily cumulative, sediment production. This trend will change with continuous cultivation, without proper conservation measures. The rivers of many parts of the tropics therefore repeatedly adjust to such fluctuations in hydrology and sediment yield. The changing characteristics of drainage basins indicate that such streams may no longer be entirely dependent on natural conditions.

17.4╇ Spatial transfer of sediment Latrubesse and Stevaux described the connection between accelerated sedimentation in the upper basin and channel form and behaviour downstream for the Araguaia Basin of central Brazil, where the ‘intensive and indiscriminate expansion of agricultural activities’ (Latrubesse and Stevaux, 2002:€109) have degraded the natural environment. Gullies have rapidly increased in size in the upper basin (Fig. 17.4). Large gullies (locally known as voçorocas), hundreds of metres long and with headcuts that are more than 20 m wide, developed in 30 years. These are formed by downcutting, piping and failure along pressure-release joints in alluvial fills. Latrubesse and Stevaux referred to a study by Silva in which 97 large and recently formed gullies were identified in about 3 km2. The sediment produced by the easy erosion of thick soils and weathered material of the upper basin leads

Anthropogenic alteration of geomorphic processes

316

Table 17.2╇ Middle Mekong Basin:€seasonal erosion and sediment transfer Period

Slopes

Tributaries

The Mekong

Dec–Apr Dry, a few showers

Increase in clearance, very little erosion of sediment transfer due to lack of rainfall

Low flow; past sediment stored mainly on bed, as bank insets and at confluences

May Rains start, rising discharge

Slopewash, gullies, small debris flows; sediment transfer to channels

Transfer of sediment pulses following showers, sediment arriving from slopes

Approximately Jun–Jul Early wet season

Considerable erosion and sediment transfer to channels by processes described for May

Transfer of tributary channel sediment to the Mekong; more sediment arrives from slopes; accumulation at confluences

Aug–Nov Late wet season

Growth of vegetation; reduced erosion; very little sediment transfer

Continuous transfer of sediment already in channels, accumulation at confluences

Sediment stored on beds, as bank insets, and at confluences with major tributaries. The river stage falls gradually and sediment transfer diminishes Increased river velocity and transport of channel sediment, increased competence Erosion of bed material, insets and bars; increased competence; bedrock erosion; localised critical and supercritical flows; water tens of metres deep in the inner channel. Continued erosion of bed material, insets and bars; increased competence; bedrock erosion; localised critical and supercritical flows; deep water in the inner channel; sediment accumulation at confluences

Source:€Gupta and Chen, 2002

317



Spatial transfer of sediment

Fig. 17.4 Gulley eroding alluvial fill in the basin of the Upper Araguaia River, Brazil. Photograph: A. Gupta

to dynamic changes in the middle river with localised aggradation and bank erosion. The extensive deforestation and land use changes in the upper basin have changed the fluvial dynamics of the Araguaia. Anthropogenic disturbances that caused the destruction of the original vegetation cover allow intense tropical rain to fall on bare slopes mantled with thick soil and weathered material. The usual sequence is a large number of small failures on hillslopes along with an increase in the number and size of gullies, the accumulation of clastic sediment in stream channels and floodplains, and an increase in the flood potential of rivers (Fig. 17.5). The destruction of forests over a large area and on steep slopes gives rise to extensive slope failures. Large landslides and debris flows occur, especially after a period of heavy rainfall (Chapter 6). A huge volume of sediment reaches the rivers at the foot of the slopes. These rivers tend to flood more frequently because of the alteration in the water balance in their drainage basins following widespread deforestation. This is followed by accelerated growth of bars and floodplains. Mapping of the slopes near the Mekong River in Lao PDR from 20-m resolution SPOT satellite images (Gupta et al., 2002) indicates vegetation clearance in structure-guided tributary valleys leading to the pulsatory passage of sediment to the Mekong as described in Table 17.2. Fresh young sand is seen in the Mekong as insets (Fig. 17.6), around rock protrusions in the channel and as river-mouth fans where tributaries join the main river. A very large volume of sediment is making its way to the coastal waters, although a large part of it may remain stored for a considerable time in the alluvial lowlands of the Lower Mekong in Cambodia and the delta in Viet Nam. Sediment plumes have been measured off the mouth of the Mekong from satellite images, the seasonal variation in the size of the plumes being related to the varying discharge of this monsoon river (Gupta et al., 2006). Such plumes occur along coasts where inland basins are undergoing considerable changes in land use and at the mouths of major rivers in South and Southeast Asia (Gupta and Krishnan, 1994). They are common along large sections of the coastline

318

Anthropogenic alteration of geomorphic processes



Fig. 17.5

Diagrammatic sketch of a river system with vegetation destruction in its upper basin



Fig. 17.6

Fresh sand in the Mekong River as inset near Luang Prabang, Lao PDR. From Gupta, 2007. By permission of Wiley in many parts of the humid tropics. The IKONOS image on the front cover of this book illustrates the discharge of such sediment-laden river water to the sea. Anthropogenic alterations of the drainage basins therefore modify the hydrology and sedimentation on slopes and along stream channels. In most cases, such alterations lead to an episodic and floodprone pattern of discharge and increased storage and transfer of sediment. Impoundments across rivers, such as dams and reservoirs, modify drainage directly, leading to a changed hydrology, reduced sediment transfer and channel modification.

319

Impoundments along rivers and their effects

17.5╇ Impoundments along rivers and their effects Dams act as barriers to the down-channel passage of water and sediment, forcing rivers to adjust their morphology and behaviour (Kondolf, 1997). Such barriers have been used for thousands of years to control rivers. About 5000 years ago, King Menes of Egypt used clay bunds to hold back the floodwaters of the Nile for basin irrigation when the annual flood receded. Since then, the number of dams has increased significantly. The International Commission on Large Dams (ICOLD) estimates that around 45 000 large dams exist in the world, about 35 000 of which have been built since 1950. According to ICOLD, a large dam is at least 17 m high from its foundation to the crest. Dams 5–17 m high and with a reservoir capacity of more than 3 million m3 of water are also listed as large dams (World Commission of Dams, 2000). Such a large number of dams cumulatively store a very large volume of water and sediment in their reservoirs. A 1987 World Bank study indicated that about 50 km3 of sediment was being trapped annually in the reservoirs of the world (Mahmood, 1987). Following the trend in Western countries, a very large number of dams, some huge in size, were constructed across tropical rivers in the second half of the twentieth century. The practice continues, although big dams are now hardly built in the developed world. As many countries in the tropics achieved independence and concentrated on development, dams and reservoirs were seen as instrumental in providing water for irrigation and power, and also for flood control. They were not only engineering marvels but were also perceived as essential symbols of national economic development. As a result, the number of dams in the tropics proliferated, their sizes also increased. Vörösmarty et al. (2003) reviewed the global impact of sediment retention from river impoundments, mostly using data from 633 large reservoirs. They identify a large reservoir as one with at least 0.5 km3 maximum storage capacity. These reservoirs impound runoff and trap sediment. Half of these reservoirs have a local sediment trapping capacity of 80€per cent or more, and the total volume is rising. The trapping tends to be more efficient in drier areas. Arid Asia, Australia/Oceania and Africa show a tendency towards high trapping. The scale of trapping is less for South and Southeast Asia and parts of South and Central America (Vörösmarty et al., 2003:€181, Fig. 5). The global sediment retention in regulated basins was a modest 5 per cent of the moving sediment in 1950. This figure had risen to 15 per cent by 1968 and to 30 per cent by 1985. Apparently, there has been a stabilisation since then. This amounts to an estimate of 4–5 Gt annually. Vörösmarty et€al. (2003) estimated that this has reduced the annual global sediment flux that reaches the oceans to 10–16 Gt. If all the reservoirs, including smaller ones, are included, this estimated total would be reduced even further. Impoundments on rivers therefore significantly modify fluxes of water, sediment, carbon and nutrients on a global scale. This volume of impeded sediment increases cumulatively each year, as does the number of large dams built or under construction, which includes huge examples, such as the Three Gorges Dam over the Changjiang and the Xiaowan Dam over the Lancang Jiang (Upper Mekong). Under normal circumstances all this sediment would have built bars and floodplains downstream and ultimately reached the coast. Rivers downstream from dams therefore need to adapt to a decrease in sediment and a controlled water release. This alters

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Anthropogenic alteration of geomorphic processes

Fig. 17.7 Examples of sediment change downstream of dams. From Williams and Wolman, 1984. Courtesy of USGS

their form and function. Rivers in alluvium and rock tend to adjust differently. We need to determine the adjustment of rivers downstream of dams.

17.5.1╇ Downstream effects of dams on alluvial rivers Williams and Wolman (1984) discussed changes downstream from 21 dams constructed on alluvial rivers in the United States. The listed changes are applicable to impounded alluvial rivers globally. The study used very good records on discharge, sediment, channel and floodplain morphology, and valley-floor vegetation. The records covered a number of years. It was possible to trace changes from the pre-dam conditions to the post-dam environment, over time. The major findings of the study are listed below. • Flood peaks generally decrease after the construction of dams, but the effect on other post-dam discharge characteristics varies between rivers. • Sediment concentrations and suspended loads decrease markedly for hundreds of kilometres downstream from dams (Fig. 17.7). • This results in bed degradation and a lowering of the longitudinal profile, although the level of degradation varies. Most of the degradation occurs during the first decade or two after the dam closure (Fig. 17.8).

321



Fig. 17.8

Impoundments along rivers and their effects

Degradation and long profile change downstream form dams. From Williams and Wolman, 1984. Courtesy of USGS • Channel width can increase, decrease or remain the same downstream from the dam. • Bed material may coarsen in the initial stage after dam closure, but this trend may change over time. • Riparian vegetation commonly increases downstream from dams, probably because of decreased peak flow. A river downstream of a dam therefore carries a more uniform and reduced discharge and less sediment. This results in the deepening of the channel and erosion of its bed. If the riparian vegetation builds up next to the river, then it binds fine material to build bigger bars and floodplains, which are often common features below a dam in an alluvial river. The down-dam channel pattern also may need to change to adjust to variations in discharge and sediment. Williams and Wolman showed that rivers may require a distance of about 500 km to ameliorate these changes. Such amelioration may also happen over a shorter distance, depending on the scale of the river and the alteration due to impoundment.

Anthropogenic alteration of geomorphic processes

322



Fig. 17.9 Changes in flood discharge and variability due to Glen Canyon Dam on the Colorado River. From Webb et al., 1999.

Courtesy of AGU

17.5.2╇ Downstream effect of dams on rivers in rock Very few detailed case studies exist on the downstream effect of dams on rivers in rock. The case study on the Glen Canyon Dam on the Colorado River upstream of the Grand Canyon is an excellent example. The pre-dam Colorado was a river with large summer floods which moved a considerable amount of sediment in the channel. The post-Glen Canyon Dam discharge has become seasonally uniform and floods no longer enter the Grand Canyon. Flow regulation has also reduced discharge variability from year to year (Fig. 17.9). This reduced the competence of the river downstream from the Glen Canyon Dam, resulting in the aggradation of debris fans at the mouth of tributary streams in the canyon and a changing channel morphology in places, particularly at rapids. Very little

323

Application of geomorphology towards a better environment

sediment is now carried by the Colorado, as most of the former high sediment load is deposited in Lake Powell, the reservoir behind the Glen Canyon Dam. The texture of the sediment is also proportionally much finer in the canyon. As a result, bed scour has occurred and the bed of the Colorado has been armoured with coarse sediment. The river also erodes sand bars which are no longer replenished, as the former sediment-carrying, bar-building floods are absent. This has resulted in a degradation of ecological habitats, changes in flora and fish species in the river, and the destruction of tourist campsites (Webb et al., 1999). An experimental flood was sent down the canyon on 26 March 1996 by releasing water from the dam to see whether it had any beneficial effect. The experiment has been repeated since then. Obviously damming rocky streams, as commonly done in the mountains, degrades rivers. Both sets of studies are from outside the tropics, but there is no reason to expect dammed tropical rivers to behave differently. However, good studies of the effect of dams on geomorphology of tropical rivers are rare. This is unfortunate, as a very large number of dams, several of which are huge, have been constructed in tropical countries as part of development projects. The Zambezi Basin, for example, carries a series of dams on both the mainstem and the tributaries:€Kariba, Cahora Bassa, Itezhi-tezhi and Kafue Gorge. Multipurpose dams and reservoirs are necessary to meet growing demands for power and irrigation, but it is also necessary to determine the effect of such impoundments on the river and cumulatively on the coastal waters of the tropics. Dams can influence river morphology and behaviour over a long distance. For example, China is constructing a cascade of eight dams on the upper parts of the Mekong River that originates in China. It is suspected that the impact of the dams will mostly be on the lower river, thousands of kilometres downstream (Gupta et al., 2006; Kummu and Varis, 2007; Kummu et al., 2010).

17.6╇ Application of geomorphology towards a better environment Only a fraction of the landforms of the world can claim to be natural, whether inside or outside the tropics (Fig. 4.8). Tropical landscapes have been settled for a long time and, since the middle of the twentieth century, the pressure of high population and the need for developing the economy have strikingly modified the landforms and operating geomorphic processes. Understanding the processes across hillslopes, along rivers and on beaches therefore requires understanding the effect of anthropogenic modifications. This is also necessary to manage rivers or ameliorate slope failures. This chapter is a general introduction to the applied aspect of tropical geomorphology. In the next two chapters we focus on specific problems:€geomorphology in urban areas, where the intensive use of land creates a special type of environment with its associated hazards; and the effect of climate change on tropical rivers. The rapid anthropogenic changes of the tropical environment provide geomorphologists with a problem-solving capability which unfortunately is seldom activated.

Anthropogenic alteration of geomorphic processes

324

Questions 1. Explain how a combination of anthropogenic alterations and fragile landforms lead to high sediment yield. Give several examples where this is prevalent. 2. How far does the effect of anthropogenic alteration in the upper valley extend down a long river? 3. What are the effects of deforestation on the slopes of the humid tropics? 4. How did Dunne demonstrate that land use is the primary control of sediment loss in small river basins in southern Kenya? What type of climate prevails over this area? 5. There is a widespread belief that shifting agriculture leads to accelerated erosion and high sediment yield. Is it just a belief or the truth? 6. What effect would the time-based variation in sediment yield have on local rivers following deforestation? 7. What effects do dams have on

(a) the longitudinal profile of a stream downstream from the dam? (b)╇ changes in suspended and bed load downstream from the dam? (c)╇ sediment that used to reach the coast?

8. How would you measure the rate of erosion of a place? 9. Can geomorphologists play an important role in managing the environment?

18

Urban geomorphology in the tropics

The mountain held the town as in a shadow. Robert Frost

18.1╇ Introduction to urban geomorphology Urban geomorphology evaluates the impact of urbanisation on landforms and the ambient geomorphological processes of an area. It focuses on alterations to the physical environment caused by the spread of towns and cities. Urban geomorphology developed from a number of case studies in the developed countries (mainly the United States) in the 1950s and 1960s. Wolman (1967) structured a three-stage sequence of changes in an area undergoing urbanisation: • Stage 1. Pre-Urban:€land under natural vegetation or agriculture; river channels adjusted to existing conditions. • Stage 2. Brief period of construction:€ land stripped bare; soil and weathered material disturbed; extensive erosion; channels receiving large quantities of sediment; channels in disequilibrium. • Stage 3. Post-construction urban landscape:€impervious surfaces of streets, parking lots, rooftops, etc; concrete drains and sewers replacing natural drainage to streams; insignificant erosion and sediment supply to channels; river channels still in disequilibrium but over time may adjust to new conditions or persist in a state of disequilibrium. Patterns of water and sediment supply from drainage basins change after urbanisation (Wolman, 1967; Leopold, 1968; Dunne and Leopold, 1978). For a brief period during construction, an impressive amount of sediment is moved from the land, building bars, dunes and tributary-mouth fans in channels and constricting their capacity, thus making such channels floodprone. The subsequent establishment of an impervious cover significantly reduces infiltration and increases runoff. The degree of imperviousness may range from near 20 per cent in low-density residential areas to about 90 per cent in the central business district. The transfer of water to drainage channels is further accelerated by the construction of efficient storm sewers that take the water off the street and convey it speedily to drainage channels The percentage of area in a city served by storm sewers may range from 0 to nearly 100 depending on the degree of urban development. This reduces the lag time between the peak in rainfall and peak in channel flood (Fig. 18.1). Dunne and Leopold (1978) suggested that on average a partly sewered basin has a lag time of about 20–25 325

326



Fig. 18.1

Urban geomorphology in the tropics

Shortening of lag time between rainfall and runoff due to urbanisation per cent of the rural value. The increasingly impervious nature of the drainage basin also reduces base flow, often dramatically. Urban channels therefore undergo a set of modifications as listed in Table 18.1. If left in their natural state after urbanisation, they flood more frequently and increase their cross-sectional area, mostly by bank erosion. Usually, however, they are modified by engineering, which involves straightening, widening and lining the channels. They become more efficient for transferring floods downstream. Leopold (1968) has shown that urbanisation increases the frequency of floods of a given size. Based on the records of the US Geological Survey he listed several estimates of increase in discharge (Fig. 18.2) from a drainage area of one square mile (2.59 km2). For example, the mean annual flood (defined as the flood with a recurrence interval of 2.33 years) increases to between two and three times its former value in a basin with 50 per cent storm sewers and 50 per cent impervious cover, a situation often found in medium-density residential areas. Urbanisation, however, does not seem to change the size of large meteorological floods. A large volume of rainfall, as from a tropical storm, is transferred rapidly into surface runoff, irrespective of any land use. It is the smaller floods that illustrate the effect of urbanisation. Such hydrological changes cause unlined urban channels to erode their beds and banks, and expand their cross-sectional areas. Such channel erosion may produce as much sediment annually as derived cumulatively over five years from an undisturbed basin (Leopold, 1968). Various measurements indicate that channel enlargement could range between 50 and 300 per cent. Most urban channels left in a natural state probably tend to look like Wolman’s description of an urban channel that reflects an erosive regime:€flood debris, eroding banks, scour holes and exposed bars. The effects of urbanisation are heightened in the humid tropics for several reasons:€the presence of thick weathered regolith that is no longer protected by vegetation; the intensity and high amount of tropical rainfall; and the ongoing progressive urbanisation of the

Introduction to urban geomorphology

327

Table 18.1╇ Geomorphological changes in a drainage basin following urbanisation Type of change

Description

Hydrological changes

Changes in flow duration modification of flow–duration curve:€lower base flows and larger and quicker floods Changes in flood frequency increase in frequency of floods of a given size increase in size of a given flood, e.g. mean annual flood Modification of runoff from individual storms decrease in lag time between rainfall and streamflow increase in size of flood peaks increase in peak velocity Changes in water quality Changes in hydrologic amenities Deposition of relatively coarse-textured sediment in channels in the early stages of urbanisation, followed by a reduction in the amount of sediment reaching the channels Changes in size, width and depth of channels Changes in depositional forms Changes in drainage density

Sedimentological changes

Morphological changes



Fig. 18.2

Changing flood–frequency curves for a hypothetical 1 square mile basin at various stages of urbanisation. Note figures are in FPS system. From Leopold, 1968. Courtesy of USGS

328

Urban geomorphology in the tropics

tropical countries. We are familiar with the first two factors. The next section reviews the case of rapid urbanisation in the tropics. In general, urbanisation increases slope instability, accelerates flooding, increases sediment production and, in certain cities, results in ground subsidence. All such impacts are discussed in this chapter. The problems are made worse if the city is located in a naturally hazardous landscape (see section 18.3).

18.2╇ Urbanisation in developing countries Urban settlements in the tropics started at least 5000 years ago, but a number of major modern cities date back only to colonial times. In developing countries, a number of cities started as the trading or military posts of a foreign power a few hundred years ago, and subsequently evolved into large cities. Quite a few industrial and port cities are even younger, having grown up in recent years, largely after independence from colonial powers. The location of a number of major tropical cities can be justified in economic terms with respect to their hinterland but not their actual sites, which often are fortuitous. Kolkata started on the eastern bank of the Hooghly River at the location where the boatmen of the British trader Job Charnock managed to moor their boat at the end of a rainy day in 1690. Nairobi grew up in the middle of a swampy plain where a tented depot of railway stores was established in 1899 as the rail line was being laid across East Africa. Kingston (Jamaica) developed when the residents of the nearby Port Royal left their homes following the destructive earthquake of 1692. Once established, a number of these centres prospered and grew into large cities with millions of inhabitants. Some are still growing rapidly. The ongoing rapid rate of urbanisation in the tropics is also a modern phenomenon. The urban population of the less-developed world, roughly synonymous with the tropics, increased from 310 000 million in 1950 to 2.38 billion in 2007, and is projected to rise to 5.33 billion by 2050 (United Nations, 2008). The urban inhabitants, however, are unevenly distributed between cities of different size. Just over half of the urban population in the less-developed region live in small urban centres with fewer than 500â•›000 people. At the other end of the scale, megacities (defined as cities with more than 10 million people) carry about 10 per cent (United Nations, 2008). The number of megacities in the world was 19 in 2007, a number which is expected to rise to 27 in 2025. The number of megacities is increasing in the tropics (Table 18.2), and such cities are expanding to cover a very large area. This phenomenon is not confined to very large urban agglomerations. Urbanisation is happening across the range and almost all cities are expanding. Table 18.2 indicates that tropical cities are growing faster than cities outside the tropics. Of the 17 megacities in 2000, 10 were in the tropics. The pattern probably also holds for smaller cities. This implies an impressive expansion of built-up areas replacing rural or natural landscapes. For example, Sâo Paulo covered an area of 180 km2 in 1930. By 1988 its area covered more than 900 km2 and its metropolitan region had spread over 8000 km2. Usually, but not always, the urban periphery records the fastest growth, an area that is not always suitable for city expansion. The urban fringe of Jakarta, on the north coast of Java,

Urbanisation in developing countries

329

Table 18.2╇ Ranked distribution of twenty largest urban agglomerations, 1950–2025 (Approximate population in millions within parentheses) Rank

1950

1975

2000

2025 (projected)

1 2

New York area (12.3) Tokyo (11.3)

Tokyo (26.6) New York area (15.9)

Tokyo (36.4) Mumbai (26.4)

3

London (8.4)

Mexico City (10.7)

4

Paris (6.5)

Osaka-Kobe (9.8)

5

Shanghai (6.1)

Sâo Paulo (9.6)

Tokyo (34.5) Mexico City (18.0) New York area (17.8) Sâo Paulo (17.1) Mumbai (16.1)

6

Moscow (5.4)

Los Angeles area (8.9)

Shanghai (13.2)

7

Buenos Aires (5.1)

Buenos Aires (8.7)

Kolkata (13.1)

8 9

Chicago (5.0) Kolkata (4.5)

Paris (8.6) Kolkata (7.9)

10

Beijing (4.3)

Moscow (7.6)

11

Osaka-Kobe (4.1)

Rio de Janeiro (7.6)

12

Los Angeles area (4.0)

London (7.5)

13 14 15 16

Berlin (3.4) Philadelphia (3.1) Rio de Janeiro (3.0) St. Petersburg (2.9)

Shanghai (7.3) Chicago (7.2) Mumbai (7.1) Seoul (6.8)

Delhi (12.4) Buenos Aires (11.8) Los Angeles area (11.8) Osaka-Kobe (11.2) Rio de Janeiro (10.8) Cairo (10.5) Dhaka (10.3) Karachi (10.0) Moscow (10.0)

17

Mexico City (2.9)

Cairo (6.5)

Manila (10.0)

18

Mumbai (2.9)

Beijing (6.0)

Seoul (9.9)

19 20

Detroit (2.8) Boston (2.6)

Manila (5.0) Tianjin (4.9)

Beijing (9.8) Paris (9.7)

Delhi (22.5) Dhaka (22.0) Sâo Paulo (21.4) Mexico City (21.0) New York area (20.6) Kolkata (20.6) Shanghai (19.4) Karachi (19.1) Kinshasa (16.8) Lagos (15.8) Cairo (15.6) Manila (14.8) Beijing (14.5) Buenos Aires (13.8) Los Angeles area (13.7) Rio de Janeiro (13.4) Jakarta (12.4) Istanbul (12.1))

Note:€(1) Increase of population and (2) the rising number of tropical cities among the top twenty over time. Source:€United Nations, 2008

is spreading at a pace of up to 18 per cent annually; a part of this expansion is replacing coastal swamps, the other part is moving inland upslope beyond the upper boundary of alluvial fans into the hills. Development also alters previously unwanted degraded land inside the city limits. For Sâo Paulo, this has meant steep vegetated slopes and wetlands next to rivers (World Resources Institute, 1996). Cities located in close proximity to volcanoes

330

Urban geomorphology in the tropics

tend to climb their lower slopes, as in the case of Mexico City or Bandung (Java). All such areal spreads come with associated geomorphic problems (Gupta, 1984). The alteration of the landforms and processes due to urbanisation, as described earlier, is common in the tropics. The data, however, tend to come from large cities, as these are the ones usually studied. We can group the problems of landforms and processes due to urbanisation into three hazard classes: 1. Standard problems of alterations of slopes and channels due to urbanisation, as mentioned in section 18.1. 2. The problems are heightened in certain cases due to the location of the cities in naturally hazardous areas, such as a tectonically active zone or where tropical cyclones are common. 3. Hazards created or accentuated by accelerated resource utilisation and city metabolism, e.g. excessive use of groundwater. A city frequently demonstrates multiple hazards. We review three major cities to illustrate such hazards, and evaluate whether these hazards could be at least partially ameliorated using knowledge of local geomorphology. Details of these three cities are in Gupta and Ahmad (1999). We will follow up these case studies with a discussion on the application of geomorphology for improving the quality of urban life. The examples that follow indicate different types of problem in connection with urban geomorphology and the practicality of solving such problems. The case studies include: • Singapore, where the geomorphological problems are small-scale flooding and slope failures, that are almost entirely anthropogenic, and which have been reasonably managed. • Kingston (Jamaica), sited in an extremely hazardous location which exacerbates slope failures and flooding, and where the management of geomorphological hazards is extremely difficult. • Bangkok, where an existing local problem has been magnified to a nearly unmanageable level by city development, which led to the unrestricted use of groundwater.

18.3╇ Three examples of geomorphic hazards and their amelioration 18.3.1╇ Urban geomorphology of Singapore Singapore was founded in 1819 by Sir Stamford Raffles, on authorisation from Lord Hastings, the Governor-General of India, as a commercial station in the Johore-Riau area for serving British mercantile interest. The location is extremely favourable. Tectonically, it is a stable area, away from any plate boundary disturbance. The relief is low, although short steep slopes do occur on the island. The average annual rainfall ranges between 1650 and 2250 mm across Singapore. Usually the tidal range is less than 2 m, and the city is

331



Fig. 18.3

Geomorphic hazards and their amelioration

Diagrammatic sketch of land use change and its effect in a granitic hill area in the humid tropics, central Singapore. From Gupta and Ahmad, 1999. With permission from Elsevier protected by a small group of islands to the south, the large Indonesian island of Sumatra to the southwest and the southern tip of the Malay Peninsula towards the north and northeast. It is safe from tropical cyclones, being almost on the equator. Most rainstorms are small (less than 10 km across), intense and of short-duration (usually not lasting for more than an hour). According to the Intensity–Duration–Frequency study of the Singapore Meteorological Service, about 90 mm of rain may fall in a 30-minute, 2-year rainstorm. The runoff from such storms has the capacity to erode land that is unprotected by vegetation. The city started on a coastal plain and the slopes of a nearby low hill (the present Canning Hill) and gradually spread inland along small river valleys and over the rolling low hills of Triassic sedimentary and granitic rocks. The natural regional landform was that of small, steep, forested hills in sharp contact with valley flats, which comprise the floodplain of a river channel in the middle. These floodplains used to act as wetlands, attenuating the downstream passage of rainstorm-created floods. An impervious cover spread over the hillslopes and arterial valleybottom roads were built along the floodplains (Fig. 18.3). There was rapid expansion after the 1960s, when planned urban development was vigorously pursued by the government. As Singapore expanded, arterial roads followed valley flats with residential expansion climbing the hills on either side. The natural propensity for flooding the valleyflats was thus heightened by urbanisation. By the 1980s, about half the island was built-up area, with the percentage rising steeply towards the centre of the city.



332

Urban geomorphology in the tropics

Fig. 18.4

Bukit Timah Basin, Singapore illustrating intensity of urban land use that necessitates a large lined canal and drainage diversions (not shown in the photograph). The cars indicate the scale of the canal which at this point has only about 6 km2 of catchment. Photograph: A. Gupta Local newspapers reported 100 floods in less than 30 years, although most of these occurred before 1985. Floods in Singapore tend to relate to urbanisation:€they are localised and last for a short time (Gupta, 1992). The urbanised nature of Singapore floods has been studied in the 6.4 km2 Upper Bukit Timah Basin. Here the densely built-up area increased from less than 10 per cent of the basin area to over 45 per cent in 36 years (1950–1986). By 1960, the stream had been engineered into a straightened earth canal. Between 1967 and 1972, this had to be replaced by a larger trapezoidal lined canal with a flood diversion arrangement to the neighbouring basin to the west, as there were six to seven floods each year that affected the arterial roads that ran on the former swampy floodplain. The calculated size of the flood with 5-year recurrence interval increased from 73.68 m3s−1 in 1950 to 91.65 in 1986 m3s−1 (Chuah, 1987). Progressive housing development on the steep slopes draining into the canal required further modifications, as floods were becoming common after rainstorms. A second diversion was built to the neighbouring basin to the east and part of the 16 m trapezoidal canal was redesigned to a U-shaped one, 20–26 m wide. The need to ameliorate the disturbance created by frequent urban floods transferred natural drainage to lined channels. This was an ongoing process, but rapid urbanisation increased the total lined channel length from 320 km in 1979 to 650 km in 1984, more than doubling in five years at a cost of $200 million for channel improvement and extension. Singapore now has about a kilometre of lined channel for each square kilometre of area. Flood alleviation in Singapore depends on the capacity of drains keeping up with urbanisation of the upper valleys, and urbanisation is continuing (Fig. 18.4). Given their catchment areas, the new lined channels of Singapore are very large, a reflection on the anthropogenic flood propensity in the humid tropics. Finding space for this enlargement is difficult in the older parts of cities and requires some ingenuity. For

333

Geomorphic hazards and their amelioration

example, this has been achieved in parts of Singapore by hiding the drainage canals under decorative pedestrian malls and walkways. In the peripheral areas of new townships, large lined canals are built from the very outset, and are designed to take runoff when the construction of the new township is complete. Huge concrete canals, however, are not very attractive, but the physical alteration of the landscape also provides the opportunity to flank them with trees, paths and grassy stretches. Such a practice provides recreational and nature corridors through the city and creates surrogate rivers. Amelioration of urban floods ideally requires good practices of geomorphology, engineering and ecology. The other principal geomorphological hazard in Singapore is the failure of engineered slopes, including both rock and fill slopes. Slopes tend to fail after regrading, quarrying, road building and residential development. Such alterations activate weathering processes and also change the groundwater regime. The circulation and accumulation of subsurface water, particularly following a rainstorm, is controlled by relict structures in the weathered material (Grade V) including iron-crusted layers. Even weathered granite, which occurs over about a third of Singapore, tends to show stress relief sheeting structures or openings of pre-existing discontinuities when excavated. The slopes then fail along definite surfaces, depending on both antecedent and trigger rainfalls (Pitts, 1992). Pitts (1992) discussed slope failures in Singapore. The slides tend to be large in number, small in size and repetitive in occurrence. The positioning of many of these slides is due to the location of bedding, joints and cleavages on backscars, sliding surfaces and release surfaces. A number of past slides are primarily due to engineering activities, such as aligning new roads parallel to the strike of sedimentary rocks, oversteepening of cut slopes and failure to provide adequate subsurface drainage. Both floods and slope failures are mainly dealt with in Singapore by post-event engineering construction of lined canals, flood diversions, porous pavements and slope protection devices. The city is supported by efficient technical and managerial capacities, and financial resources are available to tackle such hazards. The scale of the hazards is also limited, and they occur only when the local geology and geomorphology are ignored. Singapore is an example where the hazards are small-scale, the resources are available and it is possible to ameliorate the problems to a large extent. The next example is different. Kingston (Jamaica) is prone to hazards of different types and large magnitudes and its resources are limited.

18.3.2╇ Kingston:€the hazard-prone urban landscape Jamaica is located within a 200 km wide left-lateral strike-slip deformation zone of Neogene age, marking the boundary between the North American and Caribbean Plates. Neotectonics and bedrock control the geomorphology of this island. This is also an area of high rainfall from tropical storms, which may reach hurricane force. The city of Kingston is located in such an environment, where there is also a high population and widespread poverty. It is one of the most environmentally hazardous cities in the world. Kingston is a 200 km2 mosaic of coastal plains, gravelly palaeofans and steep hills. All areas are hazardous. The rocks in the hills are highly faulted and deeply weathered. The fault zone continues to be active. The epicentre of the 13 January 1993 earthquake

334



Fig. 18.5

Urban geomorphology in the tropics

Sketch map of the city of Kingston with steep hills surrounding the Liguanea Fan was north of Jacks Hill, a northern suburb of Kingston. The maximum thickness of the Liguanea gravel fan, over which most of Kingston is situated, is at least 259 m. The top 8 m of the fan appears to be an old debris flow deposit associated with palaeodrainage of the Hope River (Fig. 18.5). Landslides and liquefaction-related ground failures occur in Kingston during each major earthquake. Shepherd and Aspinall (1980) determined the recurrence interval of earthquakes of MMI VII, VIII, IX and X as 38, 97, 137 and 237 years, respectively. Slopes also fail frequently in tropical storms (Ahmad, 1994). About 200–300 mm of rain in 24 hours initiates shallow slides on more than 25° slopes. Such slopes cover about 80 per cent of the hilly areas of Kingston and this amount of rain is expected to arrive once in 5 years. Flooding is also associated with transfer of a very high amount of sediment with a high coarse fraction. The damage from hurricanes is greater but less frequent. The most common types of slope failure are shallow slides and debris flows on steep, weathering-limited hillslopes. The natural vulnerability of Kingston is further increased by three anthropogenic acts: • deforestation on slopes • occupation of hazard-prone lowlands and steep hillslopes by the less affluent part of the society; they are joined on the upper slopes by the rich community occupying plots with commanding views • reoccupation of slopes partially stabilised after a failure.

335

Geomorphic hazards and their amelioration

Such removal of vegetation and soil erosion raises the probability of numerous shallow slides and debris flows in heavy rain (Maharaj, 1993), resulting in the destruction of property and infrastructure. Furthermore, urbanisation of the Liguanea Fan has led to increased overland flow and flash floods. Geomorphological problems are difficult to resolve at this scale. The shortage of funds and resources and the difficulty in moving people from particularly hazardous areas do not help. The Office of Disaster Preparedness and Emergency Management of Jamaica is based in Kingston. Both the government and the University of the West Indies are involved in active research on geomorphological hazards and a database is growing. The implementation of the findings, however, is far more difficult than raising awareness. Kingston is not the only city with this kind of problem. A densely populated city in a particularly hazardous location is quite common.

18.3.3╇ Bangkok:€flooding and subsidence Bangkok is located in the flat delta of the Chao Phraya, about 25 km inland. The ground elevation is extremely low, varying between 0.5 and 2 m above sea level. The average annual rainfall is 1500 mm. Drainage has always been a problem in the low swampy area, especially during the wet southwestern monsoon that operates between late May and early October. The old city was located next to the river, partly on its levee. Over time the city expanded east and south into the flat swampy lands of the delta. The expansion has been phenomenal in recent years. The 1950 population of Bangkok was less than 1.5 million. The current population is over 10 million. The drainage is directed to the Chao Phraya via a series of canals, originally constructed for both drainage and transportation. Heavy rain, a high river level, the flatness of the land and high tides combine to create an environment that is difficult to drain during the wet monsoon. Widespread inundation occurs regularly in Bangkok. Inundation during the rainy season is a natural phenomenon, but Bangkok’s flood problem has been disastrously exacerbated by anthropogenic extraction of its groundwater. The urban expansion of Bangkok resulted in a rapid increase in demand for water for both domestic consumption and industrial use. In the east and southeast of the city, this demand was met by pumping water out of sand beds that function as aquifers in the Quaternary thick marine clay underneath the city. The quality of water was acceptable and as the normal water rates were not levied, the use of groundwater was a favourable economic proposition for new housing and industries such as breweries and paper manufactures. By 1985, about 1.3 million m3 of groundwater was being pumped out daily. The piezometric level in the wells, which was originally at ground level, as expected in a deltaic plain, fell to about −9 m by 1959 and locally to nearly −50 m by 1983. Saline water started to leak out of the marine clays in between the sandy aquifer beds in the groundwater (ESCAP Secretariat, 1988). Withdrawal of such a large amount of groundwater from sandy subsurface aquifers led to subsidence, in places up to nearly 10 cm each year. The cumulative effect of this subsidence created a bowl-shaped depression over eastern Bangkok, which has been estimated to be 4550 km2 in area (Nutalaya et al., 1996). This is a shallow depression, but easily flooded

336



Fig. 18.6

Urban geomorphology in the tropics

Cumulative subsidence in the Bangkok area. From Noppadol and Nutalaya, 2005. By permission of Oxford University Press and difficult to drain. Several days of waterlogging are now common in Bangkok. In the exceptionally wet year of 1983, when 1029 mm of rain fell in August and September, water covered an extensive part of Bangkok for months. The subsidence progressively damages the drainage systems, as drains and canals tend to sink with the ground. Subsidence also structurally damages roads and buildings in Bangkok (Noppadol and Nutalaya, 2005). Figure 18.6 demonstrates the level of cumulative subsidence in Bangkok. Subsidence has reached the coast at the head of the Gulf of Thailand south of Bangkok, resulting in the sinking of a part of the seashore, marine inundation and ecological destruction, especially mangroves. Further lowering of the ground surface at Bangkok, which is only a short distance inland, and simultaneous global-warming driven sea-level rise may increase the level of hazard disastrously. The scale of the problem has accelerated research on the local Quaternary geology that forms the subsurface material in Bangkok, and the surface of the city has been carefully surveyed. Nature, causes and rates of Bangkok’s subsidence are now well known and the solution obviously lies in stopping the subsidence and, if possible, raising the level of the land. This, however, is very difficult. Groundwater pumping is now controlled and water is no longer free. Perhaps the best option for flood control is a combination of careful land

337

General nature of urban geomorphological problems

zoning, control of the growth of Bangkok and its suburbs, and increased restriction on groundwater withdrawal. Subsidence, however, cannot be reversed once the marine clays below Bangkok have been dewatered.

18.4╇ The general nature of urban geomorphological problems These examples illustrate the degradation of slopes and streams as a result of urbanisation in the tropics. If the scale of degradation is limited, the problems can be resolved by careful planning and management. However, in certain cases, the hazardous nature of the city’s location may make it almost impossible. A number of cities are located in such hazardous areas, with the hazards arising from tectonic instability, volcanic eruptions or the expected passage of large storms at short intervals. Certain cities suffer from more than one of these factors. Kingston is vulnerable to both earthquakes and hurricanes. Mexico City is very close to volcanoes and parts of it suffer from subsidence due to groundwater withdrawal. The nature of local geology may cause problems, if there are vulnerable sites such as buried river channels, old landslide deposits or unmapped karst landforms in the subsurface. All these cause problems in construction and later the collapse of buildings, at times with loss of life as has happened in Kuala Lumpur (Douglas, 2005a). Problems also arise from excessive resource utilisation, such as the withdrawal of groundwater stored in the subsurface. This, as has happened in Bangkok, results in subsidence followed by increased inundation of the city and structural damage. This happens regularly in certain types of environment, such as cities in swampy basins, coastal plains and deltas. More than 70 per cent of the water required by Mexico City is drawn out of its aquifer. The groundwater is lowered about a metre annually and in the 50 years between 1935 and 1985, the surface of the central area of the city, the location with the highest demand of water has subsided 7 m (World Resources Institute, 1996). Subsidence and related hazards due to groundwater extraction are quite common and are experienced by the inhabitants of many cities. About 70 per cent of the population of metropolitan Jakarta depends on groundwater, most of which is extracted from tens of thousands of shallow wells and more than 3000 deep wells. The rapid increase of groundwater extraction after 1970 caused a significant drop in the level of groundwater, leading to saline water intrusion which particularly affected the older parts of the city near the coast. Saline water intrusion in the deeper aquifer system has advanced between 3 and 7€km inland. The usual problem of subsidence has accompanied this deterioration in water quality (Douglas, 2005b and references therein). Kolkata developed in the abandoned delta of the Ganga on levees, floodplains and swamps. As the city expanded, it extended into less desirable swampy ground, exacerbating the flood effect of the wet monsoon. The flooding is partly anthropogenic, caused by three factors:€extension of the impervious landscape, disruption of the natural drainage and the disappearance of the old swamps where floodwater used to accumulate. Furthermore, the increased demand for water is partially met by pumping out groundwater from layers of

338

Urban geomorphology in the tropics

sand and gravel, primarily from a depth of 60 and 180 m, an aquifer sandwiched between two clay beds. Large-scale groundwater extraction has caused the piezometric surface to drop by metres, leading to ground subsidence in south-central Kolkata (Biswas and Saha, 1985). The subsidence is still not at a critical stage but should not be allowed to continue. The problem of increasing subsidence associated with groundwater withdrawal has been reported so far from a limited number of cities, but as the generic causes (a Quaternary alluvial subsurface with sandy aquifer beds between impervious clay strata and large-scale demand for water) are common, subsidence-related problems may become far more widespread as urbanisation of the tropical countries continues.

18.5╇ Geomorphology and urban management The case studies illustrate the progressive degradation of the physical environment of cities with urbanisation and also the potential for the application of geomorphology for urban management. Such management requires (1) the acquisition of selective geomorphological data and (2) the ability to interface with specialists who plan and manage cities:€ urban planners, engineers and city managers. This is required globally, but the rapid urbanisation in tropics and subtropics has made it crucially important for developing countries. Geomorphological information required for managing tropical cities includes: • hard rock geology with special reference to neotectonics and rock structure • Quaternary geology, especially for river valleys, coastal plains and deltas • a descriptive account of landforms with special reference to steep slopes, breaks in slopes and lack of slope • depth of weathering and relict and precipitated structures in the regolith • ambient operation processes with special reference to drainage, landslides and subsidence • nature of rainfall, especially as caused by storms • prevalent water balance with details of storage and flux • distribution of vegetation • characteristics of resource use and extraction. The availability and quality of such information varies among cities. Douglas (2005a), for example, has reviewed the urban geomorphology of Kuala Lumpur to illustrate the type of information that is required. In a number of cities there need to be plans for special hazards, for example, Mexico City, Yogyakarta and Bandung, which are all located in the vicinity of volcanoes. In the tropics, certain cities have been well studied and even mapped for possible geomorphological hazards, but their number is small. This includes cities with possible extensive urban flooding such as Singapore and Kuala Lumpur, where the intensity and amount of tropical rainfall accentuate such floods. A number of cities on steeplands have also been observed and mapped. Kingston, Kathmandu and Hong Kong are good examples. Cities

339



Fig. 18.7

Geomorphology and urban management

Hazardous location of Arequipa City and its surroundings close to the Misti stratovolcano. From Thouret et al., 1999. (www.borntraeger-cramer.de) located near convergent plate margins, where seismicity, slope failures and, in certain cases, pyroclastic flows and lahars are hazards have also been studied to some extent. Thouret and Lavigne (2005), for example, mapped the hazards and risks in the city of Yogyakarta from the nearby volcano of Merapi in central Java. Thouret et al. (1999) identified the hazards for the Peruvian city of Arequipa, located below the Misti stratovolcano (Fig. 18.7). Only 17 km separates the city centre from the crater of the volcano; the suburbs are even closer and extend up the slopes. Arequipa is vulnerable to ash fall, pyroclastic flows, lahars and debris avalanches. Cities tend to expand into hazardous zones, Nossin (2005) referred to the expansion of the city of Bandung northwards to Lembang, building luxury settlement estates which brings it very close to the volcano of Tangkuban Perahu. Urban geomorphology should be studied in far more cities in order to provide an extensive, in-depth and reliable database. Information from such databases could then be transmitted to urban planners and city engineers in a format that they can usefully use. A map is a useful format. Information supplied prior to the expansion of the city is useful for sensitive planning, so that areas with potential hazards which degrade a number of cities at present can be avoided. The information database is useful even for existing cities, as

Urban geomorphology in the tropics

340

(a)

(b)



Fig. 18.8

(a) Hazardous urban location on an alluvial fan at the mouth of a mountain stream in northern Venezuela and (b) destruction of a city in the debris flows and large floods of 1999. Photograph:€M. C. Larsen, US Geological Survey efficient management and correct engineering techniques could ameliorate some of the geomorphic hazards. The idea has been attempted in a few places. For example, Hong Kong has a wellestablished geotechnical organisation for building such a database. UN-ESCAP, in its role as a promoter of urban and environmental geology, published a series of volumes called the Atlas of Urban Geology from its headquarters in Bangkok. Detailed studies and maps of geology, landforms and hazards are therefore already in print for certain cities. Such an approach is required, especially for tropical cities located in hazardous environments. Given the location of a number of tropical cities, an escape from calamity may be impossible. The destruction of townships such as Carmen de Uria, located on an alluvial fan in northern Venezuela in the rainstorms of December 1999, is a relevant example. Landslides (mainly debris flows) and floods from nearly 1200 mm of rain in the first 16 days of the month inundated coastal communities, killed a number of people (estimated between 15€ 000 and 30 000), and severely damaged buildings and infrastructure (Fig. 18.8). The settlements in this area are commonly on alluvial fans, the only relatively flat terrain available between the coast and the steep mountains. The fans, as expected, are areas of high geomorphic activity (Larsen et al., 2001a, 2001b). Such vulnerability exists for a number of cities, as discussed by Chardon (1999) for Manizales, which is at a hazardous location in the Colombian Andes. Even if technical capabilities exist, unsuccessful

341

Questions

communication with urban managers and planners leaves such capabilities unutilised, as evidenced in the 1985 Nevado del Ruiz catastrophe (Voight, 1996).

Questions 1. What is urban geomorphology? Why is urban geomorphology becoming progressively more important in the tropics? 2. Are cities in the tropics optimally located from the viewpoint of a geomorphologist? 3. What conclusions can we reach from Table 18.2? 4. Floods in Singapore tend to be localised and short-term. Why? 5. It has been said that in Singapore flood alleviation is based on the capacity of drainage canals to keep up with urbanisation. Is this a correct summary statement? 6. Describe the problems of a tropical city of your choice which is located in a hazardprone environment. 7. Describe the common scenario for cities using large volumes of groundwater. 8. What kind of geomorphological information is required for managing tropical cities? How could such information be transmitted to city managers and planners?

19

The future with climate change

The future ain’t what it used to be. attributed to Yogi Berra, New York Yankees

19.1╇ Climate change and the future We did not experience any significant climate change over the last several thousand years. This allows us to extrapolate the present processes and rates over a length of time and assume that we can statistically stretch any time-based record, such as river discharge, if need be. We therefore are able to compute the size of 50-year floods and other useful measures. It seems that we can no longer make such an assumption (Milly et al., 2008). The rising global temperature is changing climatic components and forcing a rise in sea level. The nature and dimension of such changes are not likely to be uniform but vary across the planet (IPCC, 2007). The general pattern of environmental changes due to global warming has started to emerge, but we are still not in a position to predict with confidence the specific effect of global warming on small areas such as a 100 km2 drainage basin or the southern slope of an isolated plateau. A generalised list of global warming-related changes in the tropics is set out in the next section. Our knowledge of this area is still rather limited and it is preferable to discuss only the robust and large-scale changes. We still do not have enough information to model the future appearance and behaviour of small-scale individual geomorphological features like a hillslope due to climate change. It is also difficult to work out the changes for a large geomorphic system. For example, a large river like the Ganga or the Amazon would have to adjust to the shrinking of glaciers at the source, the changing pattern of rainfall and storminess over most of the basin, and a rising sea level at the delta. All such changes may occur near-simultaneously. Such an evaluation is a difficult task, but even a qualitative understanding of the possible state of a river or a hillslope or a delta face would be extremely useful. A basic knowledge of geomorphology is essential for this purpose, and as mentioned in Chapter 1, the geomorphologist may have the opportunity to fulfil the role of a prophet.

342

343

A robust prediction of the effects of climate change in the tropics

19.2╇ A robust prediction of the effects of climate change in the tropics We can anticipate certain geomorphological changes as a result of global warming. In this section we describe the changes in climate which appear to be near-certain. Subsequently, we examine the possible effect of such changes on landforms and geomorphic processes in the tropics. These are the modifications that the tropics would go through in the near future. We restrict the analysis to robust changes qualitatively and do not explore the finding of various models which are currently at the frontier of research and still being finalised. The effect of the changes in climatic characteristics discussed here are summarised from IPCC (2007).

19.2.1╇ Nature of climate change Concentrations of carbon dioxide, methane and nitrous oxides (collectively known as the greenhouse gases) in the global atmosphere have increased significantly since 1750 as a result of human activities. The concentration of CO2 has increased from 280 ppm during pre-industrial days to 379 ppm in 2005. Concentration of all greenhouse gases has increased, resulting in a warmer atmosphere. As a result a number of long-term changes in climate have been observed at continental, regional, and ocean-basin scales. These include changes in arctic temperature and ice, precipitation amounts, wind patterns, and aspects of extreme weather (drought, high precipitation, and tropical cyclones). Such changes are a departure from the normal trend. For example, atmospheric warmth recorded in the second half of the twentieth century is higher than what prevailed for the last 1300 years at least. The last time temperature rose so high for an extended period was about 125 000 years ago in an interglacial time. A number of global projections expect a warming of about 0.2°C per decade for the next twenty years, an estimate that may need to be revised upward if greenhouse gas emissions continue at or above the current rate. It is therefore very likely that high temperature and heavy precipitation events would become more frequent. Model-based projections of future rainfall are more reliable for the higher latitudes than the tropics, but certain characteristics may be noted. Summer monsoon rainfall over South and Southeast Asia is generally expected to increase along with increased rainfall over East Africa. Model projections indicate an increase in precipitation and variability in the Asian monsoon and the southern part of the West African monsoon. The pattern over the drier regions in Africa or the basin of the Amazon is not clear at this stage, although a decrease in summer precipitation has been projected for the Sahel. A similar decrease in summer precipitation is expected also for Mexico and Central America, and an increase for monsoon Australia. Meteorological observations already show that droughts have become more common over the tropics and subtropics since the 1970s. Changes in the radiative flux at the surface of the Earth alter heat and moisture budgets, and thus the hydrologic cycle. As the sea surface temperature (SST) of the tropical seas and oceans increases, tropical cyclones of the future would become stronger, with larger peak

344

The future with climate change

wind speeds and heavier precipitation. For example, post-1970 observations in the North Atlantic indicate an increase in intensity of tropical cyclonic activities. Since about 1970, the number and proportion of hurricanes that reached categories 4 and 5 have increased globally, although the total number of tropical cyclones has decreased. Such changes are expected to continue even if we manage to stabilise greenhouse gas emissions, as once started this type of global warming and rise in sea level may persist for centuries, given the timescale at which climatic processes and feedbacks operate.

19.2.2╇ Warming of oceans Since 1955, the oceans have warmed, accounting for about 80 per cent of the energy flux of the Earth’s climate, and such a warming is widespread over the upper 700 m of water in the oceans. The sea level is rising. For example, the global rate of sea level rise as measured for the 1993–2003 period by sensors carried by TOPEX/Poseidon satellite is 3.1 ± 0.7 mm yr−1. This figure is close to the estimated total of 2.8 ± 0.7 mm yr−1, which is reached by summing the thermal expansion of ocean water and the contribution from the melting of land ice by global warming. Satellite measurements, which tend to be precise, further indicate that the amount of sea-level change tends to vary from region to region. For example, the rise in sea level tends to be higher for the Southeast Asian and East Asian seas and, to a smaller extent, for the tropical Atlantic. Sea level may continue to rise for several decades. Greater melting of glacial ice from Greenland and Antarctica can raise the sea level much higher. About 125 000 years ago, during the last interglacial period, the mean global sea level was 4–6 m higher. The amount of ice on Earth is currently decreasing from most mountain glaciers and the Greenland Ice Sheet. Ice thinning and breaking up also occur in Antarctica.

19.2.3╇ Identifying the changes If we select only the robust changes we come up with the following list: • retreat of mountain glaciers • melting of snow and ice earlier in the year • changes in annual rainfall and regional variations • enhanced seasonality • regional increase in droughts • increase in strength of large storms (for example, hurricanes of category 4 and 5) • higher rainfall from extreme events • a rising sea level, not globally uniform but with regional variations. These changes that rise out of global warming and are expected to continue at least for decades will affect geomorphic processes in the tropics. We should expect different rates of erosion and sedimentation, and the modification of landforms. We may explore such alterations based on the basic concept of geomorphology and analogues arising out of natural changes during the Quaternary. Such Quaternary modifications have been explored for large rivers in general by Blum (2007) and specifically for the Ganga by Goodbred (2003).

345

Geomorphological adjustments in the tropics

19.3╇ Geomorphological adjustments in the tropics from climate and sea-level changes The changes listed in the previous section will affect landforms and geomorphic processes in the tropics. It is difficult to determine such effects precisely, due to the lack of reliable quantitative measurements, but we can work out the robust changes based on our general knowledge of fluvial geomorphology and analogues from the past, especially from the Quaternary. We can use the reconstruction of past river systems (Goodbred, 2003; Blum, 2007). We will construct the changes in a hypothetical large river basin, to illustrate possible future scenarios. Given the scale of our current knowledge, it is best to keep this analysis large-scale and qualitative. Major river systems integrate signals from large areas that stretch from the mountains of the upper basin through a long mixed bedrock–alluvial valley to sea-facing deltas. Climate change forces alterations in the discharge and sediment supply to the river, causing it to go through a series of morphological and sedimentological adjustments. As a result, more sediment is either stored in the valley or removed downstream. Large river basins are polyzonal, implying different types and rates of processes in operation in different tributary basins. Thus, different parts of the basin of a large river will adjust differently to climate change, complicating the picture. Rivers are conduits that transfer water and sediment from the land to the ocean. The nature of the river channels depends primarily on four variables:€ water, sediment, texture of sediment and channel slope (Chapter 7). One or all of these four prime variables may be altered by climate change, causing rivers to adjust to new conditions. The type of change would probably be similar for all rivers, but the degree of change will also be determined by the specific physical environment of the basin. This limits the level of confident prediction. Let us assume we are dealing with a hypothetical long river that originates in a tropical high mountain like the Himalaya and flows for thousands of kilometres in a mixed bedrock–alluvial channel through a wide valley to the sea via a large delta. This region is undergoing the appropriate robust climate changes as listed in section 19.2.3. We would expect glaciers to be on the retreat in the mountains and an earlier melting of ice and snow in the spring. Normally about 80–90 per cent of the sediment of a major river is derived from such mountains. The supply of the sediment would be enhanced by increased slope failures in the upper valleys due to the melting and retreat of glaciers. Episodic high rainfall events would be common if the mountains stretch across the tracks of major storms. Thus, early into climate change huge volumes of sediment may be produced, and stored temporarily in bedrock valleys in the mountains or as alluvial fans at the highland–lowland contact. Such sediment would be gradually transferred and stored along the mixed bedrock–alluvial valley and floodplain of the river. The sediment transfer is likely to be episodic, due to the enhanced seasonality in discharge and high rainfall from an increased number of large storms. The seasonality and large storms would also modify the annual discharge hydrograph of the river with a longer

346



Fig. 19.1

The future with climate change

Probable geomorphic changes in the basin of a hypothetical large river due to climate change period of low flow and increased peakedness. We may also expect overbank flooding and transfer and storage of sediment in both cross-valley and downstream directions. This type of behaviour would make the river wider, shallower and floodprone. The river would have a high rate of sediment transfer in pulses and a long-term storage of sediment in between, if accommodation space for it is available in the valley. The lower course of this hypothetical river and its delta may need to adjust to both the changing pattern of water and sediment arriving from upstream and a rising sea level. A rising sea level would allow more accommodation space for sediment; increase coastal erosion and saline intrusion; and allow more effective inland penetration by enhanced tropical cyclones. The expected morphological changes in the delta would include:€levee breaks and crevasse splays; sedimentation in flood basins and filling of wetlands; and avulsions and the abandonment of distributary channels in low flat deltas (Fig. 19.1). The river is a source-to-sink system and will need to adjust to all these modifications resulting from climate change. The future of this hypothetical river has been derived from material discussed earlier in the book. This is a simplified illustration of the application of geomorphology in an uncertain world. Similar scenarios could be constructed for other geomorphological units such as alluvial fans, coastal plains or steep slopes.

19.4╇ The noise effect of anthropogenic changes A large part of the Earth’s surface is no longer natural. Anthropogenic alterations have modified the prevalent geomorphic processes and, in certain cases, even landforms, as detailed in Chapters 17 and 18. Such alterations are extensive and their cumulative effect could be enormous. It has been estimated that reservoirs behind dams are capable of holding back 15 per cent of total annual river discharge (Nilsson et al.,

347

Tropical geomorphology in the near future

2005). In the world, there are 45 000 dams more than 15 m high or with a large reservoir (World Commission on Dams, 2000). The number of smaller reservoirs is much bigger. Vörösmarty et al. (2003) have computed with some reservation that annually 4–5 gigatonnes of sediment is deposited in the reservoirs behind such impoundments. This is between 20 and 33 per cent of global natural flux that would reach the sea in an unmodified world. In Chapter 17, we discussed increased erosion and the accelerated production of sediment due to various development projects. Even such an increased amount of sediment is arrested locally by dams and does not travel to the coast. Syvitski et al. (2005) concluded that due to such impoundments, sediment delivery to the seas has fallen globally by 10 per cent from the natural rate, and by much more for selected regions. Geomorphic systems that are relatively untouched anthropogenically are expected to respond significantly to climate change. They have done so during the Quaternary and are expected to do so in the near future (Blum, 2007). However, given the scale of anthropogenic modifications, the effect of anthropogenic changes is probably larger in certain systems than the impact expected from climate change. Under such circumstances, the signal of climate change could be masked, at least partially, by the noise of anthropogenic modifications. Prediction of the health and efficiency of geomorphic systems in the future is therefore going to be a complicated challenge. The tropics, and indeed the world, are likely to be maintained by natural geomorphic systems that are attempting to respond to climate change while also being affected by anthropogenic alterations. Any prediction under such circumstances is difficult.

19.5╇ Tropical geomorphology in the near future Geomorphology in the near future therefore will reflect three factors:€the normal pattern of landforms and processes, their anthropogenic modification, and the effect of climate change. All these factors carry their tropical components. We have discussed in detail the natural tropical environment for most of the book (Chapters 1–16). Anthropogenic modifications are currently extensive and significant in the tropics (Chapters 17–18). As most of these modifications are essential to development, their effects are increasing progressively in tropical countries. Certain aspects of climate change are also likely to be significant in the tropics. For example, rainfall from large tropical cyclones is expected to increase in the higher humid tropics. Certain areas of the arid tropics are expected to become drier. Rainfall zones may shift (section 19.3). The combined effect of these factors for a particular location or for a large geomorphic system will be difficult to determine. Given the state of our knowledge, numerical predictions such as the level of increase in rainfall or decrease in river discharge at a location are still not very reliable. We can, however, predict some of the robust changes on a qualitative basis, based on our knowledge of geomorphology. Even that could be a very important component of tropical geomorphology in the future.

348

The future with climate change

Questions 1. What aspects of climate change due to global warming are geomorphologically crucial? 2. How would a hypothetical long river, rising in a high tectonic mountain and flowing in a mixed alluvial-bedrock valley to the sea, adjust to climate change? 3. What is more important in tropical geomorphology:€climate change or anthropogenic alterations?

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Index

aa, 257 Aalto, R., 44, 148, 249, 349 Aberdares, 234, 291 abrasion, 224, 241, 277, 281 Aceh, 17, 184, 193, 361 Adriatic, 281 Africa, 3, 5, 6, 20, 22, 23, 28, 31, 33, 34, 37, 38, 44, 49, 53, 55, 81, 119, 121, 143, 192, 209–211, 223, 227, 229, 234, 251, 254, 266, 272, 288, 291, 292, 312, 319, 328, 343, 353, 356, 358, 363, 367, 370 Africa, East, see€Africa African Plate, 13, 23 Ahaggar, 23, 212 Ahagger, 23 Ahmad, R., 330 Alaknanada, 156 Alice Springs, 215, 217, 219 alluvial fans, 9, 22, 25, 91, 129, 137, 140, 213–216, 231, 232, 247, 249, 250, 253, 329, 340, 345, 346, 356 alluvial valleys, 4, 10, 26, 130, 247, 289, 298 Altiplano, 21 Amazon, v, 3, 6, 9, 20, 21, 27, 52, 56, 79, 104, 109, 129, 131, 138, 143–152, 158, 169, 173, 182, 200, 251, 252, 264, 292, 293, 342, 343, 354, 355, 359–364, 368 Amazon Cone, 151 anabranching river, 115, 118 anastomosing river, see€anabranching river Anatolian Plate, 19 Andes, 6, 7, 17, 20, 21, 27, 52, 55, 56, 79, 83, 85, 92, 96, 98, 109, 129, 144, 146–149, 151, 201, 210, 232, 234–236, 243, 247, 249–252, 287, 288, 290, 291, 340, 352, 358, 362 andesite, 17, 64, 71, 259 Annamite Mountains, 165, 166 Annapurna, 88, 235 annual maximum series, 112 Antarctica, 5, 6, 20, 143, 287, 344 antecedent rainfall, 87, 88, 100 anthropogenic modification, 5, 11, 29, 309, 347 antidunes, 133, 134 Antisana, 290 Appalachian Plateau, 281 Appalachians, 66 aquaculture, 207

374

Arabian Peninsula, 3 Arabian Plate, 23 Arabian Sea, 25, 28, 53, 267, 293 Araguaia, 90, 315, 317, 361 Arequipa, 339, 369 arête, 241 Argentina, 236, 250, 290 Aricha, 205 Arichaghat, 159 aridity index, 209 Armero, 264 Arunachal Pradesh, 159 Ascension, 16, 254 ash cloud, 258, 265 ash fall, 259 Asia, 6, 23–25, 28, 33, 43, 44, 55, 56, 166, 192, 197, 223, 282, 284, 287, 293, 295, 296, 309, 312, 319, 351, 352, 354, 355, 357–359, 363–365, 369–370 Assam, 80, 159, 161–163, 356 Assam Plains, 159, 162 Aswan Dam, 143 Atacama, 210, 211, 262 Atlantic, 6, 16, 22, 27, 28, 34, 39, 40, 42, 52, 144, 146, 148, 150, 154, 182, 184, 190, 200, 210, 236, 252, 254, 287, 289, 292, 344, 362 Atlas Mountains, 22 Atlas of Urban Geology, 340, 355 atolls, 190 Atrai, 205 Attock, 247 attrition, 180, 224 Auranga, 5, 123, 357 Australia, 3, 5, 6, 26, 33, 44, 53, 55, 56, 65, 74, 95, 121, 123, 124, 170, 190, 192, 209–211, 215, 217–223, 226, 227, 230, 249, 250, 291–297, 319, 343, 350, 351, 354, 359, 364, 366, 367, 369–372 avalanches, 246 backswamps, 136, 137, 151, 167, 198, 199 backwash, 177, 182 Bagnold, R.E., 103, 107, 223–225, 349 Baja California, 189 bajada, 215 Bale Mountains, 291 Baltoro, 235

375

Index

Bandung, 330, 338, 339 Bangkok, 330, 335–337, 340, 355, 362, 364 Bangladesh, 37, 137, 157, 159, 162, 163, 169, 203, 205, 206, 350 bankfull discharge, 112 bar, 116, 134, 136, 137, 142, 150, 158, 163, 168, 202, 206, 248, 268, 323, 367 Barbados, 22, 272, 292, 296, 350 Barind, 204, 205 Barisan Mountains, 17 barrier island, 184 barrier reefs, 190 basal sliding, 238–241 basalt, 6, 16, 17, 25, 62, 70, 71, 81, 86, 153, 179–181, 254–256, 267, 268, 270 base level, 109, 139, 140, 216, 293, 297 Baspa, 235 Bassac, 165 Batoka Gorge, 153 Batura, 235 bauxite, 70, 81 Bay of Bengal, 25, 28, 33, 155, 158, 163, 173, 182, 184, 192, 205, 247, 293, 299 beach, 17, 41, 167, 171, 172, 175, 177–179, 181–184, 186, 187, 189, 192–194, 199, 200, 202, 203, 223, 251, 309, 312, 323 beach drifting, 182 beach ridges, 184, 199, 203 beach rock, 183 bed load, 79, 80, 87, 104–107, 113, 116, 118, 127, 132, 143, 158, 203, 251, 268, 311, 324 bed material, 108, 116, 132–134, 158, 169, 316 Beira, 153 Belize Chamber, 280 Bengal Basin, 203, 207, 366 Bengal Deep Sea Fan, 205 Bengal Fan, 207, 300, 301 Beni, 44 bergschrund, 241 berm, 182, 183 Best, J. 162, Betwani Mountains, 19, 96 Bhagirathi, 156, 157 Biafo, 235 Blackhawk landslide, 89 blind valley, 277 block and ash flows, 261, 265 Blue Mountains, 91, 108, 138, 139, 197, 248 Blue Nile, see€Nile Blum, M. D., 344, 347 Bohol Island, 282 Bolivia, 44, 236, 250, 358 Borneo, 52, 234, 283, 301 bornhardt, 8 Bornhardt, 7 Bourke, M., 217–222, 351

Brahmaputra, 3, 8, 23–25, 27, 28, 80, 106, 109, 137, 143–145, 147, 150, 152, 155, 157, 159–163, 169, 182, 189, 197–199, 203–208, 247, 251, 252, 293, 299, 301, 350–352, 356, 360, 367, 371 braided rivers, 116, 137 Brandywine Creek, 111, 372 Brazil, 46, 52, 55, 56, 90, 96, 148, 150, 151, 200, 266, 314, 315, 317, 354, 355, 361, 362, 364 Brazilian Shield, 20, 22, 27, 56 Brice, J.C., 116, 163, 351 Bruijnzeel, L.A., 55, 247, 351, 354 Brunsden, D., 246, 252, 351 Bryan, K., 216, 351 Buchanan, F., 7, 78, 351 Büdel, J., 8 Bukit Timah Basin, 332, 352 Bukittingi, 260 Burbank, D.W., 24, 88, 249, 351, 355 Burdekin, 124, 372 Cahora Bassa, 152–154, 323 Calcite, 67, 69, 270 calcrete, 77 caliche, 69, 78 California, 19, 89, 351, 362 Cambodia, 122, 164–167, 268, 317 Canning Hill, 331 capacity, 107 Cape Fold Belt, 22 Cardamom Mountains, 165 Caribbean, 9, 13, 19, 22, 28, 40, 42, 55–56, 90, 96, 121, 122, 124, 183, 190, 192, 236, 257, 272, 282, 288, 333, 349, 355, 361 Carmen de Uria, 340 caverns, 270, 280 caves, 69, 75, 181, 270, 271, 274, 275, 277, 279–281, 283, 285, 296, 364 Cayambe, 290 Central America, 9, 22, 28, 52, 55–56, 210, 215, 249, 312, 319, 343, 355, 364 Central Cordillera, 21 CFB, see€flood basalts Chaco Plain, 250, 251 Chad, 254 chambers, 280, 281 Chandina, 205 Chang Jiang, 26, 28, 33, 111, 150, 197 channel geometry, 108, 109 channel network, 119 channel pattern, 109, 113, 115, 128, 136, 159, 167, 321 channel shape, 103, 109, 113, 124, 126 channel size, 109, 111, 113, 120 channel slope, 101, 108, 109, 345 channel-forming discharge, 112 channel-in-channel topography, 124, 126

376

Index

Chao Phraya, 26, 197, 335, 369 chelation, 66 chemical weathering, 61, 63, 66, 72, 79, 80, 89, 230, see€also€weathering Chézy, 102 Chézy equation, 102 Chihuahuan, 210 Chimborazo, 254, 290 China, 33, 42, 121, 140, 164–168, 179, 191, 197, 271, 275, 278, 279, 282, 295, 299, 323, 350, 358, 359, 373 Chittagong, 204 Chocolate Hills, 282 Cimarron River, 137, 366 cirque glacier, 237, 241 cirques, 241, 290, 291 clapotis, see€standing waves clay minerals, 65, 69–72, 138, 150 climate change, 10–12, 78, 147, 149, 171, 206, 207, 288, 292, 293, 298, 299, 303, 323, 342, 343, 345–348, 351, 356, 362, 365, 369, 370 climate proxies, 131, 292, 297 climo-morphogenetic landforms, 4 coarse wood debris, 248 coast, 9, 17, 25, 28, 29, 33, 34, 37, 41, 43, 44, 52, 98, 141, 150, 153, 170–173, 175–177, 179–184, 187, 189–194, 196, 200, 205, 206, 209, 210, 226, 232, 233, 251, 257, 263, 281, 285, 308, 309, 312, 319, 324, 328, 336–337, 340, 347, 361, 370 coastal cliff, 170, 179, 180 Coastal Cordillera, 21 coastal dunes, 186 cockpit karst, 278 Cocos, 6, 20 Coleman, J.M., 8, 137, 162, 163, 352 collision edge coasts, 172 Colombia, 52, 263, 359 Colorado, 322, 371 Comilla, 205 competence, 107 compressive flows, 239 conekarst, 278 Congo, 28, 53, 143, 144, 150, 152 Congo Air Boundary, 154 congruent dissolution, 67 continental shelf, 20, 150, 195, 203, 206, 297, 360 convergent boundary, 15 convergent plate margins, 9, 99, 254, 339 coral reef, 8, 10, 172, 182, 186, 187, 189–192, 194, 295, 309, 312, 354 Cordillera Apolobamba, 236 Cordillera Blanca, 236 Cordillera Huayhuash, 236 Cordillera Real, 236 corestones, 72, 73, 77, 95, 249 corrasion, 180 corrosion, 180, 274, 276, 277, 279, 281

Cotopaxi, 254, 290 craton, 4–6, 10, 13, 19, 25, 29, 75, 145, 158, 210, 211, 229, 295, 299 creep, 84, 85, 92, 93, 95, 96, 98, 100, 224, 238, 239 crescentic dune, 225 crevasse splays, 136, 137, 202, 203, 346 critical erosion velocity, 105 critical flow, 133, 134 crusts, v, 26, 74, 77, 80, 81, 224 Cuba, 22, 38, 41, 272, 279 Curah Lengkong, 263 Curray, 301, 354 currents, 32, 44, 170, 172, 177–180, 195, 198, 207, 209, 295–297, 368 Cvijíc, J., 270 CWD, see€coarse wood debris d’Arcy-Weisbach equation, 102 dam, 10, 29, 111, 128, 154, 165, 198, 247, 248, 293, 318–324, 346, 347, 360, 372, 373 dambos, 152 Dana, J.D., 7 Danum Valley, 55, 95, 312 Dardi, 111, 124 Darjeeling, 36, 247, 352, 368 Darwin, Charles, 7 Davis, W.M., 13, 353, 369 De Martonne, 8 Dead Sea, 211 debris avalanches, 90, 262, 339 debris fall, 85, 86 debris flow, 90, 91, 98, 216, 246, 334, 353 debris slide, 85, 87, 97 Deccan Plateau, 20, 254 Deccan Traps, 25 DeCelles, P.G., 250, 251, 361 deflation, 224 deforestation, 10, 29, 37, 58, 99, 147, 149, 246, 308, 309, 311–314, 317, 324, 334, 363 delta, 4–6, 28, 150, 153, 155, 157, 158, 163–167, 173, 189, 195–198, 200–208, 245, 251, 288, 297, 299–301, 309, 311, 317, 335, 337, 342, 345, 346, 354, 356, 360, 368, 369, 371 delta front, 198, 201–203, 205, 206 delta plain, 201–203, 206 Deodhar, C. A., 268 depression, 16, 24, 33, 36, 38, 130, 151, 152, 216, 237, 241, 245, 275–277, 283, 335 desert varnish, 216 Devprayag, 156 Dhaulagiri, 235 Dhubri, 159 Dibang, 159 Dietrich, W.E., 138, 353 Dihang, see€Brahamputra discharge, 24, 28, 29, 81, 101, 105–108, 110, 111, 113–117, 119, 120, 122–127, 134, 136, 138,

Index

377

142–145, 147–149, 154–157, 160–163, 166, 195–197, 214, 217, 218, 222, 232, 248–251, 264, 277, 280, 289, 292, 293, 312, 316–318, 320–322, 326, 342, 345–347, 356, 357, 363, 364, 368 dissolution, 66–69, 77, 180, 274, 275, 296 distributary channels, 153, 158, 195, 198, 202, 206, 346 divergent boundary, 15, 20 dolerites, 86 doline, 275–278, 280–283 dolomite, 69, 274 dominant discharge, see€channel-forming discharge Dominican Republic, 277 Dorr, J. A., 251 Douglas, I., 8, 51, 58, 247, 312, 314, 337, 338, 354, 356, 367 Drakensberg, 254, 266, 291 drips, 54 dune, 17, 26, 41, 133, 134, 171, 172, 183, 185–187, 192, 193, 199, 202, 203, 207, 210, 211, 223, 233, 251, 252, 261, 292, 293, 295, 325, 350, 360 Dunne, T., 45–48, 82, 84–87, 94, 131, 138, 147–151, 313, 314, 324, 325, 349, 354, 362 Duplessy, J.C., 293, 354 duricrust, 77, 80 earthflows, 91, 246 earthquake, 8, 16, 17, 19, 24, 29, 85, 96, 161, 163, 197, 232, 247, 251, 262, 328, 333, 334, 337, 366, 367 East African Rift System, 23 East China Sea, 33 easterly waves, 35, 39, 42 Eastern Cordillera, 21 Eastern Desert, 210 Eastern Ghats, 25 Eastern Syntaxis, 80, 159, 160, 162, 252 Ecuador, 52, 256 Egypt, 110, 189, 210, 319 Eilat, 211, 213, 214, 356 El Misti, 256 El Niño, 27, 34, 43, 44, 315, 349 El Niño Southern Oscillation, see€El Nino Elephant Hills, 165 Elm, 89 Emiliani, C., 288 Emmel, F.J., 301, 354 England, 281 Enriquillo-Plantain Garden Fault, 19 ENSO, see€El Nino Enzel, Y., 295, 354, 359 ergs, see€sand sea erodibility, 49, 51, 75, 110 erosion plot, 309 erosion rate, 19, 24, 162, 309 erosional threshold, 87 erosivity, 49

Ethiopia, 291, 292 Etosha, 210 Eurasia, 5, 6, 13, 19, 23, 165, 167, 364 Eurasian Plate, 5, 13, 23 Europe, 11, 112, 287 evapotranspiration, 3, 45–47, 50, 209 exfoliation, 64 experimental flood, 323 experimental plot, 310 extending flows, 239 fall velocity, 105, 106 falls, 23, 32, 33, 35, 38, 39, 43, 47, 49–51, 85, 87, 89, 92, 105, 111, 113, 154–157, 177, 180, 207, 212, 227, 258, 263, 264, 314, 316 fan-delta, 28, 29, 197, 371 Faniran, A., 65, 67 feldspar, 65, 67–69, 71, 73, 80, 227 fengcong, 282 fenglin, 282, 359 Ferrel Cell, 4 ferricrete, 77 fetch, 173 Finke, 217 firewood, 58, 311, 312 firn, 236, 241 firn line, 236 flocculation, 187, 198 flood, 19, 22, 31, 34, 35, 37, 38, 41, 42, 44, 47, 98, 105–107, 112, 113, 119, 121, 123–127, 132, 133, 137, 138, 142, 147, 150, 155, 157, 159, 160, 163, 166, 167, 189, 196, 197, 207, 213, 214, 216–220, 222, 223, 240, 246–248, 250, 267, 268, 277, 280, 283, 287, 293, 294, 322, 326, 327, 331–333, 335, 338, 340, 342, 350, 353, 356, 357, 359, 365, 372 flood basalt, v, 20, 71, 254, 266, 269 flood diversions, 333 floodout, 218, 223 floodplain, 112, 117, 125, 132, 136–142, 147, 149–153, 155, 158–160, 167, 219, 221–223, 247, 307, 320, 331, 332, 345, 353, 354, 359, 362, 366 flows, 19, 24, 27, 33, 44, 45, 47, 55, 63, 66, 69, 80, 90–92, 94, 95, 97, 98, 100, 105, 108, 117, 121, 123–128, 131, 133, 137, 138, 144, 145, 147, 148, 151–153, 157–159, 162–167, 177, 178, 195, 198, 202, 205, 206, 214, 216, 217, 222, 239–242, 245, 247, 249, 251, 256, 257, 259–268, 270, 272, 281, 299, 316, 317, 327, 334, 335, 339, 340, 345, 352, 353 fluid threshold, 224 Fly River, 138, 203, 353 Ford, D.C., 270, 271, 273–275, 277, 279, 280, 355 forest litter, 66, 68 form ratio, see€width-depth ratio Fournier, F., 8, 51, 355 Francis, P., 255, 256, 259–264, 266, 355 French Guiana, 200

Index

378

freshwater swamp, 52, 195 friction factor, 102 fringing reefs, 189 Froude Number, 133 Gabet, E., 88, 94, 96, 100, 249, 355 Galapagos, 256 Galunggung, 262 Gandak, 137, 157 Ganga, 3, 20, 25, 27, 28, 57, 99, 109, 125, 137, 138, 140, 143–145, 147, 150, 152, 155–163, 166, 169, 182, 189, 197–199, 203–208, 251, 252, 288, 293, 297–301, 337, 342, 344, 353, 356, 367, 371 Ganga Sagar, 158 Ganga–Brahmaputra Delta, 3, 161, 173, 197–199, 203, 204, 208 Gangamopteris, 6 Gangotri Glacier, 156 Garden of Eden, 283 Gardner, T.W., 9, 156 Gaumukh, 156 Gendol, 265, 266 Ghaghara, 157 Gibbs, R. J., 157 gibbsite, 65, 69, 71, 81 Gilbert, G.K., 202 Gillieson, D., 282, 284, 355 glacial mass balance, 237 glacial mass budget, see€glacial mass balance glacial outwash, 118, 245 glaciated trough, see€glaciated valley glaciated valley, 242, 243 glacier, 85, 235–244, 288–291, 299, 300, 342, 344, 345, 364, 371 glacier surge, 240 Glen Canyon Dam, 322, 371 Glen’s Flow Law, 238, 239 global warming, 342–344, 348 Globorotalia menardii, 288 Glossopteris, 6 Godavari, 25, 115, 140, 198 Gomati, 157 Gondwana, xi, 5, 6, 20, 23, 26, 99, 152, 154, 353, 363 Gondwanaland, see€Gondwana Good Luck Cave, see€Gua Nasib Bagus Goodbred, S.L., Jr., 145, 163, 197, 204, 206, 207, 293, 297, 299–301, 344, 345, 356, 360 Goswami, D.C., 162, 163, 246, Goudie, A., 26, 77, 211 graded profile, see€profile of equilibrium Grand Canyon, 321, 322, 371 granite, 8, 17, 27, 63, 66, 72, 73, 81, 179, 180, 212, 270, 333, 365 granodiorite, 71 Great Barrier Reef, 190, 296, 352, 364 Great Escarpment, 22, 210 Great Himalaya, 23, 24

Great Rangit, 247 Greater Antilles, 19, 22 greenhouse gases, 343 Greenland, 143, 344 Grodek, T., 283 groundwater, 45–47, 160, 186, 268, 271, 272, 280, 330, 333, 335, 337, 338, 341, 350 Gua Nasib Bagus, 283 Guanxi, 282 gufeng, 282 Guilin, 282 Guinea Current, 200 Guizhou, 278, 359 Gulf of Aden, 23, 292 Gulf of Mexico, 122 Gulf of Thailand, 336 Gunung Api, 283, 371 Gunung Sewu, 282 Gupta, A., 8, 10, 40, 58, 84, 110, 115, 117, 119, 121–125, 131, 144–146, 148, 149, 153, 155, 157, 158, 161, 164–167, 193, 199–201, 218, 220–222, 224, 247, 248, 266–268, 284, 296, 298, 302, 309, 313, 315–317, 323, 330, 332, 336, 349–351, 353–359, 361–365, 367, 369, 370, 373 Guyana Shield, 21, 27 Gyala Peri, 162 Ha Long Bay, 284, 285 Hadley Cell, 4, 32 Haiti, 19, 296, 354, 366 Hawaii, 70, 256, 257 Hell, 282, 355 Himachal Pradesh, 140, 235, 365 Himadri, see€Great Himalaya Himalaya, 5, 6, 11, 16, 20, 23–25, 28, 33, 52, 80, 83, 85, 88, 96, 98–100, 109, 144, 155–163, 165, 168, 197, 232, 234–236, 243, 246, 247, 249–253, 287, 288, 290, 291, 299, 345, 349, 351–353, 355, 366, 368, 370–371 Himalayan foreland, 156, 299 Himalayan Frontal Fault, 23 Hispaniola, 22 Hispar, 235 Hjulstrom, F., 105, 106 Holocene, 5, 57, 85, 147, 151, 167, 197, 200, 203, 205, 207, 285, 289, 290, 292–297, 301–303, 311, 350, 354, 356, 359, 361, 363, 364, 368, 369 Hong Kong, 8, 72, 96, 338, 340, 350, 360, 365 hook, see€recurved spit Hope, G., 291, 334, 358 horns, 187, 242, 243 Horton, R.E., 8, 21, 93, 119, 120, 250, 251, 358 Horton overland flow, 93 Huang He, 198 Huascaran, 247 Hughli, 157, 158 Humboldt, Alexander von, 7

Index

379

hurricane, 9, 22, 36, 39, 42, 58, 89, 90, 97, 98, 100, 124, 139, 194, 196, 197, 248, 333, 334, 337, 344 Hurricane Flora, 38, 41, 139 hydration, 68, 75 hydraulic action, 180 hydraulic geometry, 113–115, 128, 361 hydrolysis, 68 hyperconcentrated flow, 84, 216 Iceland, 16 ICOLD, see€International Commission on Large Dams ignimbrite, 260 IKONOS, 57, 112, 193, 208, 257, 265, 267, 285, 286, 318, 369 impact threshold, 224 imperviousness, 325 impoundments, 23, 143, 147, 308, 319, 323, 347, 371 incongruent dissolution, 67 India, 5, 6, 20, 23, 28, 33, 36–38, 50, 53, 70, 77, 80, 110, 111, 115, 124–126, 140, 142, 157, 159–161, 163, 169, 173, 184, 188, 203, 205, 210, 249, 254, 261, 266, 267, 272, 292, 295, 330, 351, 353, 354, 356, 358–360, 364, 366, 367, 369, 371 Indian Ocean, 5–7, 13, 17, 23–25, 28, 33, 34, 40, 42, 43, 56, 121, 123, 152, 153, 157, 158, 167, 181, 183, 204, 205, 211, 232, 250, 293, 299, 350, 354, 360, 363–365 Indian Peninsula, 5, 6, 23, 25, 33, 56, 123, 157, 158, 181 Indian Plate, 165 Indian–Australian Plate, 17 Indochina, 33, 34 Indo-Gangetic Plain, 23–25 Indonesia, 17, 26, 28, 33, 43, 57, 64, 70, 81, 92, 202, 234, 249, 257, 258, 261, 262, 266, 267, 271, 278, 291, 295, 313, 314, 349, 350, 361, 369, 370 Indus, 23–25, 108, 143, 144, 147, 247, 249 Indus–Tsangpo suture, 159 inselberg, 8, 22, 73, 216, 307 Inter Tropical Convergence Zone, 32, 147, 154, 232 internal deformation, 238, 239 International Commission on Large Dams, 319 IPCC, 342, 343 Ipoh, 282 Irrawaddy, 26, 28, 109, 116, 117, 126, 128, 143–145, 197, 207, 208 ITCZ, see€Inter Tropical Convergence Zone Itezhi-tezhi, 323 Jacks Hill, 334 Jakarta, 190, 328, 329, 337 Jamaica, 22, 36, 38, 42, 91, 97, 108, 115, 124, 138, 139, 192, 197, 248, 272, 278, 328, 330, 333, 335, 349, 357, 359, 361, 362, 366, 368, 370, 371 Jamuna, 137, 142, 159, 163, 206, 350 Japan, 295

Java, 17, 57, 81, 92, 95, 183, 190, 196, 232, 233, 247, 252, 254, 257, 258, 262, 263, 265–267, 271, 278, 282, 291, 292, 301, 313, 328, 339, 361, 369, 370 Java Trench, 17 Jeje, K. K., 22, 210 jet streams, 34 Johore, 330 Kafue, 323 Kailash, 159 Kailas-Mansarovar, 24 Kalahari, 22, 210, 226, 227 Kale, V.S., 22, 210 Kaliadem, 265, 266 Kalimantan, 202, 354 kame delta, 245 kames, 245 Kanchenjunga, 235 Kangning, 278, 359 kankar, 77 kaolinite, 65, 68, 69, 71, 80, 81 Kar, A., 24 Kariba, 153, 154, 323 Karnali, 24 karren, 275 karst, 9, 22, 270–279, 281–286, 337, 359, 371, 372 karst gorge, 277 karst, tropical, 8 Kashmir, 234, 235 Kathmandu, 338, 351 Kaveri, 25, 126 Kelantan, 184 Kelud, 263 Kerala, 184 Kershaw, A.P., 295, 297, 360, 372 Khartoum, 122, 350 Khudi Khola, 88 Kilauea, 256 Kilimanjaro, 52, 234, 254, 291, 364 King, L.C., 8, 364 Kingston, 328, 330, 333–335, 337, 338, 359, 361, 370 Kinta Valley, 282, 285 Kittipong, 314, 371 knickpoint, 111 Kolkata, 328, 329, 337 Kolohai, 235 Komar, P.D., 172, 173, 175, 184, 360, 368 Kong, 72, 166, 350 Köppen, W., 31 Köppen–Geiger classification, 31 Korat, 165, 166 Kosi, 137, 157, 250, 356, 367, 371 Kota Gadang, 260 Krakatau, 25, 140 Krishna, 25, 140 Krynine, P.D., 8, 360 Kuala Lumpur, 337, 338, 354

Index

380

Kuehl, S.A., 150, 197, 203–207, 293, 356, 360, 363 Kuta Tjuta, 34, 44 La Niña, 34, 44 Ladakh, 235 lag time, 212, 325–327 lagoon, 182, 184, 185, 190 lahar, 92, 98, 246, 247, 251, 259, 261–265, 269, 339, 363, 369 Lake Eyre, 217, 218 Lake Powell, 323 Lake Tanganyika, 23 Lake Titicaca, 290 Lancaster, N., 224, 225, 227–229, 360 land use, 10, 11, 29, 119, 120, 147, 207, 213, 214, 307–309, 311–313, 317, 324, 326, 331, 332, 352, 354, 356, 357 landslide, 8, 81, 96, 309, 337, 361, 362, 366 Langkawi, 285 Lao PDR, 112, 122, 164–168, 268, 271, 315, 317, 318 lapiés, 275 Larsen, M.C., 19, 89, 97, 98, 100, 340, 361, 367 Last Glacial Maximum, 150, 290 laterite, 7, 78, 299, 366 Latrubesse, E., 152, 315, 361 lava, 16, 25, 256, 257, 260, 261, 266, 268 Lavigne, F., 339, 369 Leier, A. L., 249, 361 Lembang, 339 Leopold, L.B., 8, 45–48, 82, 84–87, 94, 110, 112–114, 116, 117, 120, 137, 138, 214, 325–327, 354, 361, 372 Lesser Antilles, 22 Lesser Himalaya, 23, 24 levees, 91, 124, 136, 137, 151, 195, 198, 199, 202, 203, 337 LGM, see€Last Glacial Maximum Lichty R. W., 137, 366 Liew, S.C., 17, 119, 124, 131, 167, 193, 266–268, 358, 361, 369 Liguanea, 334, 335 limestone, 12, 69, 75, 81, 180–182, 190, 268, 270, 271, 273–276, 278–280, 282–286, 296, 351, 364 limestone pavements, 275 Limpopo, 28, 154 linear dune, 16, 20, 37, 82, 130, 136, 172, 176, 184, 187, 188, 201, 225, 228, 245, 275 lined canal, 333 llanos, 55 logging, 51, 311–313, 354 Lohit, 159 Lombok, 64, 268, 291 long profile, see€longitudinal profile longitudinal profile, 109–111, 242, 320, 324 longshore currents, 177 longshore drifts, 184 lower flow regime, 134

Lumajang, 263 Lunkaransar, 295 Luquillo Mountains, 248 Luzon, 43, 92, 259, 261 maars, 255 MacDonnell Ranges, 217 Mackin, J.H., 109, 362 Madagascar, 6, 23, 34, 36, 43, 53, 121, 272, 356 Madeira, 44, 149, 150 Madhupur, 204–206 Magdalena, 27, 143, 144 magnitude-frequency concept, 120, 126–128 Mahakam Delta, 202 Main Central Thrust, 23 Main Range, 20, 232 Majuli Island, 162 Makalu, 235 Makgadikgadi, 153, 210 Malacca Strait, 302, 354 Malay Peninsula, 20, 37, 182, 184, 232, 250, 285, 301, 331 Malaysia, 37, 55, 57, 95, 179, 181, 279, 280, 282–284, 288, 302, 312–314, 349, 350, 354, 358, 360, 367 Mameyes Basin, 19 Mameyes River, 97 Mana Pools, 153 Manasalu, 235 mangroves, 10, 52, 172, 186–190, 194, 195, 198, 199, 202, 203, 206–208, 285, 295, 309, 312, 336 Manila, 261, 329 Manizales, 340, 352 Manning’s equation, 102, 103 Manning’s n, 102, 103, 134 Marajó Rift, 147 marginal sea coasts, 171 marine oxygen isotope series, 289 Marmore, 44 Marsh, G. P., 307, 362 Martinique, 22 Maryland, 137, 357 mass movement, 83, 85, 246, 307 mass wasting, see€mass movement Mauna Kea, 256 Mauna Loa, 256 Mauritius, 62, 254 Mavis Bank, 97, 138 Meade, R.H., 27, 28, 109, 129, 138, 143, 146, 148–150, 154, 166, 182, 232, 251, 252, 354, 359, 362, 363, 371 mechanical weathering, see€physical weathering Mediterranean, 51, 66, 212 Medlicott, H.B., 6 megacities, 328 megafans, 157, 232, 249–251, 253, 299, 300, 358, 361

Index

381

megalakes, 293 Meghna, 156, 157, 205, 299, 350 Megi, 37 Mekong, 5, 26–28, 109, 111, 112, 119, 122, 124, 127, 128, 131, 142–145, 152, 164–169, 197, 199, 200, 202, 268, 315–319, 323, 357, 358, 360, 362, 363, 368 Mekong River Commission, 122, 166, 362 Merapi, 92, 183, 232, 233, 247, 252, 261, 262, 264, 265, 267, 269, 339, 369 Merritt, D., 248, 372 Mertes, L.A.K., 131, 138, 147–151, 354, 362 Mexico, 22, 34, 35, 39, 52, 55, 189, 210, 215, 312, 329, 330, 337, 338, 343 migration, 138, 151, 163, 296, 297, 299, 308, 311, 312 Miller, J. P., 120, 125, 128, 214 Milliman, J. D., 158 Milliman and Meade, 9, 27, 28 Milliman and Syvitski, 9, 12, 27, 28, 30, 146, 158, 232 MIS, see€marine oxygen isotope series Mississippi, 137, 145, 150, 198 mogotes, 279 Mohr and Van Baren, 7 Molengraff, G. A. F., 301, 363 monsoon, 5, 24, 25, 33–35, 37, 47, 48, 50, 55, 88, 100, 122, 123, 125, 145, 154, 156–160, 166, 167, 184, 196, 198, 203, 205, 207, 250–252, 262, 267, 282, 287, 293, 295, 299, 300, 315, 317, 335, 337, 343, 353, 356, 357, 359, 365, 367, 371 montane streams, 248, 253, 349 Montserrat, 268 Moore, A. E., 90, 244, 245 moraine, 90, 244, 245 moraine ridge, 244 Morakot, 38, 41 Mosi-a-tunya, see€Victoria Falls Mount Badda, 291 Mount Elgon, 234, 291 Mount Everest, 16, 23 Mount Jaya, 234, 291 Mount Kenya, 234, 291 Mount Kinabalu, 234, 288, 291 Mount Mayan, 211 Mount Sodom, 211 mountain-building, 19, 20 Mozambique, 152 mudcapes, 201 mudflows, 91, 92, 98, 246, 259, 262, 263 mudslide, 22, 85, 87 Mulu Hills, 283 Mulu National Park, 280, 283, 284, 351, 364 Murray Darling, 293 Myanmar, 33, 56, 116, 117, 121, 128, 160, 164, 208, 271

Nahal Yael, 211–214 Nairobi, 328 Nam Ngum, 166 Nam Theun, 166 Namche Barwa, 23, 80, 159, 162, 235, 252 Namib, 22, 64, 209, 210, 226, 229, 370 Nanga Parbat, 23, 24, 235, 236, 247, 249 Nankun, 235 Nargis, 207, 208 Narmada, 25, 109, 111, 115, 124, 125, 133, 138, 293, 358, 359 Nazca, 6, 13, 20, 147 nebkhas, 228 Negev, 212 Negro, 150, 152, 355 Neogene, 23, 287, 333 neotectonics, see€tectonism Nepal, 88, 100, 235, 246, 249, 252, 351, 355, 365 nested discharges, 113 Nevado del Ruiz, 263, 264, 341, 371 New Guinea, 8, 19, 28, 52, 53, 56, 96, 138, 249, 271, 274, 275, 283, 291, 295–297, 353, 359, 367, 372 New Zealand, 11, 254, 349 Ngonye Falls, 153 Niekerk, A.W., 119, 363 Niger, 28, 143, 144, 169, 202, 226 Niger Delta, 196 Nigeria, 50, 360 Nile, 23, 28, 110, 122, 143, 144, 169, 198, 252, 293, 302, 319, 350, 362, 373 nipa palm, 202 Noakhali, 205 non-equilibrium rivers, 217, 230 non-stratified drift, 244 North American Plate, 8, 13, 19, 75, 226, 333 North Anatolian Fault, 19 North Sunda River, 301 Northern Australia, see€Australia Nossin, J.J., 92, 259, 339, 349, 364 Nyamuragira, 254 Nyiragongo, 254 Óbidos, 150 Okavango Delta, 153, 154, 210, 362 Old Alluvium, 302, 358 organic acids, 66, 68, 271, 282, 283 Orinoco, 21, 27, 28, 52, 56, 79, 80, 104, 143–146, 148, 150, 182, 200, 201, 252, 359, 362, 368, 371 overland flow, 45, 55, 95, 150, 335, 367 Pacific, 7, 13, 17, 19, 21, 28, 33, 40, 41, 43, 44, 48, 52, 55, 56, 183, 190, 191, 256, 287, 292, 353, 366, 369 Pacific Plate, 13, 19 Padma, 143, 157 pahoehoe, 257, 258

Index

382

Pakistan, 46, 210, 249, 366 palaeokarst, 270 Palmer, A. N., 248, 287 Panama, 248, 287 Pangaea, 5, 6, 23 pans, 77, 81, 89, 152 Pantanal, 56 Papua New Guinea, 96, 291 parabolic dune, 228 Paraná, 143–145, 266 Parangtritis, 183, 252 Passarge, 7 Patton, P. C., 222 peat, 52, 203, 206, 293 pediment, 82, 215, 216, 223, 307 Peltier, L., 8 Peru, 44, 181, 256, 340, 369 Philippines, 26, 28, 36, 38, 42, 43, 92, 121, 249, 256, 257, 259, 261, 262, 278, 282, 315, 355, 363, 372 Phnom Penh, 165 phyllosilicates, 64 physical weathering, 61, 63, 64, see€also€weathering Pickup, G., 112, 123, 217, 218, 220–222, 301, 350, 351, 364 Pinatubo, 43, 92, 259, 262, 363 pinnacles, 27, 283, 296 Pitts, J., 88, 100, 333, 357, 358, 364 plagioclase feldspar, 68–71, 81 planar slide, see€translational slide plane beds, 134 plantation agriculture, 312 plate tectonics, 4, 10, 13, 19, 27, 28, 145, 147, 172, 254 platform, 19, 179, 180, 190, 351 Platte River, 123, 367 playa, 215, 216 Pleistocene, 5, 10, 24, 105, 147, 190, 191, 203, 205–207, 217, 234, 235, 243, 252, 285, 288–290, 292, 295, 296, 298, 303, 350, 354, 359 Plinian eruption, 258, 259 plucking, 241 plunging breaker, 177 pocket valley, 278 polje, 277, 283 pollen, 292, 352 polyzonal, 141, 142, 144, 298, 345 pool, 111, 116, 132, 134, 136, 167, 248 porous pavement, 333 Port Royal, 328 Potter, P.E., 143, 147, 232, 353, 365 prodelta, 202, 203 profile of equilibrium, 109 protalus rampart, 241 pseudokarsts, 270 Puerto Rico, 19, 22, 43, 55, 71, 81, 89, 95–97, 100, 248, 272, 277, 279, 349, 361, 363, 367 pyroclastic surge, 261

quartz, 65, 68, 71, 73, 75, 80, 138, 142, 150, 158, 223, 359 Quaternary, 10, 12, 23, 24, 131, 140, 151, 152, 163, 165, 167, 168, 186, 191, 207, 209, 229, 230, 283, 287–290, 293, 298, 299, 302, 303, 335, 336, 338, 344, 345, 347, 350, 352, 353, 356, 358, 359, 362, 364, 366, 368–373 Queensland, 53, 295, 352, 360, 369 Quelimane, 153 Raikot Glacier, 236 raindrop, 49, 54, 55, 93 rainfall intensity, 35, 49, 94, 97 Rajasthan, 77, 287, 295 Rajmahal, 157 ranching, 312 rapids, 131, 132, 134, 153, 165, 167, 322 Rapp, A., 37, 307, 308, 365, 369 Raub al Khali, 3 recurved spit, 184 Red River, see€Sông Hóng Red Sea, 16, 23, 189, 212, 292 regolith, 73, 76, 79, 80, 82, 84, 87, 88, 91, 95, 96, 99, 275, 276, 302, 309, 314, 326, 338 reptation, 223, 225, 228 Réunion, 254 Riau, 330 Richards, K. S., 183 ridges and runnels, 183 Rieger, W.A., 112, 123, 364 riffles, 116, 132, 133 rift valley, 3, 6, 16, 20, 144, 210, 247, 254 ring of fire, 17 Rio Chagres, 248 Rio Cubuy, 97 Rio Grande, 250 Rio Parapeti, 250 Rio Pilcomayo, 250 ripples, 133, 134, 136, 217, 225, 227–230, 283, 366 River Red Gum, 221 rock avalanche, 90, 260 rock desert, 210, 216 rock fall, 85 rock glide, see€rock slide rock rib, 131, 142, 167, 168 rock slide, 85, 87 rotational slide, 87, 89, 96 Rufiji, 28 Ruwenzori, 234 Ruxton and Berry, 8, 72, 76 Sabah, 55, 95, 234, 272, 288, 291, 312–314, 354, 363, 367 Sabie, 119, 363 sabkhas, see€salts flats SACZ, see€South Atlantic Convergence Zone Saffir–Simpson Damage Potential Scale, 41

383

Index

sag ponds, 20 Sagaing Fault, 116, 117 Sahara, 22, 34, 44, 210, 211, 223, 225, 226, 287, 291–293, 362 Sahel, 44, 343 Sahul Shelf, 295–297 saline water intrusion, 337 salt flats, 23, 172, 186, 187, 189, 207 saltation, 223–225, 228 Salween, 26, 28, 144, 168, 293 Samosir, 261 San, 19, 46, 85, 86, 166, 350, 354, 361 San Andreas Fault, 19 sand sea, 211, 223, 225, 226, 228, 229 sand sheet, 226, 228, 250 sandur, 245 Sâo Francisco, 198, 202 Sâo Paulo, 328, 329 Sapper, 7 Sarawak, 272, 275, 278, 280, 283, 284, 302, 314, 351, 358, 364, 372 Sarawak Chamber, 280, 283 satellite images, 39, 193, 264, 265, 309, 315, 317 Satluj, 24 saturation overland flow, 93 savanna, see€tropical grassland Schick, A.P., 211–214, 217, 356, 365 Schlesinger, W. H., 93 Schumm, S.A., 8, 50, 87, 113, 116, 117, 119, 130, 137, 361, 366 scoria, 255 scoria cones, 255 sea arch, 181 sea-floor cores, 288, 292 sea level, 5, 17, 43, 111, 139, 144, 145, 150, 168, 171, 173, 181, 184, 186, 189–191, 195, 197, 200, 204, 207, 210, 282, 283, 285, 288, 289, 293, 295–299, 301, 303, 335, 336, 342, 344, 346, 354 sea surface temperature, 39, 292, 297, 343, 350 seasonality, 52, 77, 113, 122, 123, 125, 127, 155, 157, 247, 273, 287, 344, 345 sediment, 3, 5, 6, 9–13, 17, 19–22, 24, 27–30, 34, 50, 51, 58, 79–82, 87, 89–92, 94–102, 105–113, 115–117, 119–122, 124–139, 141–152, 154, 156–159, 162, 163, 165–167, 169–173, 177–179, 182–184, 186, 189, 192, 194–198, 200–203, 205–208, 215–217, 219, 232–234, 236, 238, 239, 243–252, 260, 263, 264, 268, 270, 274, 277, 281, 288, 289, 291–295, 297–300, 307, 309–322, 324–328, 334, 345–347, 349, 350, 352–354, 356, 357, 359–363, 366–368, 371, 372 sediment plume, 10, 309 sediment splays, 198 sediment yield, 9, 10, 28, 29, 50, 51, 95, 247, 263, 309, 310, 313–315, 324, 354, 357, 360, 371, 372 seiches, 175 semiblind valley, 277

Senegal, 28, 198 shear stress, 103, 107, 121, 124, 126, 127, 166, 238, 239, 242, 264 sheetflood, 94 sheetwash, 94 shield, 19, 79 shifting agriculture, 312, 314, 315 shifting cultivation, see€shifting agriculture Shillong Plateau, 161 shore, 170–172, 176–179, 181–183, 187, 190, 192–194 shore platform, 180, 181 Shreve, R.L., 89, 119, 120, 366 Shroder, J.F., 236, 246, 253, 366 shutter ridge, 20 Siachen, 235 Sierra Madre Occidental, 210 Sierra Madre Oriental, 210 Sierra Nevada de Santa Marta, 236 Sierra Nevada del Cocuy, 236, 370 Sihang, see€Brahmaputra Sikkim, 235 silcrete, 77, 89 Simen, 291 Simonett, D.S., 8, 19, 96, 367 Simpang Formation, 302 Simpson Desert, 217 Simpson–Strzelecki Desert, 227 Sinai, 212, 217, 226, 365 Singapore, 37, 88, 100, 265, 302, 330–333, 338, 341, 352, 354, 357, 358, 363, 364, 369 Singapore Meteorological Service, 331 Singh, S. K., 217 Sinha, R., 145, 158, 367, 369 sinkhole, see€doline sinuosity Index, 116 Siwalik Hills, 23, 24 slackwater deposit, 293, 294 slides, 87, 90 slope failure, 19, 24, 31, 34, 37, 38, 42, 44, 76, 80, 84, 86, 88, 89, 96–100, 165, 180, 246, 249, 252, 253, 291, 297, 312, 317, 323, 330, 333, 334, 339, 345 slope profile, 82 slumps, see€rotational slide smectite, 65, 69, 71, 81 snow line, 236, 290 Socompa, 262 soil, 6–8, 11, 29, 34, 38, 44–47, 49, 50, 53–55, 58, 61, 62, 64, 66–68, 70, 71, 75–80, 88, 93–97, 122, 126, 188, 207, 212, 268, 272, 273, 275, 307, 309, 314, 317, 325, 335, 351, 358, 367, 371 soil creep, 92 soil pipe, 54, 95 Solimões, see€Amazon Solo, 196 solution load, 104, 157 solution pit, 275

384

Index

Son, 157 Sông Hóng, 26, 168, 197 Sonoran, 210 South Africa, 254, 266 South America, 5–7, 20, 21, 27, 43, 55, 143, 181, 182, 210, 232, 249, 272, 290, 312, 364, 367 South American Plate, 6, 13, 20 South Asia, 9, 23, 33, 44, 55, 121, 123, 192, 298 South Atlantic Convergence Zone, 147 South Atlantic Ocean, 6, 144 South China Sea, 164, 166, 168, 179, 191, 197, 295, 299, 350, 358 South Equatorial Current, 297 Southeast Asia, 5, 7, 10, 25, 28, 33, 38, 53, 55, 57, 73, 95, 121, 124, 143, 164, 186, 191, 192, 196, 262, 264, 271, 282, 284, 288, 291, 292, 295–299, 302, 309, 312–314, 317, 319, 336, 343, 350, 354–358, 363, 364, 369–370, 373 specific surface, 63, 81 spilling breakers, 177 spit, 182, 184 Spiti Valley, 140, 365 splash erosion, 49, 54, 93–95 SPOT, 117, 315, 317 spring, 20, 152, 268, 270 squall, 37 Srepok, 166 SST, see€sea surface temperature stack, 179, 181, 192 stalactite, 281 stalagmite, 281 Stallard, R. F., 281 standing wave, 175 Stanley, D.L., 197, 297, 368 star dune, 225, 228 Starkel, L., 247, 368 steephead, see€pocket valley stemflow, 54, 55, 95, 367 Stevaux, J., 315, 361 stone line, 78 storm sewer, 325, 326 storm surge, 41, 173 storm, 29, 34, 37–39, 42, 47, 50, 63, 95, 97, 121, 126, 171, 173, 183, 184, 190, 192, 202, 212, 217, 247–250, 263, 272, 327, 331, 337, 338, 344, 345 Strahler, A.H., 8, 46, 119, 120, 128, 368 straight channel, 116, 136, 162, 250 stratified drift, 244 stratovolcano, 232, 256, 264, 268 stream power, 103, 107, 121, 124, 126, 127, 149, 166, 169 streamflow, see€discharge Strombolian eruption, 258 sturtzstrom, 89 subcritical flow, 133

subduction, 6, 16–18, 26, 71, 147, 181, 247, 251, 256, 261 subsidence, 32, 275, 328, 335–338, 362, 364 sudd, 252 Sulawesi, 271, 291 Sumatra, 3, 17, 18, 37, 57, 184, 193, 258, 260, 271, 291, 301, 309, 331, 361 Sumbawa, 43, 261 Sunda Plate, 17 Sunda Shelf, 295, 296, 298, 299 Sundarbans, 206 supercritical flow, 133, 214 surging breaker, 177 Surinam, 200 suspended load, 79, 80, 104–107, 113, 121, 127, 128, 132, 158, 187, 203, 251, 252, 268, 311 swash, 177, 182, 183 Swatch of No Ground, 205, 206, 299, 301 Sweetings, M., 271, 281 Sylhet, 205 Sylhet Basin, 205 Syvitski, J.P.M., 9, 27, 196, 347, 363, 368, 371 Taal, 261 Taiwan, 36–38, 41, 43, 96, 122, 295 Tambora, 43, 261, 264 Tandon, S.K., 145, 369 Tangkuban Perahu, 339 Tanzania, 37, 307, 308, 365, 369 Tapi, 25, 115, 359 Tapponier, P., 26, 165, 369 Tasikmalaya, 262 tectonic belts, 4, 10 tectonic terraces, 139 tectonics, see€tectonism tectonism, 9, 19, 27, 48, 149, 356, 362 tephra, 258, 259, 262 terraces, 129, 139–141, 151, 158, 165, 218, 246, 247, 249, 251, 268, 290 Thailand, 112, 122, 131, 164–166, 271, 285, 295, 302, 309, 314, 355, 369, 371 Thar, 210, 226, 295, 354, 371 thermokarst, 270 Thomas, D.S., Thomas, M., Thornthwaite, C.E., 44–46, 48, 369 Thouret, J.C., 262, 265, 266, 339, 369 Three Gorges Dam, 111, 319 throughfall, 54, 55, 95 thunderstorm, 37, 49, 50 Tibesti, 23, 254 Tibet, 24, 159, 160, 162, 352, 364 Tibetan Plateau, 24, 25, 33, 159, 160, 162, 164, 166, 197, 287, 349, 355 tidal bore, 173 tidal mudflats, 184

Index

385

tides, 154, 172, 173, 179, 183, 195, 198, 201, 203, 205, 335 till, see€non-stratified drift tillite, 6 Timor Leste, 296 Tista, 163, 247, 251, 352 Toba, 261 Todd, 217–223 Tombolo, 184 Tonlé Sap, 165 Tooth, S., 218, 221, 370 TOPEX/Poseidon, 344 topple, 85, 86 Toricelli Mountains, 19, 96, 367 Torres Sánchez, A.J., 19, 97, 98, 100, 361 tower karst, 278, 285 trade winds, 32, 33, 40, 42, 70, 154, 279, 293 trailing-edge coast, 172 transform boundaries, 15 transform fault boundary, see€transform boundaries translational slides, 87 transport-limited slopes, 79 Trengganu, 184 trigger rainfall, 87, 88, 100 tropical cyclone, 28, 35, 38–43, 47, 48, 84, 88, 95, 96, 99, 105, 107, 121, 171, 173, 192, 194, 205, 207, 208, 247, 262, 315, 330, 331, 343, 346, 347 tropical disturbances, 34, 35, 37 Tropical Geomorphology Newsletter, 9 tropical grassland, 55, 57 tropical rain forest, 10, 50, 51, 53–55, 57, 58, 66, 93, 95, 100, 190, 311, 312, 354 tropical storm, 10, 22, 38, 40, 47, 98, 99, 121–123, 157, 166, 185, 186, 190, 197, 205, 207, 248, 333, 334 Tsangpo, 24, 159, 162 tsunamis, 17, 171, 190 tuff rings, 255 Tun Kul Cave, 280 Turkey, 19 typhoon, 40, 42, 196, see€also€hurricane Uluru, 74 United States, 7, 34, 50, 112, 123, 184, 202, 215, 259, 277, 320, 325, 353 upper flow regime, 134 urban geomorphology, 325, 330, 339 urbanisation, 10, 51, 57, 307–309, 325–328, 330–332, 335, 337, 338 uvalas, 277 valley glacier, 237, 240 velocity, 39, 40, 50, 54, 87, 89, 90, 93, 101–107, 109–111, 113, 114, 123, 124, 127, 129, 133, 134, 159, 173, 174, 176–178, 214, 224, 228, 239, 240, 242, 260, 261, 264, 272, 294, 316, 327, 364

Venezuela, 21, 75, 98, 340, 352, 359, 361, 371 ventifacts, 225 Verstappen, H. Th., 255, 264, 268, 299, 370 Victoria Falls, 152, 153, 155 Vientiane, 127, 128, 165, 167 Viet Nam, 121, 165, 271, 278, 279, 283–285, 302, 317 Viles, H.A., 66, 181, 282, 370 Virgin Islands, 38 Virunga Mountains, 254 Voight, B., 263, 264, 341, 371 volcanism, 10, 19, 23, 28, 30, 98, 254, 256, 268 volcano, 17, 22, 43, 183, 190, 255–259, 261, 262, 264, 265, 268, 339, 371 Vörösmarty, C.J., 319, 347, 368, 371 Vulcanian eruption, 258 Wadi El Arish, 217 Wadi Mikeimin, 217 Wadi Watir, 217 Wallace, Alfred Russel, 7 Walvis Bay, 266 Warne, A., 197, 201, 297, 368, 371 Wasson, R.J., 158, 295, 371 water balance, 44–46 water budget, see€water balance Watts Branch, 137 wave diffraction, 175, 176 wave refraction, 175, 183, 184 wave-cut platform, 180 waves, 41, 42, 89, 133, 134, 170–180, 183–185, 187, 188, 190, 194, 195, 198, 201, 202 weathering, 6, 8, 10, 11, 20, 31, 47, 61–73, 75–81, 87, 94, 98, 124, 138, 150, 157, 161, 170, 180, 223, 230, 267, 333, 334, 338, 352, 362, 367–370 weathering-limited slope, 82 Wells, N. A., 82 Western Cape, 291 Western Cordillera, 21 Western Ghats, 25, 53, 267 White Nile, see€Nile width-depth ratio, 112, 113, 117, 119, 124, 127 Williams, G.P., 8 Williams, M., 8, 291, 292, 295, 297, 320, 372 Williams, P.W., 270–275, 277, 279, 280, 287, 372, 373 Wilson, T.J., 18 Wohl, E.E., 124, 248, 358, 365, 372 Wolman, M.G., 111, 116, 117, 120, 125, 128, 137, 217, 320, 321, 325, 326, 361, 372 Woodroffe, C.D., 190, 191, 197, 202, 373 Xiaowan, 319 Yallahs, 97, 115, 124, 138, 139, 371

Index

386

Yamuna, 157 Yangtze, see€Chang Jiang yardangs, 225 yazoo streams, 137 Yogyakarta, 265, 338, 339

Yucatan Peninsula, 22 Zambezi, 28, 143, 144, 152–155, 169, 323, 363 Zhujiang, 28, 144 zibars, see€sand sheet



Fig 2.6

The Himalaya Mountains and Ganga Plains:€a tectonically active high relief environment. From NASA/MODIS



Fig 2.7

Northern Australia, an old arid landscape. From NASA/MODIS



Fig 4.8

Vegetation clearance for oil palm plantation in eastern Sumatra, Southeast Asia. Note lines of young oil palms towards the upper left corner, canals put in for drainage and burning of vegetation in the centre. IKONOS satellite image © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2006), reproduced with permission



Fig 7.8

The Mekong eroding into quartzitic rocks along the border of Thailand with Lao PDR, upstream of Pakse, Lao, PDR. Scale indicated by trees at the bottom left. IKONOS satellite image © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2003), reproduced with permission



Fig 9.10

MODIS/NASA image of the Brahmaputra Valley and the Ganga–Brahmaputra Delta



Fig 10.18

Tsunami erosion and recovery, northern Aceh coast. (a) location of the image used for the Aceh coast; (b) coast on January 2003; (c) December 2004 (tsunami); (d) February 2006; (e) January 2007; (f) April 2008. Note removal of the beaches, destruction of the wetland and erosion along swales. Within 13 months, beaches have returned, although the wetland in the centre is under water. Vegetation has returned to the swales. The effect of the tsunami is not generally perceptible even in the 2006 image. The next two images show the continuation of the rebuilding process. IKONOS satellite images © Centre for Remote Imaging, Sensing and Processing, National University of Singapore, reproduced with permission. From Liew et al., 2010

(a)



Fig 11.9

(b)

The effect of a tropical cyclone on a major delta. A. Part of the Irrawaddy Delta. B. Same area after tropical cyclone Nargis. IKONOS satellite image. © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2009), reproduced with permission



Fig 14.2

The crater of the stratovolcano of Raung in east Java is about 2 km across. The volcano is high (3332 m) and erupts regularly; it has already been active several times in the twenty-first century. Flow and explosive deposits and screes of volcanic material mark the crater rim and the caldera floor. A second crater with a deep volcanic vent rises asymmetrically from the floor of the caldera. IKONOS satellite image© Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2009), reproduced with permission



Fig 14.6

DEM of Merapi draped by the IKONOS image of 16 June 2006, showing the southern slope of the volcano and the Gendol River Basin. A pyroclastic flow is emerging from below the billowing ash cloud. The river channel is marked by deposition of earlier pyroclastic flows. From Thouret et al., 2010. By permission of Elsevier



Fig 14.7

Lower slopes of the DEM of the Gendol Valley showing details of a volcanic flow. Segment 1: upper section of the valley; Segment 2: widening of the valley and a fan of volcaniclastic material; Segment 3: a wider valley entrenched in a volcaniclastic fan. A: the moving pyroclastic flow from upstream; B: a ridge built by a previous flow; C: an area stripped by an ash-cloud surge; D: spilling and re-routing of flow at a bend; E: avulsion of re-routed flow to a neighbouring channel; F: lobes of flow reaching overbank which buried the village of Kaliadem (KA). From Thouret et€al., 2010. By permission of Elsevier



Fig 15.7

IKONOS image of coast near Phang Nga, southern Thailand. Limestone towers are partially inundated by a Holocene rise in sea level and surrounded by mangroves. Boats provide scale. Seawater frequently invades sinkholes in karst giving rise to a spectacular landscape. IKONOS satellite image © Centre for Remote Imaging, Sensing and Processing, National University of Singapore (2007), reproduced with permission