Encyclopedia of the World's Coastal Landforms

Encyclopedia of the World‘s Coastal Landforms

Eric C. F. Bird Editor

Encyclopedia of the World‘s Coastal Landforms Volume I

With 1379 Figures

123

Editor Eric C.F. Bird, M.Sc. (Lond. and Melb.) Ph.D. (ANU) Principal Fellow in Geomorphology University of Melbourne and Director, Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected]

ISBN 978-1-4020-8638-0 This publication is available also as: Electronic publication under ISBN 978-1-4020-8639-7 Print and electronic bundle under ISBN 978-1-4020-8640-3 DOI 10.1007/978-1-4020-8640-3 Springer Dordrecht Heidelberg London New York Library of Congress Control Number: 2009938922 © Springer Science+Business Media B.V. 2010 All rights reserved: Chapters 8.2, 8.15, 9.4, 10.1, 11.2, 11.3, 18.1 and 20.5 Every effort has been made to contact the copyright holders of the figures and tables which have been reproduced from other sources. Anyone who has not been properly credited is requested to contact the publishers, so that due acknowledgment may be made in subsequent editions. No part of this work may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, microfilming, recording or otherwise, without written permission from the Publisher, with the exception of any material supplied specifically for the purpose of being entered and executed on a computer system, for exclusive use by the purchaser of the work. Cover figures: Images 5084487, 5078544 and 5333716 from photos.com. © 2010 photos.com Printed on acid-free paper Springer is part of Springer Science+Business Media (www.springer.com)

Preface Encyclopedia of the World’s Coastal Landforms is an account of the various kinds of coastal morphology found on the world’s coastlines. It is presented in 224 chapters, based on national, state, or county divisions, and arranged in a sequence that proceeds counterclockwise around the continents. The sequence begins with Alaska, proceeds around North America to Arctic Canada, then Greenland and finally, Iceland. Another counterclockwise sequence begins with Norway and proceeds by way of Europe and the Mediterranean around Africa to India, Southeast Asia, China, Korea, and the Pacific and Arctic coasts of Russia. Interruptions are made at appropriate points to include the Great Lakes, the Caspian Sea, and islands such as Britain, Madagascar, and Japan. The sequence is completed with the Philippines, Indonesia, Papua New Guinea, Australia, New Zealand, the islands of the Pacific, Atlantic, and Indian Oceans, and finally Antarctica. This work was preceded by The World’s Coastline, edited by Eric C. F. Bird and Maurice L. Schwartz and published by Van Nostrand Reinhold, New York, in 1985. An internet version was published by Kluwer Academic Publishers in 2004. This 2010 edition has also been produced with the aid of 123 contributors. It includes 1,379 illustrations and cites numerous references. The focus is on coastal geomorphology, with accounts of features seen along the coastline, illustrating sites of special scientific interest. It provides access to the large amount of data that has been assembled, to ensure that it will not just disappear over the next few years. It is now much more difficult for scientists to make broad coastal surveys as they are required to specialise more, which usually means they work on small selected coastal sites. This book will allow the scientists to place their work in a global context and make comparisons with similar coasts in other parts of the world. Coastal geomorphology has several interest groups. Some focus on the shaping of coastal landforms in relation to geology, processes, variations in climate, and the relative levels of land and sea. Others concentrate on coastline changes measured over specified periods, with analyses of their causes. Others are interested in nearshore processes and responses, particularly beaches, where morphodynamic models have been most successful. Others look to the coastal area for evidence of Quaternary history, notably changes in land- and sea level and climatic variations. Others examine the sources and patterns of movement of coastal sediment, or the array of physical-, chemical-, and biological weathering processes in the coastal zone. Coastal managers want information on how landforms are changing, what processes are at work on them, how they will change in the future, and what effects will result from the development of various kinds. The chapters presented in this version of Encyclopedia of the World’s Coastal Landforms reflect the interests of their authors, but an attempt is made to serve the wide range of interests of coastal geomorphologists. The work was designed initially for scientists needing to identify locations and distributions of coastal features and to make comparisons between coastal landforms in various parts of the world. However, it will also be of value to people seeking general information on coastal landforms in the various countries around the world, including those planning travel and exploration of coastal regions. Armchair travellers can use it to make substantial journeys around the world’s coasts, enjoying a succession of visual images. Indeed, it is unique in presenting the wide variations in scenery and environment that occur on the world’s coastline. I would like to thank Lew Ward for processing the edited material and illustrations. In assembling the work, I had help from librarians in the University of Melbourne, notably David Jones, Map Curator in the Education Resource Centre. In Black Rock I had assistance from Dr. Juliet Bird, and from Myriam, Audrey, Ingrid, and Céline Vinot. Eric C. F. Bird Geostudies, Melbourne, Australia January 2010

Contents Volume I Editor .......................................................................................................................................................... xxiii List of Contributors ...................................................................................................................................... xxv 1.0

USA – Editorial Introduction ........................................................................................................................... 1 Eric Bird

1.1

Alaska ................................................................................................................................................................ 5 Jesse Walker · Molly McGraw

1.2

Washington ..................................................................................................................................................... 15 Maurice Schwartz · Thomas Terich

1.3

Oregon ............................................................................................................................................................ 33 Paul D. Komar

1.4

California ........................................................................................................................................................ 43 Eric Bird

1.5

Texas ............................................................................................................................................................... 53 Robert Morton

1.6

Louisiana ........................................................................................................................................................ 61 Ioannis Georgiou · Mark Kulp · Michael Miner · Duncan FitzGerald

1.7

Mississippi ...................................................................................................................................................... 69 Ervin Otvos

1.8

Alabama .......................................................................................................................................................... 77 Ervin Otvos

1.9

Florida ............................................................................................................................................................ 83 Charles Finkl

1.10

Georgia ........................................................................................................................................................... 89 Miles Hayes

1.11

South Carolina ............................................................................................................................................... 93 Miles Hayes

1.12

North Carolina ............................................................................................................................................... 99 Stanley R. Riggs

1.13

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey) ........................................ 107 Eric Bird

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Contents

1.14

New York and New England ......................................................................................................................... 113 Henry Bokuniewicz

1.15

Great Lakes (USA) ........................................................................................................................................ 121 Mary-Louise Byrne

1.16

Hawaii ........................................................................................................................................................... 125 Charles H. Fletcher · Eden J. Feirstein

2.0

Canada – Editorial Introduction ................................................................................................................. 133 Eric Bird

2.1

British Columbia .......................................................................................................................................... 135 John J. Clague

2.2

New Brunswick and Nova Scotia ................................................................................................................. 141 Eric Bird

2.3

Gulf of St. Lawrence ..................................................................................................................................... 155 Eric Bird

2.4

Saint-Pierre-et-Miquelon Islands ................................................................................................................ 163 Jean-Marie M. Dubois

2.5

Newfoundland .............................................................................................................................................. 167 Eric Bird

2.6

Northern Canada ......................................................................................................................................... 169 Wayne Pollard

2.7

Great Lakes (Canada) ................................................................................................................................... 177 Mary-Louise Byrne

3.0

Pacific Central America – Editorial Introduction ....................................................................................... 183 Eric Bird

3.1

Mexico, Pacific Coast ................................................................................................................................... 185 Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez

3.2

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama ...................... 187 Anja Scheffers · Tony Browne

3.3

Pacific Coast of Colombia ............................................................................................................................ 193 Iván Correa · Robert Morton

4.0

South America – Editorial Introduction ..................................................................................................... 199 Eric Bird

4.1

Ecuador ......................................................................................................................................................... 201 H. Ayon · W. Jara

Contents

4.2

Peru ............................................................................................................................................................... 207 Vidal Taype Ramos · Eric Bird

4.3

Chile .............................................................................................................................................................. 213 José Araya-Vergara

4.4

Argentina ...................................................................................................................................................... 219 Enrique Schnack · Jorge Pousa · Germán Bértola · Federico Isla

4.5

Uruguay ........................................................................................................................................................ 227 Eric Bird

4.6

Brazil ............................................................................................................................................................. 231 Dieter Muehe

4.7

French Guiana .............................................................................................................................................. 239 J. Turenne

4.8

Surinam ........................................................................................................................................................ 243 Norbert Psuty

4.9

Guyana .......................................................................................................................................................... 245 Eric Bird

4.10

Venezuela ...................................................................................................................................................... 249 L. Ellenberg

4.11

Caribbean Coast of Colombia ..................................................................................................................... 259 Iván Correa · Robert Morton

4.12

Caribbean Coasts, Panama to Belize ............................................................................................................ 265 Anja Scheffers · Tony Browne

4.13

Mexico, Caribbean Coast ............................................................................................................................. 269 Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez

5.0

Caribbean Islands – Editorial Introduction ................................................................................................ 271 Eric Bird

5.1

Cuba ............................................................................................................................................................. 273 Ridel Rodríguez

5.2

Jamaica ......................................................................................................................................................... 279 John Norrman · Tommy Lindell

5.3

Hispaniola (Haiti and the Dominican Republic) ........................................................................................ 285 Anja Scheffers · Tony Browne

5.4

Puerto Rico ................................................................................................................................................... 289 Jack Morelock ∙ Wilson Ramirez ∙ Maritza Barreto

ix

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Contents

5.4.1

Virgin Islands ....................................................................................................................................... 295 Maurice Schwartz

5.5

Lesser Antilles .............................................................................................................................................. 299 Gillian Cambers

6.0

Atlantic Ocean Islands – Editorial Introduction ........................................................................................ 311 Eric Bird

6.1

Bahamas ....................................................................................................................................................... 313 Alan K. Craig

6.2

Bermuda ....................................................................................................................................................... 317 Eric Bird

6.3

Greenland ..................................................................................................................................................... 319 Eric Bird

6.4

Iceland .......................................................................................................................................................... 323 Eggert Larusson

6.5

Other Atlantic Ocean Islands ...................................................................................................................... 333 Jim Hansom

7.0

British Isles – Editorial Introduction .......................................................................................................... 345 Eric Bird

7.1

England and Wales ....................................................................................................................................... 347 Eric Bird

7.2

Cumbria ....................................................................................................................................................... 353 Eric Bird

7.3

Lancashire, Merseyside and Cheshire ......................................................................................................... 358 Eric Bird

7.4

The Isle of Man .............................................................................................................................................. 361 Eric Bird

7.5

North Wales and Anglesey ........................................................................................................................... 366 Eric Bird

7.6

West Wales .................................................................................................................................................... 369 Eric Bird

7.7

South Wales ................................................................................................................................................... 375 Eric Bird

7.7.1

The Severn Estuary ............................................................................................................................... 384 Eric Bird

Contents

7.8

Gloucestershire, Somerset and North Devon .............................................................................................. 386 Eric Bird

7.9

Cornwall

7.9.1

The North Coast of Cornwall .............................................................................................................. 394 Eric Bird

7.9.2

The South Coast of Cornwall ............................................................................................................... 401 Eric Bird

7.10

Isles of Scilly ................................................................................................................................................. 410 Eric Bird

7.11

South Devon.................................................................................................................................................. 413 Eric Bird

7.12

Dorset ........................................................................................................................................................... 420 Eric Bird

7.13

Hampshire ..................................................................................................................................................... 429 Eric Bird

7.14

Isle of Wight ................................................................................................................................................. 433 Eric Bird

7.15

Sussex ........................................................................................................................................................... 438 Eric Bird

7.16

Kent .............................................................................................................................................................. 443 Eric Bird

7.17

Essex .............................................................................................................................................................. 447 Eric Bird

7.18

Suffolk ........................................................................................................................................................... 450 Eric Bird

7.19

Norfolk .......................................................................................................................................................... 454 Eric Bird

7.20

Lincolnshire ................................................................................................................................................. 459 Eric Bird

7.21

Yorkshire ...................................................................................................................................................... 461 Eric Bird

7.22

Durham, Tyne and Wear............................................................................................................................... 466 Eric Bird

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7.23

Northumberland ........................................................................................................................................... 468 Eric Bird

7.24

Scotland ........................................................................................................................................................ 473 William Ritchie · Alastair Dawson

7.24.1

South-East Scotland.............................................................................................................................. 477 William Ritchie

7.24.2

Fife......................................................................................................................................................... 481 William Ritchie

7.24.3

Angus and Aberdeenshire .................................................................................................................... 485 William Ritchie

7.24.4

Fraserburgh to Moray Firth and Duncansby Head ............................................................................. 489 William Ritchie

7.24.5

The Orkney and Shetland Islands......................................................................................................... 493 Alastair Dawson

7.24.6

North and North West Scotland (Duncansby Head to Cape Wrath and Fort William) ..................... 499 Alastair Dawson

7.24.7

The Inner Hebrides ............................................................................................................................... 505 Alastair Dawson

7.24.8

The Outer Hebrides .............................................................................................................................. 515 Alastair Dawson

7.24.9

Fort William to the Clyde .................................................................................................................... 523 Alastair Dawson

7.24.10

South-West Scotland............................................................................................................................. 529 William Ritchie

7.25

Northern Ireland .......................................................................................................................................... 535 Andrew Cooper

7.26

Ireland .......................................................................................................................................................... 545 Eric Bird

7.27

Channel Islands ............................................................................................................................................ 557 Eric Bird

8.0

Europe – Editorial Introduction .................................................................................................................. 569 Eric Bird

8.1

Norway ......................................................................................................................................................... 571 Tormod Klemsdal

Contents

8.1.1

Svalbard and Jan Mayen ....................................................................................................................... 581 Tormod Klemsdal

8.2

Sweden .......................................................................................................................................................... 589 John Norrman

8.3

Finland ......................................................................................................................................................... 595 Jouko Alestalo · Olavi Granö

8.3.1

Russian Gulf of Finland ........................................................................................................................ 601 Alexey Porotov

8.4

Estonia .......................................................................................................................................................... 605 Kaarel Orviku · Are Kont · Hannes Tõnisson

8.5

Latvia ............................................................................................................................................................ 613 Guntis Eberhards · Viktor Brenners

8.6

Lithuania ...................................................................................................................................................... 619 Eric Bird

8.7

Kaliningrad .................................................................................................................................................. 623 Alexey Porotov

8.8

Poland ........................................................................................................................................................... 627 Karol Rotnicki · Joanna Rotnicka

8.9

Germany ....................................................................................................................................................... 639 Klaus Schwarzer · Horst Sterr

8.10

Denmark ....................................................................................................................................................... 649 Troels Aagaard

8.11

The Netherlands ........................................................................................................................................... 665 Eric Bird

8.12

Belgium ........................................................................................................................................................ 669 Guy De Moor · André Ozer · Irénée Heyse

8.13

France

8.13.1

North Coast of France .......................................................................................................................... 673 Eric Bird

8.13.2

West Coast of France ........................................................................................................................... 683 Eric Bird

8.13.3

Mediterranean France .......................................................................................................................... 691 Eric Bird

8.13.4

Corsica ................................................................................................................................................. 695 Eric Bird

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Contents

8.14

8.14.1

Spain ............................................................................................................................................................. 699 Eric Bird Gibraltar ............................................................................................................................................... 711 Eric Bird

8.15

Portugal ........................................................................................................................................................ 713 Carlos Morais

8.16

Italy ............................................................................................................................................................... 719 Paolo Ciavola

8.16.1

Sardinia ................................................................................................................................................ 731 Eric Bird

8.16.2

Sicily ..................................................................................................................................................... 741 Paolo Ciavola

8.17

Malta ............................................................................................................................................................. 751 George Said · John Schembri

8.18

Slovenia ........................................................................................................................................................ 761 Roger Charlier

8.19

Croatia .......................................................................................................................................................... 763 Roger Charlier

8.20

Bosnia-Herzegovina ..................................................................................................................................... 769 Roger Charlier

8.21

Montenegro .................................................................................................................................................. 771 Roger Charlier

8.22

Albania ......................................................................................................................................................... 773 Yurii Shuisky

8.23

Greece ........................................................................................................................................................... 775 Eric Bird

8.24

Crete ............................................................................................................................................................. 783 Anja Scheffers · Tony Browne

8.25

Bulgaria ........................................................................................................................................................ 785 Eric Bird

8.26

Romania ....................................................................................................................................................... 789 Roger Charlier · Constance Charlier

8.27

Ukraine ......................................................................................................................................................... 795 Yurii Shuisky

Contents

8.28

Sea of Azov ................................................................................................................................................... 799 Eric Bird

8.29

Republic of Georgia ..................................................................................................................................... 805 Pavel Kaplin · Andrei Selivanov Index Maps ................................................................................................................................................. 1471 List of Entries ............................................................................................................................................. 1483 Subject Index .............................................................................................................................................. 1489 Volume II Editor .......................................................................................................................................................... xxiii List of Contributors ...................................................................................................................................... xxv

9.0

Russian Federation – Editorial Introduction .............................................................................................. 809 Eric Bird

9.1

Russian Gulf of Finland ............................................................................................................................... 810 Alexey Porotov

9.2

Russian Black Sea Coast ............................................................................................................................... 811 Pavel Kaplin · Andrei Selivanov

9.3

The Pacific Coast of Russia .......................................................................................................................... 815 Pavel Kaplin · Andrei Selivanov

9.4

The Arctic Coast of Russia ........................................................................................................................... 823 Andrei Selivanov

9.5

Russian Caspian Coast.................................................................................................................................. 829 Svetlana Lukyanova

9.6

Kaliningrad ................................................................................................................................................... 829 Alexey Porotov

10.0

Middle East – Editorial Introduction .......................................................................................................... 831 Eric Bird

10.1

Turkey ........................................................................................................................................................... 833 Oguz Erol

10.2

Cyprus .......................................................................................................................................................... 841 Yaacov Nir

10.3

Syria ............................................................................................................................................................... 849 Eric Bird

10.4

Lebanon ........................................................................................................................................................ 851 Eric Bird

xv

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Contents

10.5

Israel (with the Gaza Strip) ........................................................................................................................... 853 Yaacov Nir

11.0

Caspian and Aral Seas – Editorial Introduction ......................................................................................... 859 Eric Bird

11.1

Iran – Caspian Sea Coast .............................................................................................................................. 861 Eric Bird

11.2

Turkmenistan ................................................................................................................................................ 867 Andrei Selivanov

11.3

Kazakhstan ................................................................................................................................................... 871 Andrei Selivanov

11.4

Russian Caspian Coast ................................................................................................................................. 875 Svetlana Lukyanova

11.5

Azerbaijan .................................................................................................................................................... 885 Pavel Kaplin · Andrei Selivanov · Svetlana Lukyanova

11.6

Aral Sea ......................................................................................................................................................... 889 Eric Bird

12.0

North Africa – Editorial Introduction ........................................................................................................ 891 Eric Bird

12.1

Egypt (Mediterranean) ................................................................................................................................ 893 Anja Scheffers ∙ Tony Browne

12.2

Libya ............................................................................................................................................................. 897 Maurice Schwartz

12.3

Tunisia .......................................................................................................................................................... 899 Eric Bird

12.4

Algeria .......................................................................................................................................................... 903 M. Larid

12.5

Morocco ........................................................................................................................................................ 907 Anja Scheffers

13.0

West Africa – Editorial Introduction .......................................................................................................... 915 Eric Bird

13.1

Mauritania .................................................................................................................................................... 917 Don Vermeer

13.2

Senegal and Gambia ..................................................................................................................................... 921 André Guilcher

Contents

13.3

Guinea Bissau................................................................................................................................................ 927 E.S. Diop

13.4

Republic of Guinea ....................................................................................................................................... 929 E.S. Diop

13.5

Sierra Leone ................................................................................................................................................. 931 Eric Bird

13.6

Liberia .......................................................................................................................................................... 935 Eric Bird

13.7

Ivory Coast ................................................................................................................................................... 939 Anja Scheffers

13.8

Ghana ........................................................................................................................................................... 943 L.A. Dei

13.9

Togo and Benin ............................................................................................................................................ 947 André Guilcher

13.10 Nigeria .......................................................................................................................................................... 949 Etop Usoro 13.11 Cameroon and Equatorial Guinea ............................................................................................................... 953 Maurice Schwartz 13.12 Gabon, Congo, Cabinda and Zaïre .............................................................................................................. 957 P. Giresse 13.13 Angola .......................................................................................................................................................... 963 André Guilcher 13.14 Namibia ........................................................................................................................................................ 969 Eric Bird · A. Goudie · H. Viles

14.0

South Africa – Introduction ........................................................................................................................ 975 Gerald G. Garland

14.1

South Africa ................................................................................................................................................. 979 Andrew Cooper

15.0

East Africa – Editorial Introduction ........................................................................................................... 987 Eric Bird

15.1

Mozambique ................................................................................................................................................. 989 Maria Eugénia Soares de Albergaria Moreira

15.2

Tanzania ....................................................................................................................................................... 995 Eric Bird

xvii

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Contents

15.3

Kenya .......................................................................................................................................................... 1003 Francis Ojany

15.4

Somalia and Djibouti ................................................................................................................................. 1007 Eric Bird

15.5

Eritrea ......................................................................................................................................................... 1011 Eric Bird

15.6

Sudan .......................................................................................................................................................... 1013 Eric Bird

15.7

Egypt, Red Sea Coast ................................................................................................................................. 1015 Anja Scheffers · Tony Browne

15.8

Jordan ......................................................................................................................................................... 1019 Yaacov Nir

16.0

South-West Asia – Editorial Introduction ................................................................................................ 1021 Eric Bird

16.1

Saudi Arabia, Red Sea Coast ...................................................................................................................... 1023 Eric Bird

16.2

Yemen ......................................................................................................................................................... 1027 Eric Bird

16.3

Southern Arabia and Oman ....................................................................................................................... 1029 Eric Bird

16.4

United Arab Emirates ................................................................................................................................ 1033 Eric Bird

16.5

Qatar ........................................................................................................................................................... 1037 Eric Bird

16.6

Bahrein ....................................................................................................................................................... 1041 Eric Bird

16.7

Saudi Arabia, Persian Gulf Coast .............................................................................................................. 1045 Eric Bird

16.8

Kuwait and Iraq .......................................................................................................................................... 1047 Eric Bird

16.9

Iran ............................................................................................................................................................. 1051 Rodman E. Snead

17.0

South Asia – Editorial Introduction .......................................................................................................... 1057 Eric Bird

Contents

17.1

Pakistan ...................................................................................................................................................... 1059 Rodman E. Snead

17.2

India ........................................................................................................................................................... 1065 G. N. Nayak · P. T. Hanamgond

17.3

Sri Lanka .................................................................................................................................................... 1071 Eric Bird

17.4

Bangladesh ................................................................................................................................................. 1077 Rodman E. Snead

17.5

Burma (Myanmar) ..................................................................................................................................... 1081 Kyaw Saw Lynn

18.0

Indian Ocean Islands – Editorial Introduction ......................................................................................... 1087 Eric Bird

18.1

Madagascar ................................................................................................................................................ 1089 Jean-Michel Lebigre

18.2

Indian Ocean Islands ................................................................................................................................. 1097 Wong Poh Poh

19.0

South-East Asia – Editorial Introduction ................................................................................................. 1111 Eric Bird

19.1

Thailand Andaman Sea Coast .................................................................................................................... 1113 Sanit Aksornkoae · Eric Bird

19.2

Malaysia – Introduction ............................................................................................................................. 1117 Teh Tiong Sa · Yap Hui Boon

19.3

Singapore .................................................................................................................................................... 1129 Wong Poh Poh

19.3.1

Thailand: Gulf of Thailand Coast ...................................................................................................... 1135 Sanit Aksornkoae · Eric Bird

19.4

Brunei (Negara Brunei Darussalam) ......................................................................................................... 1141 Gabriel Yong

19.5

Cambodia ................................................................................................................................................... 1145 Eric Bird

19.6

Vietnam ...................................................................................................................................................... 1147 D. Eisma

19.7

Philippines ................................................................................................................................................. 1151 Eric Bird

xix

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Contents

19.8

Indonesia .................................................................................................................................................... 1157 Otto Ongkosongo

19.9

East Timor .................................................................................................................................................. 1171 Wong Poh Poh

19.10 Papua New Guinea ..................................................................................................................................... 1175 Eric Bird

20.0

East Asia – Editorial Introduction ............................................................................................................ 1187 Eric Bird

20.1

China .......................................................................................................................................................... 1189 Chen Jiyu · Liu Cangzi · Dai Zhijun · Yu Zhiying

20.2

Taiwan ........................................................................................................................................................ 1197 John R. C. Hsu

20.3

North Korea ................................................................................................................................................ 1201 D. Eisma

20.4

South Korea ................................................................................................................................................ 1207 D. Eisma

20.5

Japan ........................................................................................................................................................... 1213 Kazuyuki Koike

21.0

Australia – Editorial Introduction ............................................................................................................ 1225 Eric Bird

21.1

New South Wales ........................................................................................................................................ 1229 Bruce Thom

21.1.1

Lord Howe Island – (New South Wales) ............................................................................................ 1239 Eric Bird

21.1.2

Norfolk Island .................................................................................................................................... 1247 Eric Bird

21.2

Queensland ................................................................................................................................................. 1255 David Hopley · Scott Smithers

21.3

Northern Territory ..................................................................................................................................... 1267 Eric Bird

21.4

Western Australia ....................................................................................................................................... 1279 Ian Eliot · Eric Bird

21.5

South Australia .......................................................................................................................................... 1293 C. R. Twidale · J. A. Bourne

Contents

21.6

Victoria Introduction ................................................................................................................................. 1305 Eric Bird

21.6.1

Victoria: The Western District (Discovery Bay to Mepunga) ............................................................ 1311 Eric Bird

21.6.2

Victoria: The Port Campbell Coast (Mepunga to Princetown) ......................................................... 1319 Eric Bird

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale) .......................... 1325 Eric Bird

21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)............................................................ 1337 Eric Bird

21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head) .................................................... 1349 Eric Bird

21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island.................................................... 1359 Eric Bird

21.6.7

South Gippsland, Wilsons Promontory and Corner Inlet ................................................................ 1369 Eric Bird

21.6.8

Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe) ................................................... 1377 Eric Bird

21.6.9

Victoria: The Gippsland Lakes ........................................................................................................... 1385 Eric Bird

21.7

Tasmania .................................................................................................................................................... 1395 Joanna Ellison

22.1

New Zealand ............................................................................................................................................... 1403 Terry Healy

22.2

The Chatham Islands ................................................................................................................................. 1413 Eric Bird

23.0

Pacific Ocean Islands – Editorial Introduction ......................................................................................... 1423 Eric Bird

23.1

New Caledonia and the Loyalty Islands ..................................................................................................... 1425 Jean-Michel Lebigre · Eric Bird

23.2

Fiji ............................................................................................................................................................... 1437 Patrick Nunn

23.3

Society Islands ............................................................................................................................................ 1451 André Guilcher

xxi

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Contents

23.4

Other Pacific Islands .................................................................................................................................. 1455 Joanna Ellison

24.1

Antarctic Coast .......................................................................................................................................... 1463 Eric Bird Index Maps ................................................................................................................................................. 1471 List of Entries ............................................................................................................................................. 1483 Subject Index .............................................................................................................................................. 1489

Editor Eric C.F. Bird, M.Sc. (Lond. and Melb.) Ph.D. (ANU) Principal Fellow in Geomorphology University of Melbourne and Director, Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected]

List of Contributors Troels Aagaard Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen Denmark [email protected] Sanit Aksornkoae Thailand Environment Institute (TEI) 16/151 Muang Thong Thani Bond Street Bangpood, Pakkred Nonthaburi 11120 Thailand Jouko Alestalo Department of Geography University of Turku Turku 20014 Finland José Araya-Vergara Department of Geography University of Chile Av. Portugal 84, Cas. 3387 Santiago de Chile Chile [email protected] H. Ayon Escuela Superior Politecnica del Litoral (ESPOL) Centro Nacional de Recursos Costeros Malecon No 100 y Loja Guayaquil Ecuador

Germán Bértola Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Centro de Geología de Costas y del Cuaternario Funes 3350, Mar del Plata Argentina [email protected] Eric C.F. Bird Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected] Henry Bokuniewicz Marine Sciences Research Center State University of New York Stony Brook, NY 11794-5000 USA [email protected] Yap Hui Boon Hwa Chong Institution 661, Bukit Timah Road, 269734 Singapore J.A. Bourne School of Earth and Environmental Sciences Geology and Geophysics University of Adelaide Adelaide 5005 South Australia Viktor Brenners 114 Princess Street Kew, Victoria Australia Tony Browne†

Maritza Barreto Marine Sciences Department University of Puerto Rico at Mayagüez P.O. Box 1127 Lajas, PR Puerto Rico [email protected]

Mary-Louise Byrne Geography and Environmental Studies Wilfrid Laurier University Waterloo, Ontario Canada N2L 3C5 [email protected]

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List of Contributors

Gillian Cambers Sea Grant College Program University of Puerto Rico at Mayagüez P.O. Box 9011 Puerto Rico 681 [email protected] Liu Cangzi State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China Constance Charlier C.E.O. Inc. Tucson, AZ USA Roger Charlier 2 avenue du Congo box 23 B-1050 Brussels Belgium [email protected] Paolo Ciavola Dipartimento di Scienze della Terra Università di Ferrara Corso Ercole I d‘Este, 32 44100 Ferrara Italy [email protected] John J. Clague Department of Earth Sciences Simon Fraser University 8888 University Drive Burnaby, British Columbia Canada V5A 1S6 [email protected] Andrew Cooper School of Environmental Science University of Ulster Cromore Road Coleraine BT52 1SA Northern Ireland UK [email protected]

Iván Correa Departamento de Geología Universidad Eafit P.O. Box 3300 Medellin, Colombia South America [email protected] Alan K. Craig Florida Atlantic University Boca Raton, FL 33431 USA Alastair Dawson AICSM University of Aberdeen Elphinstone Road Aberdeen AB24 3UF Scotland UK [email protected] L.A. Dei Department of Geography & Regional Planning University of Cape Coast Box UC 0180, University Post Office Cape Coast Ghana [email protected] E.S. Diop University CAD Dakar Senegal Jean-Marie M. Dubois Département de géomatique appliquée Université de Sherbrooke Sherbrooke, Québec Canada J1K 2R1 [email protected] Guntis Eberhards Faculty of Geography and Earth Sciences University of Latvia Raiņa bulv. 19 Rīga Latvia [email protected]

List of Contributors

D. Eisma Oude Molstraat 12a 2513 BB Den Haag The Netherlands Ian Eliot Faculty of Natural and Agricultural Sciences University of Western Australia [email protected] L. Ellenberg Humboldt-Universität zu Berlin Geographisches Institut Chausseestraße 86 D- 10099 Berlin [email protected] Joanna Ellison Department of Geography University of Tasmania Launceston Tasmania 7001 Australia [email protected] Oguz Erol Ankara Üniversitesi 4. Kat Koridoru Ankara Turkey 6090 [email protected] Eden J. Feirstein Department of Geology and Geophysics School of Ocean and Earth Science and Technology (SOEST) University of Hawaii 1680 East-West Rd Honolulu, HI 96822 USA Charles Finkl Department of Geosciences Florida Atlantic University Boca Raton, FL 33431 USA [email protected]

Duncan FitzGerald Department of Earth Sciences Boston University 675 Commonwealth Avenue Boston, MA 02215 USA [email protected] Charles H. Fletcher Department of Geology and Geophysics School of Ocean and Earth Science and Technology (SOEST) University of Hawaii 1680 East-West Rd Honolulu, HI USA [email protected] Gerald G. Garland Department of Geography and Urban Planning University of United Arab Emirates P.O. Box 17771 Al Ain United Arab Emirates [email protected] Ioannis Georgiou Department of Earth and Environmental Sciences Pontchartrain Institute for Environmental Sciences University of New Orleans New Orleans, LA 70148 USA [email protected] P. Giresse Laboratoire d’Étude des Géo-Environnements Marins Université de Perpignan 52, Av. Paul Alduy 66860 Perpignan France A. Goudie School of Geography and the Environment Oxford University South Parks Road, Oxford, OX1 3QY United Kingdom [email protected]

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Olavi Granö Department of Geography University of Turku Turku 20014 Finland [email protected]

David Hopley School of Earth and Environmental Sciences James Cook University Townsville, Queensland 4811 Australia [email protected]

André Guilcher†

John R. C. Hsu Department of Marine Environment and Engineering National Sun Yat-sen University 70 Lienhai Road, Gushan District Kaohsiung City 80424 Taiwan [email protected]

Mario Gutiérrez-Estrada Institute of Marine Sciences Mazatlan Mexico P.T. Hanamgond Dept. of Geology GSS College Tilakwadi, Belgaum India Jim Hansom Department of Geographical and Earth Sciences University of Glasgow Glasgow, G12 8QQ Scotland UK [email protected] Miles Hayes Department of Geology University of South Carolina Columbia, SC 29208 USA [email protected] Terry Healy Department of Earth and Ocean Sciences University of Waikato Hamilton 3240 New Zealand [email protected] Irénée Heyse Department of Geography University of Ghent Krijgslaan 281-S8 Ghent Belgium [email protected]

Federico Isla Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Centro de Geología de Costas y del Cuaternario Funes 3350, Mar del Plata Argentina [email protected] W. Jara† Chen Jiyu State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China [email protected] Pavel Kaplin Department of Geology Odessa State University Odessa 270000 Ukraine [email protected] Tormod Klemsdal Department of Geosciences University of Oslo P.O. Box 1047 Blindern 316 Oslo Norway [email protected]

List of Contributors

Kazuyuki Koike Department of Geography Komazawa University Komazawa, Setagaya-ku Tokyo 154 Japan Paul D. Komar College of Oceanic & Atmospheric Sciences Oregon State University Corvallis, OR 97331 USA [email protected] Are Kont Institute of Ecology Tallinn University Narva mnt 25 Tallinn Estonia [email protected] Mark Kulp Department of Earth and Environmental Sciences University of New Orleans 2000 Lakeshore New Orleans, LA USA [email protected] M. Larid University Center Stambouli Mustapha of Mascara Route de Mamounia 29000 Mascara Algeria Eggert Larusson School of Education University of Iceland Stakkahlíð IS 101 Reykjavík Iceland [email protected] Jean-Michel Lebigre Université de la Nouvelle-Calédonie BPR4 98851 Nouméa Cedex New Caledonia Canada [email protected]

Tommy Lindell Centre for Image Analysis Swedish University of Agricultural Sciences Uppsala University Box 337 Uppsala Sweden [email protected] Svetlana Lukyanova Faculty of Geography M.V. Lomonosov Moscow State University Moscow Russia [email protected] Kyaw Saw Lynn Anglo-Chinese Junior College 25 Dover Close East Singapore [email protected] Molly McGraw Southeastern Louisiana University Hammond Louisiana, LA 70402 USA [email protected] [email protected] Michael Miner Department of Earth and Environmental Sciences University of New Orleans New Orleans, LA 70148 USA Guy De Moor Geological Institute University of Ghent Krijgslaan 281-S8 Ghent Belgium Carlos Morais Laboratório Nacional de Engenharia Civil Av do Brasil, 101 Lisboa Portugal

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Maria Eugénia Soares de Albergaria Moreira Quinta das Mós - Rua da Cruz nr 5 Zibreira da Fé Sobral de Monte Agraço Portugal 2590–293 [email protected]

Patrick Nunn Department of Geography University of the South Pacific P.O. Box 1168 Suva, Fiji Islands [email protected]

Jack Morelock Department of Marine Sciences University of Puerto Rico at Mayagüez P.O. Box 908 Lajas, PR Puerto Rico

Francis Ojany Department of Geography University of Nairobi 30197 Nairobi Kenya 00100-GPO [email protected]

Robert Morton United States Geological Survey 10100 Burnet Rd., Bldg. 130 Austin, Texas 78758 USA [email protected] Dieter Muehe Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen Denmark [email protected] G. N. Nayak Department of Marine Science Goa University Taleigao Plateau Goa 403206 India [email protected] [email protected] Yaacov Nir 15 Shimeoni st. Rehovot 76248 Israel [email protected] John Norrman Department of Earth Sciences Uppsala University Villav. 16 Uppsala Sweden

Otto Ongkosongo Lembaga Oseanologi Nasional Jl. Pasir Putih I Ancol Timur Jakarta Utara Indonesia [email protected] Mario Arturo Ortiz-Perez Institute of Marine Sciences Mazatlan Mexico Kaarel Orviku GONSIORI str. 24-23 Talinn Estonia [email protected] Ervin Otvos Department of Coastal Sciences Ocean Springers Campus University of Southern Mississippi Ocean Springs MS 39566 USA [email protected] André Ozer Département de Géographie Physique Université de Liège 2 Allée du 6 Août Liège Belgium [email protected]

List of Contributors

Wong Poh Poh Department of Geography National University of Singapore 21 Lower Kent Ridge Road Singapore 119077 [email protected]

Vidal Taype Ramos Geophysics Commission Pan American Institute of Geography and History Av. A. Aramburú 1190-1198 Lima 34 Peru

Wayne Pollard Department of Geography McGill University 805 Sherbrooke St. W. Montreal, Quebec Canada H3A 2K6 [email protected]

Stanley R. Riggs Department of Geology Thomas Harriot College of Arts and Sciences East Carolina University 1002 Bate Building East Fifth Street Greenville, NC USA [email protected]

Alexey Porotov Faculty of Geography Lomonosov Moscow State University GSP-1, Leninskie Gory Moscow, 119991 Russian Federation [email protected] Jorge Pousa Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Universidad Nacional de La Plata (UNLP) Calle 64 (entre 119 y 120) La Plata, Buenos Aires Province Argentina [email protected]

William Ritchie AICSM University of Aberdeen Elphinstone Road Aberdeen AB24 3UF Scotland UK [email protected] Ridel Rodríguez Apartado postal 127 Holguín, C.P. Cuba 80 100 [email protected] [email protected]

Norbert Psuty Institute of Marine and Coastal Sciences Rutgers University 74 Magruder Road Sandy Hook, NJ USA [email protected]

Joanna Rotnicka Institute of Geology Adam Mickiewicz University Maków Polnych 16 Poznañ Poland [email protected]

Garry Ramedine Université des antilles et de la guyane Campus de Fouillole, BP 250 Pointe-à-Pitre, Grande-Terre Guadeloupe

Karol Rotnicki Institute of Geoecology and Geoinformation Adam Mickiewicz University Ul. Dziegielowa 27 Poznan Poland [email protected]

Wilson Ramirez Geology Department University of Puerto Rico at Mayagüez P.O. Box 9017 Mayagüez, PR Puerto Rico

Teh Tiong Sa 19 Jalan Jintan #12-25 Singapore [email protected]

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Pascal Saffache Université des antilles et de la guyane Campus de Fouillole, BP 250 Pointe-à-Pitre, Grande-Terre Guadeloupe [email protected] George Said Favorita Sir Paul Boffa street Victoria Gozo, Malta [email protected] Paul Sanlaville Le Sanderey St. Michel de Villadeix F-24380 France Anja Scheffers School of Environmental Science and Management Southern Cross University P.O. Box 157 Lismore, NSW 2480 Australia [email protected] John Schembri Favorita Sir Paul Boffa street Victoria Gozo, Malta Enrique Schnack Laboratorio de Oceanografia Costera Facultad de Ciencias Naturales y Museo Universidad Nacional de La Plata C.C. 45, 1900 La Plata, Buenos Aires Argentina [email protected] Maurice Schwartz Department of Geology Western Washington University Bellingham, Washington 98225 USA [email protected]

Klaus Schwarzer Coastal and Continental Shelf Research Institute of Geosciences University of Kiel Olshausenstrasse 40 Kiel Germany [email protected] Andrei Selivanov† Yurii Shuisky Physical Geography Department Mechnikov’s National University of Odessa 2 Dvoryanskaya St. Odessa-82 Ukraine [email protected] Scott Smithers School of Earth and Environmental Sciences Faculty of Science and Engineering James Cook University Queensland 4810 Australia Rodman E. Snead Department of Geography University of New Mexico Albuquerque New Mexico USA Horst Sterr Institute of Geography University of Kiel Olshausenstrasse 40 Kiel Germany [email protected] Thomas Terich Center for Geography and Environmental Social Sciences Huxley College of Environmental Studies Western Washington University Bellingham, WA 98225 USA [email protected]

List of Contributors

Bruce Thom School of Geosciences The University of Sydney Australia [email protected] Hannes Tõnisson Institute of Ecology Tallinn University Tallinn Estonia [email protected] J. Turenne Centre d’Ile-de-France 32, avenue Henri Varagnat 93143 Bondy cedex France C. R. Twidale Department of Geography University of Adelaide Adelaide, South Australia 5005 Australia [email protected] Etop Usoro Department of Geography and Regional Planning University of Uyo 1, Ikpa Road, P.M.B. 1017 Uyo, Akwa Ibom State Nigera Don Vermeer Louisiana State University Baton Rouge, LA 70803 USA

H. Viles School of Geography and the Environment Oxford University South Parks Road, Oxford, OX1 3QY United Kingdom [email protected] Jesse Walker Department of Geography Louisiana State University Baton Rouge, Louisiana 70803 USA [email protected] Gabriel Yong Department of Geography Universiti Brunei Darussalam Jalan Tungku Link Gadong Negara Brunei Darussalam BE1310 [email protected] Dai Zhijun State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China Yu Zhiying State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China

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1.0 USA – Editorial Introduction

The major geomorphological divisions of the United States, the Rocky Mountain ranges, the Great Plains and the Appa­lachian ranges, do not have a major influence on coastal features. The Pacific coast follows the alignment of the Coast Ranges on the western side of the Rocky Mountains, the Great Plains run southward to the lowlying coast of the Gulf of Mexico, and the Appalachians extend north-eastward to the coast of Maine. Useful background is provided by Shepard and Wanless (1971), Sherman (2005), Thornbury (1965) and Graf (1987).

1.  Alaska and Pacific Coast The Alaskan coast (> Alaska) is long, varied, and irregular, some of it steep and mountainous. It has developed where the northern extension of the Rock Mountain ranges swings westward, and is separated by valleys along intermontane basins. It has been strongly glaciated, and some of the fiords are still occupied by glaciers, as in Nunatak Fiord, though they have been receding. Tectonic activity has been vigorous along the Alaskan Peninsula and the Aleutian Island, where there are several active volcanoes along the plate convergence. South-eastern Alaska is rising rapidly, locally up to 3.5 cm/year. Mean spring tide ranges are small on the Alaskan Peninsula and the Aleutian Islands, but increase in the Gulf of Alaksa to more than 10 m near Anchorage. Much of the Pacific coast ( > Washington,  > Oregon, > California) is steep and high, and subject to continuing tectonic activity, including movements along the San Andreas Fault and tectonic disruptions associated with the subduction zone offshore, where the Pacific Plate is passing beneath the North American Plate. The continental shelf is narrow ­(20–30 km), and the coast is exposed to strong wave energy from ocean swell transmitted across the Pacific as well as storms in coastal waters. Mean spring tide ranges are generally small ­(1–3 m), but the coast is subject to occasional tsunamis, including those generated by tectonic disruptions along the offshore zone of converging plates. The steep coasts are generally vegetated slopes descending to rocky cliffs. They are interrupted by the mouths of numerous valleys, with marine inlets or coastal lagoons

(bays) enclosed by spits and barriers, mainly sandy. The spits and barriers are generally capped by dunes, and dunes become extensive in some sectors, ­notably in Central Oregon, where they have advanced several kilometres inland. Fluvial sediment from the Columbia River has been the major source of sandy beach material, which has drifted from the river mouth both north into Washington and south into Oregon. Tectonic activity has resulted in the uplift of marine terraces. These are prominent in California, where they are found at various levels up to 411 m on the Palos Verde Peninsula, rising to 600 m near Santa Barbara, then falling to 200 m near Big Sur. San Francisco Bay is a large basin dating from Pliocene times that has been invaded by the sea. Coastal rock formations from Washington State south to northern California include mélanges of sedimentary, metamorphic, and volcanic rocks assembled in a tectonically disturbed environment. The numerous stacks and islands off this coast, sometimes extending several kilometres seaward, are largely the more resistant residuals of these formations, isolated in the course of long-term coastline recession. The recession of steep, forested coastal bluffs in Washington and Oregon has been intermittent, the cumulative result of recurrent landslides and episodes of marine erosion during tsunamis. On the Olympic Peninsula in Washington, the coast had been influenced by glacial and periglacial processes during cold phases of the Pleistocene. One indication is the presence of gravelly beaches.

2.  Gulf of Mexico Coast Several rivers have fed fluvial sand and silt to the Gulf of Mexico coast, the largest being the Mississippi, which has built a major composite delta. In > Texas, there are large sandy dune-capped barriers, backed by beach ridges, lagoons, and marshes and deltaic sectors such as that at the mouth of Brazos River. The coastal plain is wide in > Louisiana, fringed seaward by beaches and cheniers perched on alluvial flats. The tide range on the Gulf coast is small, and the coast has been shaped by generally gentle Caribbean waves and occasional hurricane surges. East of the Mississippi delta, wave energy is relatively low, but

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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USA – Editorial Introduction

s­ ufficient to have formed beaches and barrier islands on the coasts of > Mississippi and > Alabama. Further east, it diminishes along the Florida coast (> Florida), and is exceptionally low in Apalachee Bay. The western coast of the Florida peninsula is low-lying with extensive marshes and mangroves threaded by inlets.

3.  Atlantic Coast Much of the Atlantic coast of the United States is a lowlying trailing edge coast with a continental shelf up to 100 km wide. Long sandy barriers and dune-capped barrier islands extend along the Atlantic coast from Florida to New Jersey (> Florida, > Georgia, > South Carolina, > North Carolina, > Central Atlantic Coast, > New York and New England, > Great Lakes (USA) and > Hawaii), backed by intermittent lagoons. The tide range is small (1–2 m) and wave action mainly south-easterly from Atlantic storms; hurricanes have occasionally produced surges causing erosion and flooding. In the absence of rivers yielding sediment to the sea, the beaches, barriers, and dunes are predominantly quartzose, with local accessions of shelly material. Longshore drift is predominantly northward, forming compound recurved spits such as Cape Henry and Sandy Hook, but there are episodes of southward drifting, the alternations contributing to the shaping of the major cuspate forelands, Cape Canaveral, Cape Fear, Cape Lookout, and Cape Hatteras. Each of these has beach and dune ridges marking stages in evolution. There are large estuarine embayments such as Delaware Bay, Chesapeake Bay, and the barrier-enclosed Pamlico Sound. Many of the coastal lagoons and inlets have muddy areas, with salt marshes. In Florida, there are intertidal sabellarid reefs, and to the south coral reefs. Beaches are more calcareous, with sediment derived from these, and mangroves appear in the salt marshes. In New England (> New York and New England) to the north the coast is hilly and irregular on structures associated with the northern Appalachians. Glaciation extended as far south as New York, and has left numerous valley-mouth inlets, submerged during the Holocene marine transgression, many of them now blocked by barriers or sand and shingle to form lagoons known as ponds. There are also glacial outwash features and drumlins, as in Boston Harbour, and moraines, including Long Island, New York. Reworked glacial drift has produced sand, pebbles, cobbles, and some boulders. There are sand and shingle spits and beaches and sand dunes, as on Long Island. Staten Island, New York, is a morainic feature close to the limit of glaciation.

4.  Great Lakes The Great Lakes (> Great Lakes (USA)) are a group of linked freshwater lakes discharging into the St. Lawrence River. Although they form a natural unit, it is necessary to split them because the boundary between the United States and > Canada passes through Lakes Superior, Huron, Erie, and Ontario (so that their southern coastlines are in the United States), whereas Lake Michigan is entirely within the United States. The present Great Lakes began to form about 15,000 years ago when the Wisconsin ice sheet was melting, and water accumulated between the receding ice margin and a drainage divide to the south. The shaping of the lakes was influenced by the topography exposed as the ice retreated and by uplift as the ice load decreased (postglacial isostatic recovery). Lake Erie is 174 m above mean sea level, and outflow at Buffalo to the Niagara River descends Niagara Falls to Lake Ontario, 75 m above sea level, and then along St. Lawrence River to the sea. The lakes are almost tideless, but show fluctuations in level related to atmospheric pressure, wind stress, precipitation and evaporation, and freshwater inflow and outflow regimes. Seasonal fluctuations are of the order of 0.35 m, but longer-term oscillations of up to 2 m have occurred. Waves are generated by local winds, and may be up to 4 m high during storms. Parts of the lakes coastline are rocky on Pre-Cambrian and Palaeozoic igneous and sedimentary formations, but glacial drift deposits are extensive, and in places these have been cut back as cliffs. Sand and gravel derived from glacial drift form narrow (<15 m) beaches, while sand has accumulated as dunes locally, notably on the SE coast of Lake Michigan. Longshore drifting results from the arrival of waves at an angle to the shore, and is predominantly eastward on southern coasts and southward on much of the shores of Lake Michigan and Lake Huron, as indicated by the deflection of stream mouths and the growth of spits.

5.  Hawaii The Hawaiian archipelago (> Hawaii) is of volcanic origin, formed by extrusion of basalt lava along a midPacific zone where a hot spot has migrated NW-SE beneath the Pacific plate. Successively formed volcanic structures have been dissected by Pacific Ocean swell and storm waves to form cliffs, some of which are high and ­spectacular. Coastal landforms include craters invaded by the sea. Boulders, gravel, and sand derived from the volcanic rocks

USA – Editorial Introduction  

form beaches. Wave energy is high, concentrated on these microtidal coasts, and wave break heavily on surf beaches. Coral reefs fringe some sectors of coast, and coralline sand and gravel is a constituent of many beaches, along with shelly and algal sediment. Some calcareous beaches are backed by dunes and dune calcarenite. There are emerged reefs and associated shore features resulting from tectonic uplift and sea level variations, and the islands have been subject to recurrent tsunamis, originating from earthquakes and volcanic eruptions around the rim of the Pacific basin. A major tsunami in 1946, caused by an earthquake off the Aleutian Islands, produced giant waves on Hawaiian coasts, as did the tsunami resulting from the Chilean earthquake in 1960.

   1.0

References Graf WL (ed) (1987) Geomorphic systems of North America. Geological Society of America, Boulder, CO Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw Hill, New York Sherman DJ (2005) North America coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 721–727 Thornbury WD (1965) Regional geomorphology of the United States. Wiley, New York

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1.1 Alaska Jesse Walker · Molly McGraw 

1.  Introduction Alaska, which spans 20° of latitude and 50° of longitude, has a coastline about 10,686 km in length. Because of its position and shape, it faces three distinct bodies of water, the Arctic Ocean on the north and northwest, the Pacific Ocean on the south, and the smaller, semi-enclosed Bering Sea on the west. Alaska occupies the northwest extension of four of North America’s major physiographic divisions (­Wah­rhaftig 1965): the Interior Plains, the rocky Mountain System, the Intermontane Plateaus, and the Pacific ­Moun­tain System

(> Fig. 1.1.1). All of these physiographic divisions extend across the Alaskan coast and each has an important bearing on coastal geomorphology and geology (> Fig. 1.1.2) and coastline character (> Fig. 1.1.3). The Interior Plains in Alaska, known as the Arctic Coastal Plain, is composed almost entirely of Quaternary sediments most of which are unconsolidated and some aeolian. The coastal zone of the Brooks Range (the Alaskan portion of the Rocky Mountain System) consists of Mesozoic and Palaeozoic sedimentary and metamorphic rocks. Most of the Alaskan Bering Sea Coast belongs to the Intermontane Plateaus. It is composed mainly of

⊡⊡ Fig. 1.1.1 Physiographic divisions of Alaska.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Alaska

⊡⊡ Fig. 1.1.2 Coastal geology of Alaska.

Quaternary sediment although there are several sections where Mesozoic and Palaeozoic sedimentary and metamorphic rocks predominate. In addition, a few locations have Quaternary and Tertiary volcanics and intrusive igneous rocks of Tertiary age. The geology of the Pacific Mountain System is most diverse. Virtually, all rock types are present. In the Aleutian Islands, Quaternary volcanics predominate; along the central Gulf of Alaska coast, sedimentary and metamorphic rocks of Quaternary, Tertiary, and Mesozoic ages are all present; and along the lengthy coastline of southeast Alaska there is a mixture of Tertiary and Mesozoic intrusives and Mesozoic and Palaeozoic sedimentary and metamorphic rocks. The continental shelf of Alaska varies greatly in width being quite narrow along the Pacific Ocean and quite wide in the Bering Sea and the Arctic Ocean. With a width of only about 150 km opposite Demarcation Point (the

northwest corner of Alaska), it widens westward toward the Siberian Coast. The northern half of the Bering Sea is sufficiently shallow so that during glacial epochs, it was subaerial, connecting eastern Siberia with western Alaska. Along the Pacific coast of Alaska, the continental shelf is bordered seaward by the Aleutian Trench, which results from the subduction of the Pacific Plate beneath the North American Plate. These contrasts are also evidenced in the amount of tectonic activity affecting the coastal zone.Along the Arctic Ocean and Bering Sea coastlines, there is minimal earthquake and volcanic activity, whereas along the Gulf of Alaska coast, it tends to be intense. The climate of coastal Alaska varies with latitude. Three distinct types of climate, polar, continental, and marine, are represented along the Alaskan coastline. Polar climatic conditions are characterised by average low ­temperatures less than −10°C and occur along the Arctic

Alaska

1.1

⊡⊡ Fig. 1.1.3 Coastline character of Alaska.

and Chukchi Seas, roughly from Demarcation Point to Point Hope. The Arctic Ocean has minor influence on temperature moderation and precipitation along the coast. The prevailing winds are easterly and are strong along the coast, but decrease inland. Temperatures within this region range from a mean annual high of −2.3°C in Kotzebue to a mean annual low of −15°C along the Arctic Ocean coast. Total average annual snowfall ranges from 760 mm along the Arctic coastline to 1,260 mm in Kotzebue. The continental climate is characterised as wet with cold winters having average temperatures of less than −3°C. Generally, both summer and winter temperatures are extreme and precipitation is light. Except for a few isolated locations, surface winds are light and from the north and east. These conditions occur along the southern Chukchi Sea and Bering Sea, from Point Hope to Unimak Island. Temperatures within this region range from a mean annual

high of 5.6°C at Port Heiden to a mean annual low of −9°C in Wales. Average annual snowfall ranges from 970 mm in Wales to 1,670 mm at Moses Point. The marine climate is characterised by wet, mild winters and cool summers. This type of climate is found along the southern coast, from the Aleutian Islands to the Alexander Archipelago. Temperatures vary little from place to place in southeast Alaska, but precipitation changes based on exposure to the influence of the warm Japan Current and by the proximity of the Gulf of Alaska and by altitude. Strong northerly winds occur along the Gulf coast. Winds are generally from the southeast in the Alexander Archipelago. Most of the annual precipitation is in the form of rainfall, which is heavy for most of the region. Average annual snowfall ranges from 100 mm in Sitka to 6,360 mm at Whittier. Temperatures are moderate with an average annual range of −2° to 6°C in Anchorage and from −4° to 10°C in Sitka.

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Alaska

Tides vary significantly along the Alaskan coast. The coastal tide range generally increases from north to south. The tide range for Barrow, on the Arctic Ocean, is about 0.4 m; that at Port Heiden, on the Bering Sea coast, is about 12 m. High tide ranges are found along the Gulf of Alaska coast and in the Alexander Archipelago. The maximum diurnal tide range for Anchorage is about 9 m and at Whittier it is about 12 m. Sea ice affects most of coastal Alaska. The Arctic Ocean remains frozen roughly from October to June. The ice may reach thicknesses of 200 cm, and 2–3 m pressure ridges are common. The Bering Sea freeze-up generally begins in mid-November. Ice break-up occurs in late May and early June. Ice thickness varies with latitude along the coast. Maximum thicknesses of 161 cm and 82 cm have been calculated for the Kotzebue and Kuskokwim areas, respectively. Bering Sea ice typically extends to the upper portion of Bristol Bay. Sea ice is generally confined to the Cook Inlet along the Gulf of Alaska coast and begins forming in November and is completely broken up by mid-April. Two of the major conditions affecting coastline development in Alaska are permafrost and glaciers. Permafrost is especially significant along the northern coast of Alaska while glaciers are common along the southern coast. Although only the southern part of Alaska and the Aleutian Islands are free of permafrost, the continuous variety (the type that impacts coastal environments) is confined to the coastline north and east of the Bering Sea. In that area permafrost, with its included ice wedges, affects the character and amount of coastal erosion. In the zone of discontinuous permafrost, coastline erosion frequently exposes relict permafrost patches. Along the north coast of Alaska, permafrost that formed during lower sea stands is present in nearshore, sub-sea locations. Except for a few glaciers present in the Brooks Range of northern Alaska, glaciers today are confined to southern Alaska where many still impact the coastline. Most of the Alaska coast, even during the period of maximum glacial extent, was untouched directly by glaciers. Along the Arctic Ocean and Bering Sea coasts, sediment from glaciers (as for example via the Colville River draining the Brooks Range glaciers) reached the sea. During the Pleistocene, glaciers along the Gulf of Alaska coast advanced out over the continental shelf. Today, they are confined to the coastal plain or coastal mountains. Many glaciers in southeastern Alaska have their calving termini in fiords (Molnia 1986; Trabant et al. 2003). In Alaska, geologic and oceanic settings combine to provide five coastal units: Arctic, Bering Sea, Aleutian Islands, Gulf of Alaska, and Southeast. Within these five zones, coastal processes and resultant landforms are

numerous and highly varied. Some forms are small and ephemeral such as sea-ice kettles that develop seasonally along the Arctic coast; some are large and active like the Yukon River delta; some, such as the lava coasts in the Aleutian Islands chain, are the result of catastrophic events; others, like the extensive tidal flats in Cook Inlet, are subjected to daily modification and yet others are large and durable as exampled by Chatham Strait, a 400 km long fiord in southeast Alaska.

2.  The Arctic Coast The Arctic coast of Alaska extends from Demarcation Point on the Canada–Alaska border west to Point Barrow and then southwest to Cape Prince of Wales. Point Barrow, the location of the most northern town in the United States, is also the dividing point between the Beaufort Sea to the east and the Chukchi Sea to the west. From the international border west to Point Barrow, the coast is characterised by low, ice-rich cliffs, except where interrupted by the numerous rivers that flow across the coastal plain. Some of these cliffs face the open sea although most are separated from it by a series of shallow lagoons and low barrier islands. Most cliffs are 1–5 m high although toward the east some are as much as 9 m high (Walker 1991). These cliffs are bordered by narrow beaches. The barrier islands, most numerous along the eastern half of the Beaufort Sea coast, are generally narrow and short and separated by wide inlets. The western half of the Beaufort Sea coast is irregular and is composed of a number of large bays. One large delta, the Colville River delta, and a number of small deltas are present along the coast (Walker 1974). Coastal erosion yields sediment and carbon inputs (Jor­genson and Brown 2005). The Chukchi Sea coast generally has higher relief than the Beaufort Sea coast. Southwest of Point Barrow, tundra cliffs are usually between 10 and 18 m high (Harper 1978) although at Cape Lisburne, where the Brooks Range extends to the shore, they are more than 300 m high (> Fig. 1.1.4). An unvegetated talus slope extends down into the sea. Lengthy spits and barrier islands, separated by narrow inlets, occur along the middle section of the coast between Point Barrow and Cape Lisburne and along the northwest facing portion of Seward Peninsula. Shorter barrier bars are present along much of the rest of this coast. In some places, as at Cape Krusenstern and Point Hope, extensive beachridge systems have developed (> Fig. 1.1.5). Point Hope is a large cuspate formation that has grown out into the Chukchi Sea, flanked by long barriers that enclose lagoons and a swampy coastal plain with rounded lakes. One of

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⊡⊡ Fig. 1.1.4 Cliffs south of Cape Lisburne. (Photo by J.L. LaBelle, Courtesy ARLIS.)

⊡⊡ Fig. 1.1.5 Beach ridges at Point Hope. (Photo by C.D. Evans, Courtesy ARLIS.)

the most distinctive features of the Arctic Coast is a large embayment located north of Seward Peninsula. Sediment transported by the Noatak and Kobuk, two of the largest rivers in Alaska, is rapidly filling this embayment. In arctic Alaska, a low tide range and a long-lasting seaice cover combine to limit the amount of ­wave-generated

coastal erosion. Even during summer, especially along the Beaufort Sea coast, sea-ice is seldom far from shore and effectively reduces fetch and dampens wave action. Nonetheless, coastal retreat can be rapid. When the sea-ice retreats from the shore, usually during summer and fall, storms may cause severe erosion. The two towns of Barrow

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and Wainwright are exposed to a long fetch when the sea-ice retreats from the shore. Nearly all of the arctic coast is within the zone of continuous permafrost and exhibits such permafrost-related forms as ice wedges, ice-wedge polygons, and thaw lakes and permafrost-related processes as thermal melting and thermal erosion. Ice wedges, which may occupy as much as 25% of the upper several metres of the ground, and icewedge polygons, which form a dense network over much of the Arctic Coastal Plain (> Figs. 1.1.6 and >1.1.7), are frequently subjected to wave attack. Differential thaw and erosion result in the undercut, low-cliff, serrated-type coastline that characterises much of the Arctic. Thaw lakes when tapped by ocean waves become part of the shallow foreshore and provide a coastline that is usually arcuate in form. The Cape Lisburne area is dominated by cliffs as much as 300 m high and represents the western end of the Brooks Range. About 50 km south of Cape Lisburne is Point Hope, a 15 km long sand spit. An extensive series of beach ridges has built up at Cape Krusenstern south of Point Hope.

3.  The Bering Sea Coast The Bering Sea Coast extends south from Cape Prince of Wales to the west end of the Alaska Peninsula. It is

dominated by two large embayments (Norton Sound to the north and Bristol Bay to the south) that are separated by the extensive Yukon–Kuskokwim deltaic complex and the southwest extension of the Kuskokwim Mountain Range. Norton Sound, with a relatively smooth coastline, has narrow beaches that are backed by glaciated hills. The Yukon–Kuskokwim delta, a large triangular area with an 800 km long coastline, consists of many lakes, marshes, and abandoned distributaries. Presently, the Yukon River, which contributes the majority of the sediment that reaches the Bering Sea, discharges north into Norton Sound where interdistributary mudflats are extensive. South of the Yukon, the coast is dominated by beach ridges, many of which are undergoing rapid erosion. The Kuskokwim River, in contrast to the Yukonl, flows into a 100 km long estuary. The Yukon–Kuskokwim delta is the seventh largest delta on earth, covering about 54,000 sq km. It is a vast plain with both wet and dry tundra and numerous lakes. Tidal flats, typically 100–1,000 m wide, flank the Kuskokwim delta (Dupre 1977). Although most of the coast between the Yukon and Kuskokwim is made of deltaic deposits, that in the SW part of the deltaic complex, where Quaternary volcanic activity formed hills, includes gravel beaches and rocky headlands. Nelson Island near the Kuskokwim Delta and Nunivak Island about 50 km offshore, have such coastlines.

⊡⊡ Fig. 1.1.6 Ice-wedge polygons during snow melt and break-up.

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⊡⊡ Fig. 1.1.7 Progressive ice-wedge degradation in the Colville River delta.

In contrast, much of the coast of Saint Lawrence Island has low sandy beaches often backed by sand dunes. On its south coast sand and gravel barriers enclose Koozata Lagoon and Sekinak Lagoon, their inner shores having numerous cuspate spits marking early stages in lagoon segmentation. South of the Kuskokwim Estuary, the Kuskokwim Mountains meet the sea with high rocky cliffs and numerous stacks. Most of the Bristol Bay coast is composed of glacial and fluvial sediment and has sandy beaches, extensive mudflats, and lagoons. The tide range along the Bering Sea coast, although variable, in general increases from less than 1 m in Norton Sound to more than 5 m in Bristol Bay. In contrast, the amount of sea-ice decreases southward; ranging from a continuous cover during more than half of the year in the north to only occasional occurrences, even in mid-winter, in the south. Thus, whereas the coastline north of the Kuskokwim is protected by sea-ice during winter, that in Bristol Bay is subjected throughout the year to many of the storms that cross from the Pacific Ocean into the Bering Sea. As a result, most of the unconsolidated coasts in Bristol Bay have been eroded into low cliffs.

4.  The Aleutian Islands The Aleutian Islands begin at the west end of the Alaska Peninsula and extend in a gentle arc for nearly 2,000 km

across the north Pacific Ocean to Attu Island. The islands are the subaerial portions of a 10–100 km wide ridge that flanks the western half of the Aleutian Trench. This trench (over 3,000 km long and up to 7,500 m deep) is the subducting edge of the Pacific Plate. Active subduction is responsible for the occurrence of frequent earthquakes in south Alaska and for the creation and maintenance of the 60 centres of volcanic activity that extend from near Anchorage to the west end of the Aleutian Islands. Of these 60 centres, more than 40 have been active since 1760 (Selkregg 1976). Many of the coasts in the Aleutian Islands are composed of lava that flowed into the sea or of ash that has fallen on the shore or into the water and then washed on shore; other coasts have been eliminated or altered by explosive eruptions. One of the best examples of such a coast is that on Bogoslof Island, which lies near Bering Canyon (the longest submarine canyon in the world). The volcanic history of Bogoslof Island has been followed since 1769 (Shepard and Wanless 1971). The coastlines of most of the islands are irregular. Rocky cliffs alternate with boulder beaches although in some protected coves sandy beaches are present. Abrasion platforms are found at present-day sea level as well as at various heights below and above (up to 200 m) sea level. In the Aleutian Islands, wind-generated waves, which are high throughout the year, combine with high tide ranges to cause extensive coastal erosion. In places, erosion rates are sufficient to have produced stacks along the

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shore and hanging valleys at the mouths of many streams. Other factors affecting coastal development in the islands are the vast quantities of kelp, which grows well in the turbulent water (and diminishes wave energy) and the occasional tsunami some of which have carried driftwood to heights of more than 30 m (Black 1974).

5.  The Gulf of Alaska Coast The Gulf of Alaska Coast, which extends in a clockwise direction from the eastern Aleutian Islands to Cape Spencer, includes a number of distinctive coastal types. The Pacific facing portion of the Alaskan Peninsula, in contrast to its low-lying Bering Sea counterpart, is highly irregular and has numerous rocky headlands that alternate with formerly glaciated inlets. Many of the headlands are represented offshore by rugged islands and rocky stacks. One of the largest islands in Alaska is Kodiak, which is separated from the northern part of the Alaska Peninsula by a 50-km wide strait. The coastline of Kodiak is also irregular and rugged. Fiords, especially along its northwest coast, are numerous. Cook Inlet, a deep structural basin over 300 km long, is a tidal estuary that opens into the Gulf of Alaska and has some of the world’s highest tide ranges. Near its head, extensive tidal flats and marshlands occupy the coastal edge of a low coastal plain. The plain narrows gradually on the northern side of Cook Inlet and toward its mouth merges into mountains that descend directly into the sea. In contrast, on the south side of Cook Inlet, pocket beaches alternate with low cliffs along most of its extent. Associated with the high tide range of Cook Inlet are strong tidal currents; the two help keep ice from forming a solid cover over the Inlet during winter. The coast from the mouth of Cook Inlet east around Prince William Sound is deeply indented with fiords and possesses many ice-scoured headlands and islands. Much sediment is transported into the fiords by the short, turbulent streams that originate in the alpine glaciers that are still present in the coastal mountains. From Prince William Sound southeast to Cape Spencer, the coastline is smooth in outline and is characterised by deltas, beach and dune ridges, outwash plains, moraines, and glaciers. In some locations, tidewater glaciers are present and represent some of the most rapidly changing coastlines to be found in Alaska. One of the major features of this coast is Malaspina Glacier, said to be the world’s largest piedmont glacier (Hayes et al. 1976). Although all of the Pacific Ocean coast of Alaska may experience earthquakes, some of the most severe occur around the Gulf of Alaska and many have a direct effect on

the coastline. The March 27, 1964 Alaskan Earthquake, the most intense ever recorded in North America, depressed, elevated, or otherwise modified some 16,000 km of coastline (Shepard and Wanless 1971). Montague Island was raised by nearly 10 m; other coastlines were lowered, some by as much as 3 m. Homer Spit, in Kachemak Bay, sank nearly 2 m, and spruce forests on low-lying coasts in Turnagain Inlet were replaced by salt marshes. There was slumping on coastal bluffs near Anchorage, forming debris cones that extended up to 600 m into the sea. Such sudden emergence establishes a new zone on which subaerial processes can operate, resulting in a replacement of marine biota, and shifts seaward the zone of marine processes. Submergence, on the other hand, leads to the incursion of saltwater into coastal lakes and marshes, the loss of beaches, spits, and bars, and the subjection of former subaerial surfaces to wave attack. Much of the Gulf of Alaska coast shows evidence of these changes (Waller 1966; Hamilton and Shennan 2005). Earthquakes are a major (although not the only) cause of landslides in the coastal zones of Alaska. Some landslides occur offshore especially at the front of the subaqueous portion of deltas, whereas others happen along the steep subaerial slopes leading down to the shore. The most famous slide was triggered by a July 9, 1958 earthquake near Lituya Bay. A horizontal movement of 6.6 m and vertical movement of 10.7 m caused a rockslide that dumped more than 30 million m3 of rock into the bay. The wave that the rockfall created swept 530 m up the opposite side of the bay destroying all vegetation in the process. Trim­ lines from such waves are conspicuous along its coast. Earthquakes also trigger calving of ice at the terminus of glaciers; for example, during the 1958 earthquake Turner Glacier at the head of the Yakutat Bay was reduced in length by 750 m. Hubbard Glacier is another example of calving glaciers along this shoreline (>Fig. 1.1.8).

6.  The Southeast Alaska Coast The Southeast Alaska coastline extends from Cape Spencer south to Tongass at Dixon Entrance on the Alaska–Canada border. Although the general length of the Southeast Alaska Coast, as given in the United States Coast Pilot (1981), is only 450 km (4.3% of Alaska’s total), the tidal shoreline, is more than 45 times longer. Such a variation is because of the large number of inlets, straits, bays, fiords, and islands that characterise the region. The coast of most of southeast Alaska is rugged with steep cliffs many of which descend far below sea level. However, there are low coastal plains near the head of some of the fiords and

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⊡⊡ Fig. 1.1.8 Hubbard Glacier. (Photo by W.D. Harrisson, Courtesy ARLIS.)

strandflats and reefs border many islands and headlands, especially those exposed directly to the ocean. Many of the geomorphic details are closely related to the intensity of wave and tidal action and lithology. Most of the coastline is protected from the open sea and thus not generally subjected to the high energy storm waves that characterise this part of Alaska. Nonetheless, tides of 7 m and more in some of the fiords and bays are responsible for strong tidal currents and intensive scouring action. In contrast, the gross geomorphic character of southeast Alaska is the result of the glacial modification of a tectonically active, complex mountain terrain. At the maximum of the Plaeistocene Glaciation, ice covered most of southeast Alaska and extended past the present coastline onto the Continental Shelf. During waxing, and to some extent waning, stages, the dominant trend of glacial erosion was controlled by northwest–southeast trending fault structures. Such structural control made possible the formation of exceptionally long fiords. However, during the time when glacial ice dominated the landscape, the flow and therefore erosional trend was seaward across the structural grain. Although today’s glaciers are only minor remnants of their Pleistocene predecessors, they remain important agents in coastal modification. Southeast Alaska, tectonically active since the early Palaeozoic, is presently being uplifted at one of the fastest rates in the world (Skelkregg 1974–1976). In places, uplift is more than 3.5 cm/year, a rate that helps account for the pres­ence of glaciomarine deposits at heights of 230 m above

sea level near Juneau. Some of this uplift is thought to be the result of rebound that occurs with deglaciation. Although general rebound is still occurring, localised rebound is partially responsible for a rapid uplift that is being recorded near retreating glaciers (Hicks and Shofnos 1965).

References Black RF (1974) Geology and ancient Aleuts, Amchitka and Umnak Islands, Aleutians. Arctic Anthropol 11:126–140 Dupre WR (1977) Yukon delta coastal processes study. University of Houston, Houston, TX Hamilton S, Shennan I (2005) Late Holocene relative sea-level changes and the earthquake deformation cycle around upper Cook Inlet, Alaska. Quat Sci Rev 24:1479–1498 Harper JR (1978) Coastal erosion rates along the Chukchi Sea coast near Barrow, Alaska. Arctic 31:428–433 Hayes NO, Ruby CH, Stephen MF, Wilson SJ (1976) Geomorphology of the southern coast of Alaska. In: Proceedings of the 15th Coastal Engineering Conference, Honolulu, HI, pp 1992–2008 Hicks SD, Shofnos W (1965) The determination of land emergence from sea level observations in southeast Alaska. J Geophys Res 70:3315–3320 Jorgenson MT, Brown J (2005) Classification of the Alaskan Beaufort sea coast and estimation of sediment and carbon inputs from coastal erosion. Geo Marine 492:32–45 Molnia BF (1986) Glacial history of the northeastern Gulf of Alaska-a synthesis. In: Hamilton TD et al (eds) Glaciation in Alaska - The geologic record. Alaska Geological Society, pp 219–236 Selkregg LL (ed) (1974–1976) Alaska regional profiles, 6 vols. University of Alaska, Anchorage, AL Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw-Hill Book Company, New York

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Trabant DC, Krimmel RM, Echelmeyer KA, Zirnheld SL, Elsberg DH (2003) The slow advance of a calving glacier; Hubbard Glacier, Alaska, U.S.A. Ann Glaciol 36:45–50 Wahrhaftig C (1965) Physiographic divisions of Alaska. United States Geological Survey Professional Paper 482, Washington, DC Waller RM (1966) Effects of the earthquake of March 27, 1964, in the Homer Area, Alaska. United States Geological Survey Professional Paper 542-D, Washington, DC

Walker HJ (1974) The Colville River and the Beaufort Sea: some inter­ actions. In: Reed JC, Sater JE (eds) The coast and shelf of the Beaufort Sea. The Arctic Institute of North America, Washington, DC, pp 513–540 Walker HJ (1991) Bluff erosion at Barrow and Wainwright, Arctic Alaska. Z Geomorphol 81:53–61

1.2  Washington Maurice Schwartz · Thomas Terich

1.  Introduction The coastline of the state of Washington is 4,296 km long and consists of three segments: the eastern coast and islands from the Canadian border in the north down to the straits and islands of Puget Sound in the south, the northern coast of the Olympic Peninsula along the Strait of Juan de Fuca, and the outer, western, coast facing the Pacific Ocean. Along the eastern side of the Olympic Peninsula is the Hood Canal, a 104 km long inundated glacial trough. An account of the coastal features of Washington State was provided by Shepard and Wanless (1971) and a geomorphic classification of the ocean coast by Terich and Schwartz (1981). Locations of places mentioned can be found in the Washington Atlas and Gazetteer published by the DeLorme Mapping Company, P.O. Box 298, Freeport, Maine 04032, USA http://www.delorme.com. Further topographic de­­ tails are shown on the relevant 1:24,000 Washington 7.5 min maps produced by the United States Geological Survey, Denver, Colorado 80225, USA Reference can be made to the Washington State Department of Ecology collection of air oblique photographs of the coast at http://apps.ecy. wa.gov/shorephotos/. The geology of the Washington State coast is dominated by the folded and faulted Tertiary formations of the Olympic Peninsula and the extensive Pleistocene glacial drift and associated deposits of the Puget Sound region, which extend northward across the Canadian ­border and also occupy valleys and mantle coastal slopes along the Pacific seaboard. There are Mesozoic and Palae­ ozoic formations south of Bellingham and in the San Juan Islands. There was intensive Pleistocene glaciation, producing rugged mountains and wide mantles of glacial drift deposits in the lowlands. Shepard and Wanless (1971) commented that although Ozette Lake (12.8 km long and over 90 m deep) lies in a basin scoured by an ice sheet, the coast between Cape Flattery and Quillayute River shows little evidence of glaciation: there are no fiords. They ­suggested that the ice was relatively stagnant in the coastal fringe, where it melted to deposit a glacial drift mantle.

Cliffs and bluffs along much of the coast and around islands are cut into this glacial drift, and sand and gravel derived from these has produced beaches and spits, especially on the islands, along the eastern coast, and around Puget Sound. Bedrock outcrops, mainly of Mesozoic and Tertiary formations, are found on headlands and along shores where the glacial drift mantle has been removed by marine erosion, and are more prominent on the Pacific Ocean coast. Stacks and natural arches along the Washington coast between the Quillayute River and Point Grenville are the outcome of differential erosion of Eocene (Quinault Formation) volcanic rocks, siltstones and sandstones and the Miocene Hoh Rock Assemblage, which consists of a basic mélange of chaotically mixed blocks of hard sandstone and basalt in a matrix of softer mudrocks and an overlying steeply tilted and overturned formation of sandstone, siltstone and conglomerate. Evidence that the Hoh mélange rose discordantly as diapirs or piercement structures through the overlying sedimentary beds can be seen in sea cliff exposures where the mélange is bounded on both sides by strata steeply dipping outward; and in offshore sub-bottom profiles indicating a doming of otherwise horizontal beds. Due to the easily eroded soft claystone and broken siltstone within it, the Hoh mélange forms embayments along the coast between headlands of the more resistant sedimentary beds of the upper Hoh Rock Assemblage and the Quinalt Formation. Some large blocks of mélange hard sandstone and basalt still remain as erosional remnants in stacks offshore. The coastal climate is cool temperate and humid, with extensive winter snow in the mountains. The Pacific Coast region has an average annual rainfall of about 2,000 mm increasing to 2,875 mm in the rain forest area inland of La Push. Along the forested coastal strip the principal trees are Western hemlock (Tsuga heterophylla), western arborvitae (Thuja plicata), Douglas fir (Pseudotsuga menziesii), Sitka spruce (Picea sitchensis), and Lodgepole pine (Pinus contorta). Westerly winds prevail, but easterlies are frequent, particularly in winter. On the Pacific coast onshore winds are highly variable, between SSW and NNW. East of  the Olympic Mountains rainfall is reduced in a rain shadow

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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region, the Sequim district being well known for its dry climate (annual rainfall down to about 400 mm), but the eastern coast is also relatively wet (1,000–1,200 mm). Tides are typically mesotidal, from about 2 m at neaps to 4 m at springs. Mean spring tide ranges on the eastern coast at Bellingham and Anacortes are 2.2 m, at Seattle 3.1 m and rising to 4.0 m at Olympic in Puget Sound, but on the Strait of Juan de Fuca they diminish to 1.8 m at Port Angeles, and on the Pacific Ocean coast are 2.2 m at Cape Flattery and 2.1 m at Willapa Bay. Never­ theless, tidal currents can be strong through narrow straits, as in Deception Pass, between Fidalgo Island and Whidbey Island, where they attain 18.5 km/h (10 knots), while currents of about 3–7 km/h (1.6–3.8 knots) are measured off Cape Flattery. Wave action along the eastern coast and in Puget Sound is determined by local winds across often short fetches in straits and between islands, but on the Pacific coast westerly ocean swell accompanies locally generated and occasional storm waves. Net longshore drifting varies with configuration (Schwartz et al. 1985): Sandy Point, north of Bellingham, has grown southward but Ediz Hook and Dungeness Spit on the Strait of Juan de Fuca have grown eastward as the result of north-westerly wave action, and the latter has also been shaped by waves from the east and north east. On the Pacific coast the prevailing south-westerly waves generate net longshore drifting northward, as can be seen in the development in that direction of the Long Beach Peninsula and the northerly deflection of the mouths of such rivers as the Copalis, Moclips, Quinault, Queets, and Hoh. There are occasional local reversals, as at La Push, where wave refraction around two large stacks has caused a spit at the mouth of the Quillayute River to grow towards the south. Occasionally tsunamis generated by tectonic disturbances along the nearby plate boundary have produced giant waves on the Pacific coast. The Cascade Subduction Zone is located 85 km offshore, where the Juan de Fuca Plate has been subducting under the North American plate during the last 7,500 years. There is stratigraphic evidence of rhythmic estuarine deposits along the shores of Willapa Bay and Grays Harbour, interpreted as indicating up to six major earthquakes (magnitude 8 or 9) along this zone in the last 3,500 years, with an average return period of 500–540 years (Atwater and Hemphill 1997). At Willapa Bay serial exposures of buried forest and marsh soils, with the uppermost soil mantled by ­tsunami-deposited sand, date the most recent tsunami event as occurring in 1700 (Atwater 1996). Cor­roborating this event are historical records in Japan of a devastating tsunami reaching the coast there on January 26, 1700

(Atwater et al. 2005). This predates European settlement of the Washington coastal area, but in the 1800s coastal native people described nearby Neah Bay as being flooded in the “not very remote past without any swell”, while canoes were lifted into trees and many people were drowned. The Washington coast has been subject to changing land and sea levels, with tectonic uplift in the Olympic peninsula and along the Pacific coast associated with movements on the nearby subducting Juan de Fuca Plate (Atwater 1996) and Quaternary isostatic movements related to land depression by glacial ice loads and recovery and after deglaciations as well as oscillations of sea level. In consequence there are terraces at various levels in the coastal region, active emergence along the shores of Juan de Fuca Strait and the eastern coast and islands, and active submergence south of Cape Flattery and in the southern part of Puget Sound. Cliffs have been cut into glacial drift and bedrock formations on coastal sectors exposed to strong wave action, especially on the Pacific coast, where there are also forested bluffs, usually with some basal cliffing, and these show occasional recession by slumping and landslides, especially during storms and periods of heavy rainfall. The Pacific coast of the Olympic Peninsula is one of the great scenic coasts of the world. Differential erosion of Tertiary continental and marine sediment and volcanics has produced sea caves, stacks, arches and high cliffs. Long term coastline recession is indicated by numerous stacks and islands in coastal waters (mostly within 2 km of the shore), and it is possible that episodes of substantial cliff recession have taken place during occasional tsunamis. Many of the stacks and small islands off the coast south from Cape Flattery are residuals of harder rock, including Eocene volcanics, sandstone and siltstone, the Miocene Hoh Rock Assemblage and the Pliocene Quinault Formation. These are found sporadically within a softer sandstone and mudstone matrix in mélange formations outcropping along the coast, and are exposed and left standing offshore as the softer rocks are removed by marine erosion. Bluffs cut in glacial drift on the islands and along the eastern coast also recede as the result of recurrent local landslides. Shore platforms are not as well developed on the coast of Washington State as on other Pacific coasts (see Australia, New Zealand, Japan), but they are found on the Cape Flattery peninsula (Bird and Schwartz 2000). They include subhorizontal high tide benches produced largely by weathering processes and seaward sloping intertidalsubtidal shore platforms cut by abrasion where waves

Washington

move sand and rock fragments to and fro. Where sediment accumulations exist along the inner edges of shore platforms, they consist mostly of coarse gravel and cobbles; sandy beaches are to be found only in the occasional embayments along this stretch. Similar features are seen on some headlands along the Pacific coast and the Strait of Juan de Fuca, but shore platforms are poorly developed on the eastern coast and around the islands. On the eastern coast and island shores there are sand and gravel beaches, often with cobbles and boulders, derived from glacial drift cliffs and the sea floor. On the shores of the Strait of Juan de Fuca and along the Pacific coast sand and gravel is also supplied by rivers, and heavy driftwood concentrations are found along the shore. As on other coasts dominated by glacial drift deposits (Denmark, for example) there are numerous spits, some straight, some recurved and some cuspate. Dungeness Spit and Ediz Hook, respectively 8 and 5.6 km long, are two large spits on the southern side of the Strait of Juan de Fuca. Damming of a river and barricading the foot of a cliff, both west and updrift of Ediz Hook has caused considerable erosion at the base of this highly industrialised spit. The U.S. Army Corps of Engineers has recently initiated a maintenance program consisting of foreshore revetment and periodic beach nourishment. Smaller spits have formed on island and peninsula shores, particularly at points where the coastline changes direction (Finlayson 2006). On Cape Flattery beaches are confined to coves, but southward along the Pacific coast there are long sand and gravel beaches between rocky headlands. South of Moclips the coast consists of broad, prograded sandy beaches fronting dunes. Provenance of the accreted sand along this sector has been attributed to the Columbia River to the south, whence sand has been carried northward by longshore drifting. With the construction of large dams on the Columbia during the twentieth century, it is speculated that the progradation regime may have ceased and possibly will turn to one of erosion. The beach that prograded alongside the breakwater built on the northern side of the Columbia River mouth at Cape Disappointment is now eroding, and erosion has followed earlier progradation alongside breakwaters at Point Brown and Westport, north and south of the entrance to Grays Harbour. Coastal dunes are best developed on the Long Beach Peninsula, where there are multiple rows of beach and dune ridges, each of which was built parallel to the prograding sandy coastline. Locally there are blowouts and small transgressive dunes, as on Leadbetter Point at the northern end of North Beach Peninsula, but the extensive dune fields of the Oregon coast have not developed here.

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Similar dune morphology has developed along the coast south of Copalis Beach to Ocean Shores and Point Brown, and between Westport and Cape Shoalwater. Otherwise there are only minor dunes, chiefly behind beaches near the mouths of rivers and on some of the spits in northern Washington. One small dune at Fort Worden State Park, near Port Townsend, has been planted with dune grass to maintain it in an intensively used area. The southern part of the Pacific coast is backed by a gently rolling plain, with two large estuarine lagoons, Willapa Bay and Grays Harbour, developed by the submergence of the mouths of valleys cut by rivers during phases of lower sea level, followed by the growth of barrier spits on the ocean shore, constricting their entrances. Northward migration of the main Willapa Bay inlet channel is believed to be the cause of severe beach erosion on the northern shore, at Cape Shoalwater, where the coastline has receded up to 3.7 km over a 76 year period. Reference has been made to the breakwaters built to stabilise the entrance to Grays Harbour. There are small estuaries at the mouths of rivers on the Pacific coast, notably the Quillayute and Queets, and the Pysht River on the Strait of Juan de Fuca, and the growth of the compound Dungeness Spit has almost enclosed the waters of Dungeness Bay as a coastal lagoon. Crockett Lake on Whidbey Island is an example of a coastal lagoon enclosed by a wave-built barrier of sand and gravel. Several deltas have been built on the eastern coast where rivers have deposited large loads of sediment brought down from mountainous catchments. They include the deltas of the Lummi and Nooksack, the Samish, Skagit and Stillaguamish deltas, the Snohomish delta near Everett, the much modified Puyallup delta (largely converted to Tacoma docks) and the Nisqually delta at the southern end of Puget Sound. The Skohomish River has built a marshy delta at the Great Bend on Hood Canal. On the Strait of Juan de Fuca coast the Elwha delta is prominent. Marshes and swamps (sloughs) have formed on the more sheltered parts of estuary and lagoon shores, and in straits and bays protected from strong wave action by ­peninsulas, islands or spits. There are corridors of fresh water swamp between the north-south dune ridges on Long Beach Peninsula, some of which are used to cultivate cranberries.

2. The Coast of Washington from North to South Point Roberts is a rectangular peninsula, an isolated segment of the United States, cut off because the Canadian

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boundary, the 49th parallel of latitude, runs westward across Boundary Bay to intersect it between Maple Beach and Boundary Bluff. It has beaches of sand and gravel bordered by multiple sand bars. Drayton Harbour is a NW-facing horseshoe-shaped bay bordered on the western side by steep wooded bluffs, and partly enclosed as a lagoon by the long Semiahmoo spit that has grown north-east toward Tongue Point opposite the harbour structure at the town of Blaine. This spit shelters salt marshes and mudflats exposed at low tide in Drayton Harbour on its landward (eastern) side, while the more exposed western shore consists of a beach of well-rounded pebbles and sand that continues southward below wooded bluffs and cliffs 20–25 m high, with occasional landslides in grey clayey sand (glacial drift). At the southern end of Semiahmoo spit, adjacent to the bluff and beside a sewage treatment site, is a large Indian midden over 6 m thick, dating from 4,000 bp onwards and rather overgrown. The beach passes in front of slumping bluffs south towards Birch Point (>Fig. 1.2.1). Birch Bay is similar in outline to Drayton Harbour, but faces SW and has no enclosing spits. Mean spring tide range is about 3.6 m, and at low tide extensive sand bars are exposed in front of a gravel beach segmented by groynes. The beach outline is an asymmetrical (log-spiral) curve, produced by the refraction of waves from the NW, and although there are episodes of northward and southward drifting net longshore drift is zero. A succession of emerged coastlines in the hinterland of Birch Bay, marked

by a stairway of glacial rebound terraces, follows the same asymmetrical curve. South of Neptune Beach is Sandy Point, a large spit that has grown southward as the result of wave action sweeping sand and gravel along the coast (>Fig. 1.2.2). In the 1960s canals were dredged within the spit to form a marina, and there are now many houses behind the beach and around the boat channels. Unfortunately, beach erosion has become a serious problem, and residents have built sea walls and boulder ramparts on the shore to protect their properties. The spit is low-lying, and occasionally overwashed and flooded during storms. In the 1920s it was temporarily breached, and the storm of Easter 1977 swept sand and gravel in from the beach and damaged several houses. Within Lummi Bay, to the south, are the eroded (and now embanked) remains of the Lummi delta, which decayed after the diversion of the Nooksack River eastward into Bellingham Bay in 1877 (see below). Portage Point is at the south-eastern end of the Lummi Nation Indian Reservation peninsula. The Portage (>Fig. 1.2.3) is a tidal divide (intertidal tombolo) where a road is exposed at low tide, giving access to Portage Island. Brant Point is a sandy spit with shoals curving out in the lee of Portage Island into Bellingham Bay. There is a cliff cut in glacial drift on the southern end of Portage Island at Point Frances (>Fig. 1.2.4) which is typical of the intermittently active cliffs in this region. On the northern side of Bellingham Bay the Nooksack River has a growing delta, and the bay water is discoloured

⊡⊡ Fig. 1.2.1 Slumping cliff in glacial drift north of Birch Point. (Courtesy Geostudies.)

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⊡⊡ Fig. 1.2.2 Air view of Neptune Beach and the spit that has grown southward at Sandy Point. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.3 The intertidal tombolo linking Portage Island to the Lummi Peninsula. (Courtesy Washington Department of Ecology.)

by brown silt and clay when the river floods. The Nooksack River has a northern branch, Lummi River, which flows to the sea in Lummi Bay on the northern side of the peninsula. In the mid-nineteenth century the Nooksack had become choked with a massive log jam just below the divergence of the two rivers, and some of the water and sediment discharge was carried down Lummi River into Lummi Bay. In 1876–1877 the log jam was removed, and the Lummi River diked off. After that the Nooksack delta began to grow out into Bellingham Bay (Shepard and Wanless 1971). The San Juan Islands are scattered, mainly forested islands of various sizes and shapes in the southern part of the Strait of Georgia. They represent the partial marine submergence of a mountain range, dissected by glaciated

valleys that are now fiords. San Juan Island is the largest of the group, about 29 km long with an average width of 10 km, a hilly island rising to the grassy granite slopes of Mount Young, 198 m high. It has steep coasts and rocky shores, declining to a narrow peninsula in the south-east, extending to Cattle Point. On the southern coast False Bay is a shallow rounded embayment with a wide intertidal zone. To the north of San Juan Island is Waldron Island, consisting of a south-eastern upland of Upper Cretaceous shaly sandstones and conglomerate culminating southward in the steep-sided narrow Point Disney, which has vertical cliffs up to 150 m high on its western flank. On the eastern coast is rocky Mail Bay, and to the north a lowland of glacial drift rising to 30 m above sea level. Stratified glacial

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⊡⊡ Fig. 1.2.4 The cliff at Point Frances on the southern coast of Portage Island. (Courtesy Washington Department of Ecology.)

deposits outcrop in 30 m cliffs near Point Hammond at the north-eastern end of the island, and there are several small promontories of exposed bedrock, the most prominent at Fishery Point in the north-west. From here North Bay curves out to a sharp cuspate point of glacial drift projecting westward to Sandy Point. To the south Cowlitz Bay is another curved bay that extends south to Point Disney, and near the south-eastern end is intertidal Mouatt Reef, rising a metre above low tide. To the north of San Juan Island is Spieden Island, a narrow ridge rising to 116 m. Nearby is Flattop Island, which has a steep rocky north-west coast and a surface that is not flat, but slopes south-east with the dip of the rock formations. Gull Rock, off the north-west point of the island, is split by a chasm cut out of shale between hard conglomerates. The hard rocks form reefs, notably White Rocks, one of which rises 10 m above high tide level. To the west is Stuart Island, almost split by elongated Reid Harbour, with Turn Point at its western end looking across Haro Strait to the similar Canadian islands that border Vancouver Island. Also to the north of Waldron Island is Bare Island, a small island and seabird colony rising 12 m above sea level and almost unvegetated. Narrow inlets have been cut in steeply dipping shales on the east and west coasts, and there is a cap of glacial drift as well as much guano. Orcas Island is M-shaped, with a smooth ­north-eastern coast following the geological strike and to the south

s­ teep-sided ridges separated by deep inlets, East Sound, West Sound and Deer Harbour. Mount Constitution, the highest peak in the San Juan Islands, rises to 759 m. Point Doughty is low and narrow, projecting into President Channel, and Freeman Island is a small island to the south that has a broad flat rock platform exposed at low tide. There are broad flat rocky reefs off the north coast, including Parker Reef, exposed at low tide. Lopez Island is a lowlying island to the south of Orcas Island. On the mainland coast Bellingham Bay is a broad bay opening southward and bordered by steep bluffs and occasional cliffs. Relics of the wooden supports for the Bellingham and British Columbia Railroad stand offshore along the SE coast, and at Bellingham the Georgia Pacific water treatment lagoon is prominent on the shore. Aerating sprinklers and introduced bacteria and fungi were used to break down chemicals generated by the paper works and discharged into the sea, but this reduction of pollution in Belling­­­ ham  Bay resulted in the return of the shipworm Teredo, which bored into the wooden pilings that support piers and attacked the rafts of logs that used to be floated in by sea. Fairhaven is a historic neighborhood, dating from the mid-nineteenth century. The coast has been modified by reclamation. Towards the end of the nineteenth century large quantities of sand and gravel were sluiced from Poe’s (Post) Hill Point, which was an 18 m high promontory, and deposited in the bay to form new land that was then called Commercial Point for the development of

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shipbuilding yards, boat piers and factories. Depletion of the beach at Marine Park, Fairhaven, was countered in 2004 by emplacing an artificial beach of shingle. South of Fairhaven Chuckanut Mountain (>Fig. 1.2.5) descends to a steep forested coast on Cretaceous sandstones, passing southward across the Oyster Creek Fault to Palaeozoic metasediments. Below the steep slopes mudflats are exposed at low tide on the shores of Samish Bay. The view south from Chuckanut Mountain is of a wide coastal plain which includes the Samish River delta (>Fig. 1.2.6). ⊡⊡ Fig. 1.2.5 Steep coast on Cretaceous sandstones in Larrabee State Park, below Chuckanut Mountain south of Fairhaven. Honeycomb weathering is well developed on these feldspathic sandstones (Mustoe 1982). (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.6 Coastal plain formed by reclamation of swamps and intertidal areas on the Samish delta. (Courtesy Geostudies.)

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Samish Island, a wooded flat-topped ridge (two former islands linked by a sandy tombolo), was attached to the mainland by reclamation of intertidal land between earthen dikes by Dutchmen more than a century ago. Offshore is a group of islands including high Cypress Island and low-lying Guemes Island. Sand and gravel eroded from cliffs has supplied beaches and spits shaped by wave action on the shore of Lopez Island (>Fig. 1.2.7). South of the Guemes Channel is Fidalgo Island, the inner part of which is separated from the mainland only

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by a marshy isthmus followed by the Swinomish Channel. Shannon Point is a cliffed promontory at the north-­western end with glacial drift (Vashon drift (c. 13,000–15,000 years bp.) of the Fraser glaciation) over strongly folded and metamorphosed Mesozoic sediment and volcanics and a crystalline basement of Devonian rocks. Fidalgo Head on the west coast is a grassy bluff with outcrops of serpentine rock, above the swirling waters of Burrows Channel, looking across to the San Juan Islands. To the south is Deception Pass (>Fig. 1.2.8).

Deception Pass is located between Fidalgo Island and Whidbey Island and is a deep and narrow strait through which the tides swirl at up to 18.5 km/h (10 knots). It is bordered by steep rocky cliffs. When it was discovered by Vancouver in 1792 he thought it would lead through to a harbour, but when he found otherwise he named it from having been deceived. West Point (>Fig. 1.2.9) has been truncated by strong tidal currents. Whidbey Island is about 145 km long, parallel to the mainland coast and extending southward to Seattle. On the ⊡⊡ Fig. 1.2.7 Cuspate spit, Flat Point, Lopez Island, San Juan Islands, with Shaw Island on the right. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.8 Deception Pass, showing strong tidal current. (Courtesy Geostudies.)

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west coast high cliffs of sandy glacial drift face the Strait of Juan de Fuca, with beaches and a looped barrier (>Fig.  1.2.10). Point Partridge and Admiralty Head are prominent headlands, and on the latter is Fort Casey, where an old searchlight tramway protrudes 4  m from the cliff below an artillery fort, indicating recent recession of the cliff. This part of Whidbey Island has the Weird Pits, which are kettle holes up to 30 m deep which formed as the result of collapse following the melting of chunks of ice left in the glacial drift after the glaciers receded. Near Point Partridge ⊡⊡ Fig. 1.2.9 West Point, Deception Pass, with a strong ebb current flowing past the rocky outcrop and driftwood on the beach. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.10 Ebey’s Landing on the west coast of Whidbey Island has a looped barrier of sand and gravel enclosing Parego’s Lagoon, backed by bluffs 80 m high. (Courtesy Geostudies.)

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one of these has been excavated by the sea to form a cove with a little lagoon. On the east coast of Whidbey Island a succession of promontories separates coves and harbour bays. Bluffs with a scrub or forest cover alternate with cliffs exposing glacial drift deposits, and there is occasional slumping in the bluffs. At low tide a gravelly foreshore is exposed at the village of Coupeville, founded in 1852 on slopes descending into Penn Cove. Skagit River has built a delta on which the river splits into North Fork and South Fork, on either

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side of Fir Island, and there are relics of many distributary channels. Camano Island, parallel to Whidbey Island east of Saratoga Passage, has wooded bluffs and occasional cliffs. Camano Head (100 m) is one of the tallest, steeply sloping bluffs in the archipelago, with tanglewood scrub sliding down the slope. South of Stanwood the Stillaguamish River has built a delta into Port Susan, a bay sheltered from westerly winds and waves by bluffs in the lee of Camano Island. The coast southward along the Tulalip Indian Reservation has wooded bluffs and sand and gravel beaches, with some minor spits. Mission Point is a hooked sand spit sheltering Tulalip Bay. A delta has been built where the Snohomish River and Steamboat Slough curve out on either side of Smith Island, north of the town of Everett. This town is built on low hills beside the valley of the Snohomish River. The waterfront harbour on the shores of Port Gardner Bay is sheltered by a chain of low islands extending north to Jetty Island, partly artificially formed by the dumping of spoil from dredging. Reclamation of the Everett waterfront has formed a series of docks and marinas and a large U.S. naval station at the southern end. Puget Sound was cut by glaciers and meltwaters which formed deep, steep-sided channels, into which the sea flooded to form marine inlets that wind and branch deep into the state of Washington; inlets bordered by forested bluffs. The islands are generally steep-sided and flat or gently undulating, possibly as the result of marine

­ lanation of the glacial drift deposits at higher sea levels. p The rise and fall of the tides in Puget Sound is about 2.1 m in the north, increasing to nearly 3 m in the south, and at low tide sandy and gravelly beaches and muddy fringes are exposed. Spits are found on many of the islands (>Fig. 1.2.11), and there are intertidal and subtidal shoals. South from Everett bluffs and cliffs form the eastern coast of Possession Sound, and extend past Mukilteo to Norma Beach, where the coastal suburbs of Seattle begin. Bluffs and small headlands with some coves and beaches line the Seattle coast. Seattle has spread across the isthmus of ridges and valleys that trend south to north on the eastern side of Puget Sound, has a coastline of wooded bluffs with headlands and bays, and some of which have been reclaimed for port development. West Point is a cuspate spit that carries a City of Seattle sewage treatment plant, a small Metro lighthouse marking the northern point of Elliot Bay, and a Coast Guard building. Elliot Bay provided a good harbour in deep water, below a steep wooded hillside with ravines and springs. This part of the coast has been much modified by land reclamation; in 1902–1930 Denny Hill was removed and the earth and rock sluiced from it was dumped to reclaim former tideflats in Elliot Bay. Tacoma is a large city on a wide plateau near the southern end of Puget Sound. Commencement Bay is broad, with the delta of Puyallop River at its head, modified to form Tacoma docks. The coast southward is dominated by ⊡⊡ Fig. 1.2.11 Trailing spit extending south east from Smith Island in the Strait of Juan de Fuca. (Courtesy Geostudies.)

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steep bluffs cut in glacial drift, extending past the narrow promontory at Point Defiance. Towards the southern end of Puget Sound the Nisqually River has built a shrubby deltaic plain with extensive marshes. The head of Budd Inlet, to the west, has been enclosed to form Capitol Lake, alongside Olympia. Bainbridge Island, on the western side of Puget Sound opposite Seattle, is a broad plateau of glacial drift bordered by bluffs, separated from the mainland by Agate Passage. Restoration Point has an emerged shore platform 6 m above the modern one, probably uplifted tectonically. Point Monroe has a looped sand barrier enclosing a lagoon at the northern end of Bainbridge Island. To the south is Vashon Island, similar to Bainbridge Island, attached to Maury Island by way of an isthmus at Portage. On the west coast of Puget Sound opposite the cliffy southern end of Whidbey Island is Point No Point, a rounded or lobate spit like many others on Puget Sound, shaped by waves delivering sand and gravel but blunted by strong tidal currents. Hood Canal, the long channel that diverges from Admiralty Inlet near Port Ludlow is generally 3–5 km wide as it runs southward past Hoodsport to the Great Bend, and 1.5–3 km wide as it swings north-east to its marshy head at Lynch Cove, near Belfair; it is bordered by steep wooded slopes that are sometimes disrupted by landslides, the scars of which persist for some years. Inflowing streams have built small deltas, such as that of the Dosewallips River, and a large delta has been built by the Skokomish River south of the Great Bend near Union. Longbranch Peninsula, between Case Inlet and Carr Inlet, has springs flowing from cliffs of gravel and clay on its west coast. The Hood Canal opens northward into Admiralty Inlet near Port Ludlow, sited on a slight hill overlooking a harbour. To the north, Marrowstone Island and Indian Island are islands of glacial drift that run north-south. They are fringed by bluffs and minor cliffs, and have beaches of sand and gravel and spits, such as the cuspate spit at Marrowstone Point. Port Townsend on the east coast of the Quimper Peninsula, looks across Admiralty Inlet to Whidbey Island, and has bluffs fronted by a reclaimed shore area (Zelo et al. 2000). Beach sand drifts northward past Fort Worden. Point Wilson at the northern end of Quimper Pen­ insula, on the western side of the entrance to Admiralty Inlet and Puget Sound, is a sand and gravel spit with low dunes. It has been shaped by waves arriving from the south-east, across Admiralty Inlet, and from the west, along the Strait of Juan de Fuca. The Quimper Peninsula, bordered to the west by Discovery Bay is broad and flat-topped with steep forested

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bluffs. At Fairmont, a village at the head of Discovery Bay, salt marshes and mudflats are exposed as the tide falls. Offshore is Protection Island, which became a national wildlife refuge in 1982. West of Discovery Bay is the broad Miller Peninsula, and then Sequim Bay, which is essentially a lagoon almost enclosed by Gibson Spit (>Fig.  1.2.12) which has grown southward on the western side and Travis spit, which has grown eastward. Strong tidal currents flow through the gap. Dungeness Spit (>Fig. 1.2.13) is a sand and gravel spit nearly 10 km long, a long narrow ridge of sand and gravel, piled with driftwood, with lobes of washed-over sand, runs out eastward to the distant lighthouse (Schwartz et al. 1987). Sediment from eroding cliffs in glacial drift to the west has drifted alongshore to build the spit, which is still growing eastward. It is backed by a lagoon, Dunge­ness Bay, containing truncated Cline spit and broader Graveyard Spit, running sharply back from the middle of Dungeness Spit, and intervening marshes and lagoons (>Fig. 1.2.14). Dungeness Spit has been supplied with sand and gravel from eroding cliffs in glacial drift to the west. The steep cliffs cut in sand and gravel between Dungeness Spit and Port Angeles are receding, and pose a problem for houses built on cliff-top land. Cliff recession occurs as the result of basal wave attack, but there is also slumping, particularly after wet weather when the glacial drift deposits become overloaded with groundwater and are unstable. Port Angeles stands in the lee of the Ediz Hook spit. Ediz Hook is a spit 5.6 km long that has grown out from the coast east of the Elwha delta, supplied with sand and gravel from eroding cliffs and from the Elwha River. In recent years erosion has become severe along the shores of this spit, partly because of the reduction of sand and gravel supply from cliffs that have been stabilised and partly because of the damming of the Elwha River, so that the shore is now heavily armored with large rock boulders, which preserve the spit artificially as a breakwater for Port Angeles. The Elwha River is incised in a deep valley. It has built a delta, but this is eroding because dams have intercepted most of the sediment previously brought down the river to the coast. West from the mouth of Elwha River the coast consists of bluffs and cliffs, interrupted at the mouths of many small river valleys. Pysht is at the mouth of two valleys, and to the west there is a rocky headland and a prominent sea stack to the west on Pillar Point. Clallam Bay is a village with a small fishing harbour, where a walkway leads out from the coastguard station to a steep headland and Sekiu is a village on the western shore of Clallam Bay, with a sheltered harbour. To the west are cliffs fronted by shore platforms exposed at low tide.

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⊡⊡ Fig. 1.2.12 Gibson Spit (left) and Lagoon at Washington Harbour and Travis Spit (right) on Miller Peninsula constrict the entrance to Sequim Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.13 Dungeness Spit. (Courtesy Journal of Coastal Research.)

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⊡⊡ Fig. 1.2.14 The eastern end of Dungeness Spit, showing the broad Graveyard Spit and the narrower Cline Spit built by easterly wave action. Directions of longshore drifting are indicated. Strong tidal currents flow through the constricted passage in and out of Dungeness Bay, to the left: they attain 56 cm/s as the tide rises and 60 cm/s (about 1 knot) as it ebbs, and the channel floor has a gravel lag deposit, in contrast with mud under the lagoon and sand near the shores. Dungeness River, on the right, reaches the sea by way of an outlet deflected westward by longshore spit growth. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.15 Old Hat island at Seal Rocks, east of Neah Bay, Strait of Juan de Fuca. (Courtesy Geostudies.)

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Seal Rocks, east of Neah Bay, include high Old Hat islands (>Fig.  1.2.15), surrounded by subhorizontal high tide shore platforms produced largely by weathering processes (Bird and Schwartz 2000). Neah Bay, to the east of Cape Flattery, has a harbour protected by a breakwater linked to Waadah Island, and has a crescent-shaped beach of pure white sand connecting Bahada Point. Tatoosh Island (>Fig.  1.2.16) is a flat-topped cliffedged island offshore, the northwesternmost point of the 48 coterminous United States at the western end of the Strait of Juan de Fuca. Cape Flattery has massive sandstone cliffs with basal rock ledges, looking out to Tatoosh Island. The peninsula consists of Lower Tertiary (mainly Eocene) stratified sandstones and siltstones and, locally massive or conglomeratic. The cliffs show weathering features, including tafoni and caves, structural ledges, and basal subhorizontal high tide shore platforms (Bird and Schwartz 2000). High cliffs run south from Cape Flattery towards Waatch Point, where there are wide rocky shore platforms cut by abrasion as waves move sand and gravel to and fro. Hobuck Beach at the mouth of Waatch River curves round Mukkaw Bay towards Anderson Point. Sooes River flows northward behind a barrier spit to a deflected outlet. Anderson Point has subhorizontal high tide shore platforms around islands and rock stacks. Shishi Beach, 4.8 km of sandy shore, runs south from Portage Head in the Makah Indian Reservation; there are numerous grey

jagged reefs and stacks offshore. Point of the Arches is a cliffy headland with caves and tunnels and many stacks offshore. Between Point of the Arches and Cape Alava is wilderness coast in the Olympic National Park, accessible only to walkers. Cape Alava is the most westerly point in the coterminous United States. South from Cape Alava, near Ozette, the coastline is a series of beaches a few ­hundred metres long with a headland at each end. The headlands are 30–50 m high, steep and thick with salal bush, and impassable at high tide. Alava Island stands as a mesa in the ocean, first and biggest of a group of sea stacks, reefs, isles and boulders known as The Flattery Rocks. One of the first headlands south of Cape Alava is Wedding Rock, a dark-basalt cliff with broken rocks at its base. These rocks have served as a canvas for ancient artists, with petroglyphs carved on the rocks. Sand Point, down a 15 km walking trail from Ozette, is notable for its shore pools exposed when the tide is out. Ozette Lake is elongated parallel to the coastline, and one report suggests that it was impounded by a bouldery glacial moraine on its western shore (Alt and Hyndman 1984). Its great depth (90 m) indicates that it occupies an ice-scoured depression. The coast continues southward past Point Johnson and several other rocky promontories. Hole in the Wall is on a small headland penetrated by a wave-cut cave. There are many outlying sea stacks of dark basalt. Rialto Beach, a curving bank of pebbles and cobbles backed by heaps of driftwood worn smooth and clean by the abrasive waves, runs out from the mouth of Quillayute River at La Push. The harbour is ­protected by a large ⊡⊡ Fig. 1.2.16 An air view of Tatoosh Island, looking towards Cape Flattery and showing high-tide shore platforms.

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breakwater, strewn with driftwood, linked to one of several forested flat-topped islands. To the south are First, Second and Third Beaches, extending down to rugged Teahwhit Head. 27 km of steep forested wilderness coast extends between La Push and Ruby Beach, which takes its name from the polished black stones that can be found on the shore. This part of the coast is accessible only on foot, along Olympic National Park trails that run behind prominent Hoh Head and a succession of rocky capes and curving bays, with many islands and stacks offshore (>Fig. 1.2.17). Destruction Island is a bold, steep-sided island, 1.5 km long and 5 km offshore (Shepard and Wanless 1971), one of the very many islands and stacks off the oceanic Washington coast. To the south Kalaloch Creek swirls out past a slumping meander cliff to the sea. Southward the coastline straightens, and is fringed by a beach of sand and gravel. At Queets the beaches are again numbered rather than named: First Beach, Second Beach and Third Beach. Tunnel Island is penetrated by a large cave. There have been recurrent tsunamis on this coast. Cape Elizabeth is a flat-topped promontory with cliffs cut in soft stratified siltstone and conglomerate (Quinault Formation) which run southward to the mouth of Quinault River. A 1902 photograph shows that a double arch in Eocene sandstone and conglomerate existed here, but before 1970 the inner arch collapsed. South of Taholah is Point Grenville, a rocky promontory of marine volcanic ⊡⊡ Fig. 1.2.17 Hoh Head, a cliffed promontory of hard sandstone that has persisted while adjacent sectors of weaker mélange have been cut back. (Courtesy Washington Department of Ecology.)

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rocks, rising abruptly from the ocean, into which it extends a short distance in a semicircular shape. Outlying to the southeast are two pyramidal rocks 20–25 m high. Moclips is a village on top of bluffs with pine and cedar forest at the mouth of Moclips River. The straight coast ­continues, with bluffs and cliffs rising behind a long sandy beach. Roosevelt Beach is a firm sandy surf beach in front of cliffs of soft dark sand. The barrier beaches continue southward to the mouth of the Columbia River. Roosevelt Beach narrows southwards past Copalis River to Copalis Beach, where the steep wooded bluffs, locally slumping, give place to dunes behind a narrow shingle beach with much driftwood. At Sampson the beach and dunes become a barrier spit on the northern side of Grays Harbour. On it are extensive seaside resorts, including Ocean Shores. Grays Harbour is a large estuarine lagoon which narrows eastward to the mouth of Chehalis River, where Aberdeen and Hoquiam are log-shipping and lumber-milling towns. At Point Brown, on the northern shore, North Jetty was built to stabilise the entrance. A wide sandy beach, backed by dunes, built up after the jetty was first built, but is now eroding away, and at Ocean Shores a large condominium block which was threatened by shore erosion and protected by massive boulder rampart now projects on to the beach. Westport on the broad sandy spit that borders the southern part of Grays Harbour stands behind the sand dunes, with long jetties built early in the twentieth century

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to stabilise the entrance to Grays Harbour and disperse the sand shoals that impeded navigation: at first the jetties led to widening of the sandy beach on either side of the entrance, but in the last few years there has been rapid erosion on the southern shore, and the Corps of Engineers has had to deposit about 600,000 cubic metres of sand and pebbles to make a wide isthmus that prevents the sea breaking through into Half Moon Bay and threatening Westport town. Cape Shoalwater is a low promontory on the northern side of the entrance to Willapa Bay, and has been subject to severe marine erosion. Comparison of historical maps shows that Cape Shoalwater was been cut back 3.7 km over a 76 year period, and that since 1967 it has been receding about 30 m/year. Since the early charts were made the entrance to Willapa Bay has widened, and the erosion is partly due to the deep-water channel impinging on the Cape Shoalwater shore (>Fig. 1.2.18). Willapa Bay, once known as Shoalwater Bay, is a large coastal lagoon with a wide shoaly entrance, an irregular outline, rocky headlands and marshy coves, and the high forested Long Island; the average tide rises and falls about 2 m. It has been calculated that Willapa Bay exchanges 12 million gallons of water per tide: on the west coast only the Columbia River, San Francisco Bay, and Puget Sound exchange more than this. Long Beach Peninsula is a barrier spit, consisting of sand brought down to the sea by the Columbia River, which has accumulated on the coast north of the river

mouth. A wide beach is backed by grassy and forested ­parallel dune ridges, with Willapa Bay on the landward side, the dunes running from south to north, and separated by elongated hollows containing lakes and bogs. Long Beach extends north from Cape Disappointment to Leadbetter Point at the entrance to Willapa Bay. South of Seaview the dunes come to an end, interrupted by cliffy North Head and the rocky Cape Disappointment promontory. From Cape Disappointment there is a view of the long North Jetty, with much driftwood on the sandy beach that has grown very wide on its northern side. The breakwater was built to stabilise the northern side of the mouth of the Columbia River, and comparison of early maps and charts with the modern configuration indicates that a wide beach has formed on its northern side. Nearby is a deep, steepsided rocky inlet. Large quantities of sand were carried down the Columbia River and swept into the sea in times of flood, and ocean waves have carried this onshore to build the beaches and dunes of North Beach to the north (and Clatsop Plains, in Oregon to the south). In recent decades the sand supply has diminished, possibly because of dam construction upstream, and beaches that were formerly built up with sand washed in from the sea floor now show  signs of erosion. The explosive eruption of Mount St. Helens in 1980 sent large quantities of sediment down the Toutle River into the Cowlitz River and so into the Columbia River at Longville. The sediment reaching the

⊡⊡ Fig. 1.2.18 Coastline erosion at Cape Shoalwater. (Courtesy Geostudies.)

Washington

Columbia River was largely fine-grained (silt and clay) rather than sand. At Longville the river shallowed overnight from, and dredging was necessary to 12 to 4 m restore navigability. Turbid water flowed downstream to the mouth of the Columbia River, but there was no sand accretion as a result of the Mount St. Helens eruption. Along the northern side of the Columbia River steep slopes descend to the estuary shore. There is a succession of bays and rounded promontories, and Grays River and a number of smaller streams flow into the Columbia estuary from the north. Upstream the estuary becomes shoaly, with marshy islands extending at intervals up to the Lewis and Clark bridge at Longview.

References Alt D, Hyndman D (1984) Roadside geology of Washington. Mountain Press Atwater BF (1987) Evidence for great Holocene earthquakes along the outer coast of Washington State. Science 236:942–944

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Atwater BF (1996) Coastal evidence for great earthquakes in western Washington. In Rogers, AM et al. United States Geological Survey Professional Paper 1560:77–90 Atwater BF, Hemphill Haley E (1997) Recurrence intervals for great earthquakes of the past 3500 years at Northeastern Willapa Bay. Washington. United States Geological Survey Professional Paper 1576 Atwater BF et al (2005) The orphan tsunami of 1700: Japanese clues to a parent earthquake in North America. United States Geological Survey Report 1707 Bird E, Schwartz M (2000) Shore platforms at Cape Flattery. Washington. Washington Geology, 28: 10–15 Finlayson DP (2006) The geomorphology of Puget sound beaches. Dissertation, School of Oceanography, University of Washington, Seattle, WA, 216 p Mustoe GE (1982) The origin of honeycomb weathering. Bull Geol Soc Am 93:108–115 Schwartz ML, Fabbri P, Wallace RS (1987) Geomorphology of Dungeness Spit, Washington, U.S.A., J Coastal Res 3:451–455 Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw-Hill, New York Terich TA, Schwartz ML (1981) A geomorphic classification of Washington State’s Pacific Coast. Shore Beach 49:21–27 Zelo I, Shipman H, Brennan J (2000) Alternative bank protection methods for Puget Sound shorelines. Washington State Department of Ecology publication #00-06-012

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1.3 Oregon

Paul D. Komar

1.  Introduction

2.  The Oregon Coast

The Pacific coast of Oregon extends for nearly 700 km from the Washington border on the Columbia River in the north, to the border with California in the south (>Fig. 1.3.1). The Oregon coast is characterised by considerable geomorphic variability, with long stretches of rocky shore and major headlands formed of resistant volcanic rocks, which isolate sections of sandy beaches ranging in lengths from small pockets to littoral cells that extend for 10 s to over 100 km distances along the coast; the longest is the 250 km beach from Florence to Coos Bay, backed by the Oregon Dunes Recreation Area, the most extensive coastal dune sheet in North America.

Important to the geology and geomorphology of the Oregon coast is its tectonic setting, being located within the zone of collision of three of Earth’s tectonic plates, the oceanic Juan de Fuca and Gorda plates, and the continental North American plate. The crust of the ocean plates is formed at the spreading ridges and is then carried eastward toward the continent, where being denser the ocean crust slides beneath the less-dense continental crust and is subducted. In most locations where plate subduction occurs major earthquakes are generated by the plates scraping together, but there has not been a subduction earthquake off the Oregon coast since the settlement by Euro-Americans. However, evidence has been found that such earthquakes occurred in the prehistoric past, evidence that includes estuarine marsh sediment buried by layers of sand, which had been transported far inland by huge tsunami waves that had accompanied each earthquake (Atwater 1987). Based on the numbers of such layers discovered by geologists along the coast, it  has been concluded that catastrophic earthquakes and  tsunami have occurred repeatedly in the past, with intervals ranging from 300 to 600 years. Carbon-14 dating of the buried marshes indicated that the most recent event occurred about 300 years ago, with its exact date having been established by Satake et al. (1996) to have been 26 January 1700, based on the arrival of the generated tsunami along the coast of Japan where it destroyed a number of villages. From the size of the tsunami waves that reached Japan, it was concluded that the earthquake must have been about magnitude 9, with the Oregon event having been comparable to the Sumatran subduction earthquake and tsunami in 2005. As the oceanic plates are being subducted beneath the continent, sediment that had accumulated on the seafloor as the plate moved slowly toward the coast is scraped off and added to the continental mass. This is the origin of the  Tertiary mudstones and siltstones that form the sea cliffs along the Oregon coast, while the rocky headlands are composed of volcanic rocks whose origins are also

⊡⊡ Fig. 1.3.1 The coast of Oregon and its principal communities.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Oregon

a­ ssociated with this tectonic setting. Nearly all of western Oregon has been formed by continental accretion of ocean sediment and a series of volcanic seamounts and islands, or by entire blocks of the sea floor. The oldest rocks found in western Oregon date back to the Palaeocene and Eocene geological periods, i.e. 40–60 million years ago. About 5 million years ago, during the Pliocene, the mountains of the Coast Range progressively emerged from the sea and western Oregon as we recognise it today came into existence. Upon emergence from the sea, the rain, rivers and ocean waves went to work to erode what formerly had been deep-sea crustal rocks, the basalts of former volcanoes, and ocean sediment now hardened into rock. These processes etched out the land, cutting away the weaker rocks, leaving behind the more resistant to form the peaks of the Coast Range and headlands along the coast. As on the coasts of Washington and northern California, there are numerous small islands and stacks in Oregon’s coastal waters, representing remnants of resistant rocks left behind as the coast retreated (>Fig. 1.3.2). The modern morphology of the coast is the product of the erosion that has taken place during the past 5 million years. Due to the tectonic setting of the Oregon coast, there are significant land elevation changes that affect the ­relative sea level, the change in the elevation of the land relative to that of the sea, which is globally rising at a rate between 1 and 2 mm/every year. The studies of the buried marsh deposits in estuaries that documented past occurrences of subduction earthquakes also supported the conclusion

that nearly all of the Oregon coast abruptly subsided during those tectonic events, often by 1–2 m (Atwater 1987). In contrast, much of the coast is presently rising, with that south of Florence being uplifted faster than the global rise in sea level. This has been documented through analyses of benchmarks used by surveyors, which the government re-surveys every few years to determine their elevation changes. South of Florence the tectonic rise of the land represents a net rate of land emergence above the sea of the order of 1 mm/every year, while along the north coast the relative sea level rise is on average −1.5 mm/every year (the negative value signifying that the global rise in sea level is faster than the change in land elevations). The consequence of this difference is that during historic times coastal erosion has been substantially less along the southern Oregon coast where the land is emerging from the sea, evident in the vegetated sea cliff at Bandon (>Fig. 1.3.2), in contrast with the significant erosion that is occurring along the northern Oregon coast (Komar and Shih 1993). The change from abrupt subsidence of the land during a subduction earthquake, compared with the progressive aseismic uplift now experienced, is interpreted in terms of the present accumulation of subduction strain between earthquake events, causing the slow rise of the land while the plates are locked together; the sudden release of that strain at the time of an earthquake results in the immediate subsidence of the land. The long-term net tectonic uplift of the Oregon coast, together with the cycles of sea level rise and fall during the

⊡⊡ Fig. 1.3.2 The sea cliff and offshore stacks at Bandon on the southern Oregon coast. The numerous stacks are interpreted as evidence for the rapid erosion of this shore following the 1700 earthquake when this area abruptly subsided, while the near absence of erosion today is attributed to the ongoing tectonic uplift of the land, which exceeds the rate of rise in sea level.

Oregon

Ice Ages, has given rise to marine terraces that in places form stairways up the flanks of the Coast Range. Along the south coast the highest and oldest terrace in the series reaches elevations up to nearly 500 m. The lowermost terrace, extending along much of the coast, dates back about 80,000 years. The Pleistocene terrace sands (former beaches and dunes), together with the ­underlying Tertiary mudstones and siltstones, are being eroded by the waves, forming sea cliffs that back sand beaches. Many of Oregon’s coastal communities are situated on this nearly level terrace; cities such as Cannon Beach, Lincoln City and Newport have suffered property losses as the cliffs progressively retreated (>Fig. 1.3.3). State Parks have also been impacted, with the loss of picnic grounds and camping facilities. In total, sea cliff erosion and property losses affect hundreds of km of the Oregon coast (Komar 1997). Landslides are common along the Oregon coast. The sea cliffs cut into the marine terraces are particularly susceptible to mass movement on various scales, with the occasional formation of large landslides (Komar 2004). This susceptibility is due in part to the consistency of the cliffs, landsliding being most active and destructive where the cliffs are composed of Tertiary mudstones and with their layers dipping toward the ocean. Most destructive in terms of losses of private property has been the Jump-Off Joe Landside in Newport (Komar 1997). Its initial movement began during the winter of 1942–43, affecting about 15 acres along with the loss of 15 homes. An attempt was made to re-develop the site in 1980, with plans to build

⊡⊡ Fig. 1.3.3 Sea cliff erosion at Gleneden Beach south of Lincoln City, the cliff being composed of uplifted Pleistocene terrace sands, with their erosion being the main source of sand to the beach.

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condominiums on the down-drop block of the landslide. This plan was prevented when the State rejected the developer’s request to construct a sea wall along the toe of the slide, to prevent further wave ­erosion and its continued seaward movement. The developer instead constructed the condominiums on a small remnant of the marine ­terrace to the immediate north of the Jump-Off Joe Landslide, beyond which was a second older landslide of comparable size. As the condominiums approached completion in 1981, slippage in the remnant terrace undermined the foundation, leading to their des­truction (>Fig. 1.3.4). With the developer having gone bankrupt, the  city had to cover the expenses of tearing down the condominiums. Landslides are also found on the basaltic headlands, or more precisely within the loose debris shed from the headlands that had accumulated along their flanks. Massive landslides associated with headlands have affected a few private homes, but in particular park lands such as Ecola State Park on Tillamook Head north of Cannon Beach. A number of large inactive landslides are found along the coast, believed to have formed at the time of the 1700 ­subduction earthquake; a few experienced renewed movement when their forest cover was removed by commercial logging. The differential rates of erosion and long-term retreat of the different rock types found along the Oregon coast have given rise to its irregular shore, with headlands of resistant basalt separating embayments formed by the

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⊡⊡ Fig. 1.3.4 The condominiums constructed on the remnant of terrace adjacent to the Jump-Off Joe Landslide, with slippage of the bluff having stressed the foundations so the windows were broken, evident in this 1981 photo.

more rapid retreat of the less resistant mudstones and siltstones. These embayments are the principal sites of beach sand accumulation, each constituting what is termed a littoral cell, basically a stretch of beach that for the most part is isolated by the large headlands that prevent the exchange of sand between adjacent cells: >Figure 1.3.5 is a typical example of a littoral cell. The individual cells contain different quantities of sand, apparent in the widths of their beaches, with the quantities depending on the existence of sand sources such as rivers or sea cliff erosion, which differ from cell to cell. A few cells have virtually no modern-day sources, the sand in its beaches instead being relict, having reached the cell thousands of years ago. In a study of the mineral contents of the sands within the cells we found that the metamorphic rocks of the Klamath Mountains in southern-most Oregon and northern California had been a prime source of sand, even to the cells along the central to northern Oregon coast where the numerous headlands now make it impossible for the Klamath sand to reach those beaches (Clemens and Komar 1988). We concluded that the Klamath derived sand had been transported north­ ­ward during times of lowered sea levels in the Ice Ages, unhindered by the presence of headlands when the shores were seaward on what today is the continental shelf. With the melting of glaciers and the rise in sea level, those shelf sands were pushed landward within the migrating beaches, having become trapped between headlands in the littoral cells roughly five thousand years ago. With the beaches being contained within littoral cells bounded by large rocky headlands, they are in effect pocket

beaches, even though they may have long shore lengths of 10 to over 100 km. In general, during the summer months the waves arrive predominantly from the west to ­northwest, and this causes a southward displacement of the sand within the cells. In contrast, the waves of winter storms mainly arrive from the southwest, moving the sand back to the north. As a result, there tends to be a seasonal northsouth oscillation of beach sand within the littoral cells, but with the long-term net littoral drift effectively being zero. The existence of this net-zero longshore transport of beach sand was evident when jetties were constructed during the early twentieth century on the inlets to estuaries and bays. In contrast to jetty construction in Southern California and along most of the U.S. East Coast, which did block a net longshore transport of beach sand so that it accumulated to one side of the jetties while erosion occurred in the down-drift direction, the jetties along the Oregon coast generally resulted in sand accumulation both to their north and south sides, locally where the shoreline is partially sheltered from the waves by the jetties (Komar 1997). Thus, jetties constructed on the Oregon coast have not been a problem in terms of having induced erosion and property losses. However one exception was dramatic, where early in the twentieth century the community of Bayocean was lost to erosion, but that occurrence can be attributed to only a single jetty having been constructed, not the usual pair, with the single jetty having caused the beach sand to be swept through the inlet into Tillamook Bay, lost from the sand spit on which the community had been developed (Komar 1997).

Oregon

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⊡⊡ Fig. 1.3.5 The Lincoln City Littoral Cell, with Cascade Head to its north and Cape Foulweather forming its southern boundary. Siletz Spit is positioned midway along the cell, with the remainder of the beaches backed by sea cliffs, the erosion of which is now the main source of sand to the beach as seen in (>Fig. 1.3.4); sand carried down the Siletz River is deposited within the Bay, so that it represents a minor contributor to the beach.

While the beaches along the Oregon coast are predominantly backed by sea cliffs, a number of sand spits are found separating the ocean from the estuaries and bays, as seen in (>Fig. 1.3.5) for Siletz Spit. These spits individually point either north or south, and some in close proximity and within the same littoral cell point in opposite directions; therefore, the direction of spit extension does not provide evidence for the existence or direction of a net ­littoral drift. While in total the progressive loss of property along the Oregon coast resulting from sea cliff erosion has probably been greater, more dramatic has been the erosion of developments on the sand spits, beginning with Bayocean and more recently with the major erosion that has occurred on Siletz Spit (>Fig. 1.3.6), Alsea Spit, and Netarts Spit, with the latter having impacted a State Park. These variable quantities of sand within the littoral cells and the resulting widths of their beaches depend on the existence of modern-day sources as well as the volumes of relict sand. This determines the rates of backshore property erosion, which varies from cell to cell, governed by the capacity of the fronting beach to buffer those properties from wave attack. The Oregon coast is one of the world’s most dynamic environments, with the extremes of its waves and tides accounting for occurrences of erosion like that seen in (>Fig. 1.3.6). These extremes have a direct connection with the Earth’s evolving climate, including the intensification of storms and the waves they generate, which may be due to global warming, and the periodic occurrences of

major El Niños. The Oregon coast is noted for the severity of its winter storms, which typically generate waves having deep-water significant wave heights (the average of the highest one-third of the waves) greater than 10 m, and with the significant wave heights during the most severe storms having reached 15 m, at which time the highest individual waves would have been about 25 m, the height of a 10-storey building. Daily measurements of waves off the Oregon coast have been collected by buoys since the mid-1970s. Of concern, those measurements demonstrate that the wave heights on an average have been progressively increasing, with the significant wave heights of the strongest winter storms back in 1975 having been about 9 m, having increased from 12 to 15 m in recent years (Allan and Komar 2006). While our analyses were based on wavebuoy measurements, data for the storm intensities in terms of wind speeds and atmospheric pressures extend further into the past, and demonstrate that the increases in wave heights likely began at least as early as the mid-twentieth century. Although the exact cause is uncertain, the increases in storm intensities and wave heights may be associated with global warming, though it has also been suggested that particulate pollution in the atmosphere may be important, drifting across the Pacific from China. The increases in deep-water wave heights and periods measured by the buoys off the Oregon coast have produced parallel increases in the processes active on its beaches, in particular the sizes of the breaking waves and elevations

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⊡⊡ Fig. 1.3.6 The erosion of Siletz Spit during the winter of 1972–73, with one home being lost and others on promontories of riprap placed to protect them. The erosion was caused by the coincidence of high storm waves and elevated tides, enhanced by seaward flowing rip currents that locally cut embayments into the beach.

reached by the swash runup of the waves when they reach the shore (Komar and Allan 2002). The runup elevations on beaches are particularly important in that they combine with the high tides to produce erosion of dunes and cliffs backing the beaches. Our analyses have shown that the progressive increase in wave heights has resulted in parallel increases in average swash runup levels at rates that are greater than the global rise in sea level, a factor that undoubtedly has played an important role in the increased property erosion experienced along the Oregon coast. Oregon’s tides are classified as mixed: there usually are two highs and two lows each day, but with the highs reaching different levels. With an average range of about 2 m and a maximum spring tide range of 4 m, they are further classified as mesotidal. These are the predicted astronomical

tides based on the forces of attraction of the moon and sun on the ocean’s water. The actual measured tides on the Or­­ egon coast can differ significantly from those ­predictions, with the difference primarily being of interest when the measured tide is substantially higher than predicted, since such occurrences can result in beach and property erosion. One cause of elevated measured tides is the occurrence of a storm surge, created by the onshore-directed winds and low atmospheric pressures of major storms. Measurements of surges on the Oregon coast show that they elevate tides of the order of 1.0–1.5 m, and although they are much smaller than storm surges produced by hurricanes along the U.S. East Coast, due to the low slopes of Oregon’s beaches its storm surges shift the mean-water shoreline landward by some 25–40 m, increasing the impact of the storm-wave runup on shore-front properties. The most significant climate event in terms of its ­erosion impacts along the Oregon coast has been the ­occurrence of a major El Niño, like those during the winters of 1982–83 and 1997–98 (Komar et al. 2000; Allan and Komar 2006). Particularly noteworthy is that an El Niño significantly elevates the measured tides, on average by about 0.30 m but achieving a maximum difference of about 0.60 m between the measured and predicted tides. This is documented by assessments of the monthly averages of the increased tidal elevations, with the maximum occurring during December and January, corresponding to the months that tend to have the greatest numbers and intensities of storms. The increased water elevation itself can be accounted for in part by the thermal expansion of the coastal water, which even in normal years is warmer during the winter due to the presence of cold water in the summer caused by upwelling, the water achieving still higher temperatures and levels in an El Niño winter. Another component results from the northward-flowing coastal currents, with their deflection to the right by the Coriolis force acting to pile water up along the shore; again, which tends to be stronger during an El Niño, resulting in elevated monthly-averaged water levels. Of im­­ portance to the resulting erosion of the Oregon coast, this  increase in water levels spans the entire winter, in effect representing a sudden increase in mean sea level, even though it later returns to normal when the El Niño ends. This rise in the monthly-averaged water levels during an El Niño elevates the water at all stages of the tides, so for several months there is an appreciably enhanced ­probability that the runup of storm waves on beaches will impact shore-front properties. The occurrence of a major El Niño also tends to, on average result in a degree of increase in the heights of winter storm waves that reach the Oregon coast, but far more

Oregon

important is that the waves arrive more from the southsouthwest than during normal years. This is because the tracks of the El Niño storms are shifted more to the south as they cross the North Pacific, passing over the shores of California rather than Oregon and Washington as they do during normal winters. The result is that unusually large quantities of the beach sand within the littoral cells are transported northward during an El Niño winter, resulting in ‘hot spot’ erosion at the south ends of the cells, to the north of the headlands, as depicted in >Fig. 1.3.7. Furthermore, the enhanced northward longshore sediment transport also tends to deflect the inlets to bays and estuaries, forming ‘hot spot’ erosion sites to the north of those inlets (or where present, north of jetties that act as mini-headlands). Following the El Niño winter the sand slowly returns to the south within the littoral cells, eventually reestablishing the long-term equilibrium sand volumes, although this return may take several years during which the erosion within the hot-spot zones persists. Erosion during the major El Niños of 1982–83 and 1997–98 extended along the entire U.S. West Coast, with that in northern California, Oregon and Washington ­having primarily occurred in hot-spot zones as depicted in >Fig. 1.3.7. Furthermore, along the Oregon coast the ­1997–98 El Niño erosion was significantly expanded ­during the following winter of 1998–1999 when several storms generated waves that exceeded what then had been ⊡⊡ Fig. 1.3.7 The contrasting reversing longshore sand movements within Oregon’s littoral cells during a Normal Year versus a major El Niño when strong storms from the southwest result in an enhanced northward transport and the development of ‘hot spot’ erosion sites.

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projected to be the 100-year extreme event (later to be reanalysed and increased to a 16-m significant wave height). The combined impacts of those two winters were extensive, to both private properties and State parks, in what we characterised as having been a ‘one-two punch.’ >Figure 1.3.8 shows the erosion at Cape Lookout State Park on Netarts Spit, an example of hot-spot erosion in that the Park is located at the south end of the littoral cell, to the north of Cape Lookout. The erosion began during the 1982–83 El Niño (Komar 1997), at which time the high tree-covered dunes were eroded away, so the public bathrooms within the campground landward of the dunes were threatened. Although they were ­temporarily protected by the ­placement of riprap, it proved insufficient during the 1997–98 El Niño and major storms that swept across the campground during the following winter, damaging the bathrooms to the extent they had to be torn down. The erosion of sea cliffs and foredunes along the Oregon coast has led to the proliferation of riprap revetments, with an example being those on Siletz Spit seen in >Fig. 1.3.6. The use of a ‘hard’ structure was undesirable in Cape Lookout State Park following the El Niño erosion, so it was decided to follow a ‘design with nature’ approach by constructing an artificial dune containing a core of sand-filled geotextile bags and planted with native dune grass, fronted by a ­constructed cobble berm (‘dynamic revetment’), in essence a nourished cobble beach as found naturally along

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⊡⊡ Fig. 1.3.8 The 1982–83 El Niño erosion at Cape Lookout State Park, with the loss of the high tree-covered dunes and vertical supports of a failed log sea wall seen in the background. The bathroom was lost during the 1997–98 El Niño.

the Oregon coast, thus acting to dissipate the swash of the waves before they impact the rebuilt dunes. The resulting ‘structures’ (>Fig. 1.3.9) have the appearance of their natural counterparts along the coast, and thus far have provided protection to the State Park (Allan and Komar 2004). The use of ‘hard’ structures – riprap revetments and sea walls – along the Oregon coast has had adverse impacts, especially where they have been used to prevent the erosion of sea cliffs. In some littoral cells the cliff erosion is the primary source of sand to the beach, in locations where the cliffs are composed of uplifted Pleistocene beach and dune sand; for example, the erosion at Gleneden Beach (>Fig. 1.3.3) within the Lincoln City Littoral Cell (>Fig. 1.3.5) south of Siletz Spit. There has been a considerable proliferation of structures within that cell, to the extent that the cliff-erosion source of sand to the beach has largely been cut off. Based on the previous rates of cliff retreat within that cell, it was estimated that the quantities of sand added to and building up the beach effectively balanced the relative rise in the sea level at that site (Shih and Komar 1994). With the loss of that sand source due to the installation of shore-protection structures, this equilibrium no longer exists, with the expected prolonged detriment of the beaches within that cell. Impressive accumulations of dune sands are found on the Oregon coast estimated to be present along about 45% of the coast, either in the form of foredunes backing the

beaches or contained within the massive Oregon Dunes Recreation Area that extends from Florence, south to Coos Bay. People have had a major role in altering the vegetationcover of the dunes, which in turn has affected their morphologies, with mixed consequences. When EuroAmericans first settled the Clatsop Plains south of the Columbia River in the nineteenth century, the extensive dune fields were covered by dense grasses. Those native grasses could be eaten by livestock, and overgrazing quickly reactivated the dunes so that by the 1930s some three thousand acres of sand had become mobile. In 1934 this area was planted with European beach grass, which livestock will not generally eat. Its introduction has had unforeseen consequences as it rapidly spread along the coast. On the positive side, it captured sand blowing landward from the beaches, building up substantially higher foredunes than had previously existed with the native grasses, providing a greater degree of protection from erosion and flooding to backshore properties. However, this growth of the foredunes has had negative environmental consequences, particularly to the nesting of Snowy Plover that need areas of open sand; in recent years extensive efforts have been undertaken to locally remove this invasive dune grass, to provide nesting habitat for this endangered shore bird. The arrival of the European beach grass has also had negative consequences for the Oregon Dunes Recreation Area. A century ago those dunes existed as an unvegetated

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⊡⊡ Fig. 1.3.9 The reconstructed dunes with a core of sand-filled geotextile bags, and cobble berm at Cape Lookout State Park, used as an alternative form of shore protection (Allan and Komar 2004).

sand surface extending from the ocean shore to the ­precip­itation ridge of the dunes at their landward edge (Cooper 1958). Sand was free to blow inland from the beach to support the continued growth of the high dunes. However, since the arrival of the European beach grass in the 1930s, large foredunes have grown immediately landward from the beach, where they now capture the sand blowing inland, preventing it from reaching the inland dunes. The impact of that loss was first noted in the area immediately landward from the foredunes, where the ground level was lowered to the water table, permitting the growth of shrubs and other vegetation where the high dunes had previously existed. The aerial extent of the active dunes has substantially decreased, and there is concern regarding their long-term preservation. An attempt was made to remove the beach grass and foredunes along a portion of the Oregon Dunes, but that experiment has not been followed up by a larger-scale implementation of this potential solution.

References Allan JC, Komar PD (2004) Environmentally compatible cobble berm and artificial dune for shore protection. Shore & Beach 72:9–18

Allan JC, Komar PD (2006) Climate controls on US West Coast erosion processes. J Coast Res 22:511–529 Atwater BF (1987) Evidence for great Holocene earthquakes along the outer coast of Washington state. Science 236:942–944 Clemens KE, Komar PD (1988) Oregon beach-sand compositions produced by the mixing of sediments from multiple sources under a transgressing sea. J Sediment Petrol 56:15–22 Cooper WS (1958) Coast dunes of Oregon and Washington. Geol Soc Am Memoir 72:169 Komar PD (1997) The Pacific Northwest coast: living with the shores of Oregon and Washington. Duke University Press, Durham, NC, p 195 Komar PD (2004) Oregon’s coastal cliffs: processes and erosion impacts. In: Hampton MA, Griggs GB (eds) Formation, evolution and stability of coastal cliffs – Status and trends. US Geological Survey Professional Paper 1693:65–80, Washington, DC Komar PD, Allan JC (2002) Nearshore-process climates related to their potential for causing beach and property erosion. Shore & Beach 70:31–40 Komar PD, Shih SM (1993) Cliff erosion along the Oregon coast: a ­tectonic – Sea level imprint plus local controls by beach processes. J Coast Res 9:747–765 Komar PD, Allan J, Dias-Mendez GM, Marra JJ, Ruggiero P (2000) El Niño and La Niña: erosion processes and impacts. In: Proceedi­ngs of 27th International Coastal Engineering Conference, Sydney, New South Wales, Australia, pp2414–2427 Satake K, Shimazaki K, Tsuji Y, Ueda K (1996) Time and size of a giant earthquake in Cascadia inferred from Japanese tsunami records of January 1700. Nature 379:246–249 Shih SM, Komar PD (1994) Sediments, beach morphology and sea cliff erosion within an Oregon coast littoral cell. J Coast Res 10:144–157

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1.4 California

1.  Introduction The California coast is about 2,900 km long, which includes the San Francisco Bay and the offshore islands. Some 70% of the coastline is rocky or has cliffs,  although pocket and fringing beaches are found commonly below the cliffs. The rest comprises either sandy fringing beaches backed by dunes, as south of San Luis Obispo, barrier beaches fronting lagoons and wetlands, as near Eureka, or the wetlands around San Francisco Bay. As on the Oregon coast there are numerous islands and stacks in the coastal waters of northern California that are scattered residuals of hard rock left by the recession of cliffs cut into softer sediment. About 85% of the coast is actively eroding. Although long stretches of exposed west-facing shores remain largely natural, extensive stretches around San Francisco, Los Angeles, and San Diego have been much altered by development. Of the original 80,000 ha of wetlands (excluding San Francisco Bay), 52% have been destroyed by dredging and filling, and a further 40% have been moderately or severely damaged. The coast embraces five geomorphic provinces, the Klamath Mountains, Northern Coast Ranges, Southern Coast Ranges, Transverse Ranges, and Peninsular Ranges. These provinces are structural units whose geology and relief reflect post-Palaeozoic interaction between the ­westward-moving North American lithospheric plate and subducting plates farther west. California’s structural framework began to emerge in late Jurassic and Cre­taceous times as the Nevadan Orogeny caused the ancestral Kalmath Mountains, Sierra Nevada, and Peninsular Ranges to rise from shallow seas and an island arc along the western margins of the North American plate, accompanied by westward thrusting, volcanism, and batholith emplacement. Subduction of marine sedimentary and ­volcanic rocks in the deep eugeosynclinal trench along this plate margin, followed by uplift, produced the Franciscan Formation, a heterogeneous jumble of readily erodible greywackes, shales, and metamorphic rocks found throughout the Coastal Ranges. During late Cretaceous and Cainozoic times, great

thicknesses of clastic sediment and submarine volca­­nic rocks accumulated in subsiding basins along the ­continental margin and strike-slip faulting became significant. Orogenic activity culminated in the uplift of coastal mountains from Miocene to mid-Pleistocene times. Coastal lowlands were restricted to subsiding intermontane troughs (Ventura Basin), to pull-apart structures (Santa Maria Basin), or to the seaward ends of strike valleys (Eureka Basin). During Neogene times, the North American plate approached and then overran the East Pacific Rise, a complex spreading centre in the Pacific plate. As that portion of California west of the San Andreas Fault transferred to the western limbs of this spreading centre, the Peninsular Ranges mini-plate (including Baja California) accelerated northwest from 2 to 5 cm/a year causing the Transverse Ranges to rotate into their present east–west alignment. Reflecting the influence of en echelon transform structures in the subducted East Pacific Rise, extensive strike-slip faulting with locally intense thrust faulting came to dominate coastal zone structures and landforms. Tectonism has persisted throughout Quaternary times, as shown by deformed and faulted marine terraces and the  subsiding coastal basins. A spectacular flight of 13 marine terraces are found on the Palos Verdes Peninsula, the highest at 380 m above sea level being equated with the nearby mid-Pleistocene deposits (Woodring et al. 1946; Wehmiller et al. 1977). A similar terrace sequence are found on San Clemente Island to the south, where the straight cliff along the NE coast is probably a fault scarp. On San Nicolas Island farther west, the 240-m terrace   dates back to the mid-Pleistocene age, while a 120-m terrace is dated as far back as 190,000 years. Fragments of high terraces are indeed found throughout the coast-from the deformed sequence near Eureka, along the seaward slopes of the Coast Ranges, to the extensive mesas around San Diego where the 90–150-m Linda Vista terrace and its fossil barrier beaches are considered to be a million years old. However, it is the last (Sangamon) interglacial terraces that provide the most continuous record of Quaternary tectonism.

Edited version of a chapter published by Anthony R. Orme in The World’s Coastline (1985: 27–36). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Terraces at 6 m and -14 m near San Diego have been dated at 120,000 and 80,000 years  (Kern 1977); a 23–45-m terrace at Palos Verdes has been dated at 85,000 years; 30-m terraces along the Malibu coast have ages of 130,000 and 105,000 years; while a Holocene terrace forms part of an impressive sequence west of Ventura, where uplift of 10 m/1,000 ayear has characterised the past 0.6 million years (Yeats 1978; see Orme 1980 for comprehensive bibliography). Terrace uplift and deformation is also impressive between Santa Cruz and San Francisco. Evidence for coastal submergence (by eustatic, isostatic or tectonic movements) is also compelling. The Ventura Basin, for example, subsided tectonically at rates up to 9.5 m/1,000 year from 2.0 to 0.6 million years ago, allowing a 7-km thickness of Plio-Pleistocene sediment to accumulate beneath the Oxnard Plain. More recently the Flandrian marine transgression entered the Golden Gate about 10,000 years ago and, rising at a rate of 2 cm/a year, flooded laterally across San Francisco Bay as rapidly as 30  m/a year until 8,000 years ago (Atwater et al. 1977) (>Fig. 1.4.3). The rate of sea level rise then declined and for the past 6,000 years has averaged 0.1–0.2 cm/a year. During this time, however, Holocene salt marsh deposits within the Bay have undergone 5 m of tectonic and possibly ­isostatic subsidence, a rate of 0.8 mm/a year superimposed  onto the Flandrian transgression eustatic effects. Subsidence is even more pronounced along the San Andreas Fault zone through the Bolinas Lagoon (Berquist 1978). Over the past century, tide gauge data show that the sea level has been rising at an average rate of 1.5 mm/a year along much of the coast (Hicks and Crosby 1974). Thus, despite ample evidence for Quaternary uplift and deformation, recent coastal changes must be viewed against the Flandrian transgression and a continuing sea level rise. Tides, winds, waves, currents, tsunamis, mass wasting, sediment discharge, and human activity all influence the patterns of coastal erosion and deposition. Mean spring tide ranges are between 1.5 and 1.8 m along the coast maximum tides attaining 2.5 m. Tidal currents are strong through narrow straits. At the Golden Gate, the flood inflowing tide reaches 1.7 m/s and the ebb 2.3 m/s, augmented by the head of stream discharge through San Francisco Bay. At the mouth of Tomales Bay, the ebb current clashes with incoming swells to form dangerous 3-m sneaker waves over the bar. Prevailing winds reach the coast from the northwest, producing waves and determining dune orientation. North of Point Conception 30–50% of ocean swell arrives from the northwest and most ­others from WNW. South of Point Conception, changing coastal orientation, strong refraction, and offshore islands

cause 70% of swells to pass up the Santa Barbara Channel from due west, while 80% of swells approach Los Angeles from the WSW. Here also, southerly swells set up by late summer hurricanes off western Mexico, by Southern Hemisphere winter storms, and by local winter depressions passing along more southerly tracks may cause erosion on south-facing beaches. Wave heights at Ventura average 1 m but range from 0.3–7.0 m (Orme 1982). The predominant northwest swells set up strong longshore currents, up to 2 m/s, and longshore drift from north to south, although northward drift may be favoured by shore configuration and reversing currents. The effect of the 1,000 km wide cold California Current (mean velocity 0.1–0.3 m/s; net discharge 11 million m3/s) is more climatic and ecological than geomorphic, but fog affects weathering and soil moisture on coastal slopes and the ocean current and its inshore counter current can transport fine sediment flushed out of coastal rivers. Tsunamis occasionally affect the coast, an example being the 1964 Alaska earthquake being followed by a series of up to 7 m waves on the coast at Crescent City. Stream sediment discharge, so important for nourishing California’s beaches, reflects the availability of rock waste, the precipitation and runoff patterns, and the extent to which fluvial transport has been blocked by dams. Along the northern California coast, erodible rocks, steep slopes, and high precipitation and runoff combine to ­produce frequent landslides and high erosion rates. The Klamath, Eel and Russian Rivers together account for 77% of all fine-grained sediment discharged to the sea north of Point Conception. Sediment yielded by the Sacramento-San Joaquin river system are mostly deposited before reaching the ocean. Farther south, Franciscan rocks and Cainozoic sediment are similarly erodible and prone to mass movement, but whether this debris reaches the coast depends on the winter precipitation regime, notably on the frequency of high magnitude storm flows. For example, most of the 9 million tonnes of sand delivered to the shore by the Santa Clara River between 1933 and 1938 arrived in 6 days of floods in 1938. Further, sediment discharged by this river in the 1969 floods was 47.6 million tonnes, compared with 1 or 2 million tonnes in relatively dry years. The coastal sediment budget of southern California thus sees years of plenty (notably 1969, 1978, 1980 and 1983) and years of famine (the 30 years preceding the 1969 floods). Mass movement from sea cliffs and coastal slopes is also an important sediment source, notably where fractured or poorly consolidated rocks are found, but this source has been both aggravated and curtailed by road and other construction

California

work, for example along the Malibu coast west of Los Angeles and on the Big Sur coast. In many places, for one reason or another, Pleistocene landslides and slumps have been reactivated in recent times, often through changes in groundwater hydrology. Offshore California is represented north of Point Conception by a shelf and slope, and southward by a continental borderland. The relatively narrow northern shelf widens to 50 km off the Golden Gate but is dissected by several submarine canyons, notably Delgada, Noyo, and Bodega canyons off northern California and Sur, Lucia, and Arguello canyons farther south. The largest of all is the Monterey Canyon, which heads within 0.8 km of the shore but dissects the shelf and slope to a depth of over 3,500 m. A large lunate sand bar composed of terrestrial sediment lies off the Golden Gate. The continental borderland off southern California comprises a series of fault blocks and troughs, with some closed basins and rugged emergent islands. These structures trend east-west off the Transverse Ranges and northwest-southeast farther south, and their origins must be related to the Cainozoic behaviour of the adjacent land areas (Emery 1960). At least 32 distinct submarine canyons dissect southern California’s offshore area.

2.  The Coastline of Northern California South from the Oregon border the coast consists of low cliffs and bluffs fringed by grey sandy beaches, with a few ⊡⊡ Fig. 1.4.1 Stacks and bluffs in the bay south of Trinidad Head. (Courtesy Geostudies.)

1.4

islands offshore. Lake Earl is a lagoon north of Point St George, where the bluffs steepen and are undercut by low cliffs. The northern California cliffs and bluffs are cut in pulverised eugeosynclinal Franciscan rocks consisting of greywackes and shales, and though they are forested they are prone to frequent mass movement and rapid erosion. These rocks strike NW-SE and have been eroded into narrow valleys that follow the strike and open into bays separated by elongated ridges that end in headlands. On this exposed coast winter storms generate waves 4–7 m high  at Point St George where the coast turns southeast, and Cres­ cent City stands on a low plateau edged by cliffs, with a series of headlands and bays. The tsunami of 27 March 1964 sent four waves, the highest over 6 m, in through the town, demolishing 29 city blocks. Sandy barriers enclose Freshwater Lagoon, Stone Lagoon and Big Lagoon, separated by spurs of high ground. The lagoons have rushy marshes and are backed by steep slopes. At Trinidad the coast is steep and partly cliffed, and to the south the coast is fringed by numerous islets and stacks, consisting of hard chert, greenstone and sandstone eroded out of a matrix of soft mudstone in the Franciscan Mélange (>Fig.  1.4.1). One of these harder components protrudes from a cliff of slumping clay at Luffenholtz, and will in due course become a headland, then an island. The steep coast ends as the Little River flows out at the northern end of a long sandy beach. Humboldt Bay is a large coastal lagoon behind a dunecapped sandy barrier, with a central outlet bordered by jetties. The shores are partly marshy (> Fig.  1.4.2), and

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⊡⊡ Fig. 1.4.2 Flat-topped cliffs and stacks indicate the coastal terrace at Ten Mile River. (Courtesy Geostudies.)

shoreline erosion is extensive, possibly as a result of the stabilisation and dredging of the entrance to allow ships access to Eureka. The River Eel carries large quantities of sand and gravel down to an estuary largely filled with shoals, but it has not yet built a delta, and the river mouth opens between paired spits. A long steep and cliffy coast, trenched by deep valleys incised into the coast ranges, extends south, past Cape Mendocino to Punta Gorda, then southeast to Point Delgarda and Cape Vizcaino. There are ­sloping cliffs and sandy beaches, as at Union Landing. Headlands end in lines of stacks near Westport and there is a wide coastal terrace which extends on to flat-topped islands at Brutal Point. The even-crested cliffs end at Ten Mile River (> Fig. 1.4.2) and a long sandy beach is backed by dunes, some drifting inland. At Fort Bragg the cliffs return, fronting a coastal terrace, and there are again many rocky stacks. Glass Beach is unusual in that it is derived from broken glass, porcelain and other garbage dumped from the cliffs, a procedure that was halted in 1970. To the south, on the coastal terrace, is Mendocino, with cliffs in sandstone and clay and reefs with kelp at the mouth of Big River. There are beaches of grey sand and pebbles at the mouth of Little River, and the coastal terrace continues southward, backed by a steep rising slope. It was cut when sea level stood higher relative to the land, and rocky knolls that rise from the terrace were formerly stacks. Navarro River is one of several rivers that have incised deep steep-sided valleys across the coastal terrace. It has a small estuary, encumbered by sand spits with some lumber. At Alder Creek the cliffs pass

inland as bluffs behind a sandy beach and dunes, which extend to the mouth of Hathaway Creek at Point Arena. The sloping cliffs at and south of Point Arena are receding evenly, cut in soft shale over sandstone and rising to a broad coastal terrace. In places the basal sandstone forms shore platforms, usually coinciding with bedding planes, while interspersed outcrops of hard grey gnarled rock protrude as headlands and are eventually isolated as reefs and stacks. Inland, the Mendocino Range runs parallel to the coast, and there are segments where the coastal terrace fades out to a steep coast, as at Saunders Landing. Rivers deliver gravelly loads to the shore, and beaches are mixtures of sand and pebbles. South from Stewarts Point the cliffs are cut in yellow sandstones which show honeycomb weathering, and shore platforms are well developed near Walsh Landing. Fort Ross stands on a coastal terrace that slopes gently seaward, incised by short narrow steep-sided valleys and ending in cliffs and coves (> Fig. 1.4.3). The coastal terrace is interrupted by a steep sector, the Sonoma Coast, but resumes to be crossed by the Russian River, which descends from a deep gorge cut through the recently uplifted Coast Range. The estuary is blocked by a barrier of grey sand (> Fig. 1.4.4), through which the river generally maintains an outlet that is enlarged during flood discharge. Near Peaked Hill, former stacks rise abruptly from the coastal terrace, which is backed by slopes marking a degraded cliff. Shell Beach and Duncans Beach are wide in front of cliffs, but to the south the beach narrows and rapid cliff recession threatens houses at Sorano del Mar.

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1.4

⊡⊡ Fig. 1.4.3 Coastal terrace and cliffs at Fort Ross. (Courtesy Geostudies.)

⊡⊡ Fig. 1.4.4 Sand barrier at the mouth of Russian River. (Courtesy Geostudies.)

From Salmon Creek the sandy beach resumes, and becomes a barrier backed by dunes and curving out to Bodega Head. On the landward side is Bodega Harbour, with a narrow outlet at the western end of a curving sand spit, and the coast then runs close to the line of the San Andreas Fault along Bodega Bay. Dillon Beach is wide and sandy, receiving ocean swell, but Tomales Point, offshore, marks the beginning of the long narrow Tomales Bay, where wave energy is much reduced. There are successive

headlands and bays along the coastline that runs southeast, past the salt marshes in the Chileno River estuary (> Fig. 1.4.5) and along the seaward slopes of the Balinas Ridge. Bluffs pass into cliffs on the more exposed promontories, south of Marconi, and at the head of Tomales Bay there is a marshy delta at the mouth of Olemo Creek. The inner shore of the Point Reyes Peninsula is fairly straight, close to the San Andreas Fault, and consists of Mesozoic granite. Lateral movement along this fault is

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⊡⊡ Fig. 1.4.5 Salt marsh in Chileno River estuary, Tomales Bay. (Courtesy Geostudies.)

causing the Point Reyes Peninsula to move northwest along the coast at an average rate of about 5 cm/a year, much of the movement occurring during earthquakes, as in 1906 when there was sudden displacement of up to 6 m. The outer coast at Tomales Point is exposed to oceanic waves, and cliffs cut in red sandstone underlain by hard grey Sierran granite extend south past McLures Beach. The high energy Pacific swell generates major rip currents here. Much of the west coast of the Point Reyes Peninsula has a wide sandy beach backed by dunes, forming a barrier that impounds Abbotts Lagoon in a gently incised valley. At the southern end, Point Reyes is a steep-sided ridge running west to east, and on the south coast is the broad curve of Drakes Bay, with cliffs cut in soft Pliocene sediment (>Fig. 1.4.6) and beaches that become barriers and spits in front of lakes (Drakes Estero and the adjoining Estero de Limantour) impounded in rias. Drakes Head is a cliffed promontory behind the barrier spits, and an outlet from the esteros has an ebb tide delta that refracts incoming waves. The coast steepens on the flanks of forested ranges down to Bolinas Point and Duxbury Point. U.S. Coast Survey records indicate that these points have been receding at about 0.5 m/a year. The San Andreas Fault runs southeast from Tomales Bay along a valley that ends in Bolinas Lagoon, which is almost enclosed by a sandy spit. The steep forested coast then continues to Point Bonita, on the northern side of Golden Gate, the strait at the mouth of San Francisco Bay.

3.  San Francisco Bay San Francisco Bay occupies a late Pliocene to mid-­Pleistocene structural trough within the Coast Ranges. The trough has a Franciscan basement bounded by ­northwest-trending strike-slip faults and is largely filled with Neogene and Quaternary marine and fluvial sediment. Continuing tectonism apart, this trough has been ­further deepened by fluvial erosion and invaded several times by Quaternary transgressions (Atwater et al. 1977). The bay is connected to the Central Valley through the Carquinez Strait, with a bedrock channel incised by the ancestral Sacramento River to 60 m below sea level, and to the Pacific Ocean through the Golden Gate, with a channel that is scoured to 104 m below sea level. San Francisco Bay is a natural sump for an 80,000 sq. km drainage area comprising the Sacramento-San Joaquin system. Some 200 years ago the bay waters covered 1,800 sq. km, but they have been reduced to 1,100 sq. km, and 70% of the bay is shallower than 4 m. This reduction is due in part to massive fluvial sedimentation during and after a period of hydraulic gold mining in the Sierra Nevada foothills in the latter half of the nineteenth century, and in part to extensive coastal land reclamation. Gilbert (1917) estimated that 1,816 million m3 of debris were eroded in the foothills by mining and natural processes between 1850 and 1914, of which 50% was deposited in the valley downstream, 48% reached the bay and only 2% reached the ocean. He also estimated that it would take 50 years (until 1964) for the upstream debris to reach the bay, after which the yield

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⊡⊡ Fig. 1.4.6 Cliffs in Drakes Bay, Point Reyes Peninsula. (Courtesy Geostudies.)

would diminish to about 6 million m3/a year, a figure that accords well with recent estimates. Wave action within San Francisco Bay is limited by the short fetch, and breaker heights average at 0.3 m. Mean spring tide range is about 1.8 m, and tidal currents rarely exceed 1 m/s within the bay. The salinity averages 27–29‰, with dilution by inflow of about 570 m3/s from the Sacramento-San Joaquin River. The relatively low-energy shores have salt marshes around mean high tide level dominated by Spartina leiantha, Distichlis spicata, and Salicornia ambigua zones, upper tidal flats between mean high and mean low water and lower tidal flats below mean low water (Pestrong 1972). About 200 sq. km of marshes and mudflats are exposed at mean tide level.

4.  The Coastline of Southern California It is convenient here to consider Southern California as the coast, south from San Francisco. The Southern Coast Ranges extend 400 km southeast from the Golden Gate to Point Arguello and are dominated by massive deformation and faulting, strike valleys, and elongated ridges. Between the San Andreas and Nacimiento fault zones, metamorphic and granitic rocks form the 60-km-wide Salinian basement. These rocks reach the coast in rugged slide-prone cliffs south of Pacifica. Pillar Point is a southward projection bordering Half  Moon Bay, which has been shaped by ocean swell refracted around the point and has a beach showing lateral

gradation in grain size, coarsening and steepening as exposure to wave action increases southward (Bascom 1951). A generally steep coast with valleys incised into the ranges extends past Pigeon Point and Point Ano Nuevo, where the coast curves to Santa Cruz in Monterey Bay. Only one of the natural arches cut in Santa Cruz mudstone is still standing; the others have collapsed to leave stacks offshore. Late Cretaceous and Tertiary sediment reach the coast north of Monterey Bay in deformed marine terraces and eroding sea cliffs (Bradley and Griggs 1976). The terraces have been rising at 1.6–2.6 cm/century and long-term sea cliff erosion rates range from zero in resistant granodiorites to 0.3 m a year or more in softer Tertiary  sediment, reaching 0.6 m/a year locally (Griggs and Johnson 1979). Extensive sand dunes around Monterey Bay are ex­­ plained in part by the abundance of sediment formerly brought down by the Pajaro and Salinas rivers. For 150 km south from Monterey the Santa Lucia Range dominates the coast, often reaching 1,000 m within 2 km of the shore. Where Franciscan rocks reappear west of the Nacimiento Fault, landsliding is a perennial problem. South of Ragged Point, the mountains trend inland, emerged marine terraces reappear, and broad structural basins are veneered with extensive dunes of late Pleistocene and Holocene age, notably behind the 25 km of sandy beach between Pismo Beach and Point Sal. Sand for these dunes comes mainly from the Arroyo Grande and from the Santa Maria and Santa Ynez rivers. Cliff erosion threatens homes and roads at Cayucos, Avila Beach, and Shell

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Beach, despite partial shelter from northwesterly swells. In Morro Bay salt marshes are protected by a long barrier spit. In the hinterland is a series of transgressive dune formations (Orme 1990). The Santa Ynez Mountains, which rise to 1,430 m extend 125 km east from Point Arguello to Ventura, and are a south-dipping homocline of mostly Tertiary marine and fluvial sediment. Quaternary uplift has produced much deformed marine terraces along the coast. Though protected by offshore islands, refracted westerly ocean swell have produced average erosion rates of 0.15 m/a year in the relatively soft Neogene sediment along the Santa Barbara coast. A small barrier-lagoon system survives at Carpinteria. The Oxnard Plain is a triangular lowland, formed mainly by deposition from the Santa Clara River and Calleguas Creek, where the Ventura structural trough plunges westward into the Santa Barbara Channel. The coastline once comprised a sandy barrier beach with low dunes backed by lagoons and marshes, breached by rivers during winter floods, but there has been much construction along the shore. Breakwaters and groynes have been built at Santa Barbara, Ventura and Port Hueneme, and there have been problems with interception of eastward longshore drifting of sand by the prevailing westerly swell. About 750,000 m3/year of sediment moves along this coast in one of several littoral cells on the coast of Southern California, each with sand supplied mainly from rivers passing alongshore until it is lost into the head of a submarine canyon (Emery 1960). To the south of the Santa Barbara Channel the Channel Islands (Anacapa, Santa Cruz, Santa Rosa, San Miguel) are underlain by a metamorphic or granitic basement and mantled with thick late Cretaceous and Tertiary rocks, notably Miocene marine shales and volcanic formations. The Santa Monica Mountains, which rise to 949 m, extend 50 km along the coast between Point Mugu and Santa Monica and are bounded to the south by the active Malibu thrust fault, the impact of which is seen in seismic activity, deformed terraces, fractured coastal rocks, and frequent mass movement. Much of the coast is typified by  unwise housing development on the backshore or beneath crumbling cliffs, creating inevitable problems during storms and southerly swells. Access is restricted to many beaches, as at Malibu. The Peninsular Ranges extend SSE for 200 km behind Los Angeles to the Mexican border. This mini-plate, ­comprising basement rocks, granitic plutons and post-­ batholithic sediment, has been diverging from the North American mainland at a mean rate of 6 cm/a year since it shifted on to the western limb of the East Pacific Rise some 4–5 million years ago.

Along the coast deformed marine terraces such as those on Palos Verdes (> Fig. 1.4.7), San Onofre Moun­ ­tain  and around San Diego, and active faults as in the Inglewood-Rose Canyon structural zone that outlines the coast between Newport Bay and San Diego, attest to continuing tectonic activity. The beach fringed Los Angeles coast is heavily urbanised, and Long Beach is a port and industrial urban complex where coastal land subsidence has followed oil extraction. Barrier and lagoon systems were once common along this coast, but most have been modified, and some obliterated, by reclamation. Ballona Lagoon, once an outlet for the Los Angeles River, is now the marine  and Newport Bay, until 1915 the outlet for the Santa Ana River, was heavily developed. Offshore is Santa Catalina Island, which has warped emerged marine terraces. Much of the coast south of Newport is cliffed, between valleys where rivers supply sand to the shore, sometimes through small barrier lagoons. This is another littoral cell in which beaches have been described as ‘rivers of sand,’ fluvial sediment drifting southward along the coast until it disappears into a submarine canyon off San Diego. Sand is also derived from receding cliffs, partly by marine erosion and slumping and partly by the effects of runoff. Erosion is a significant problem along much of this coast. Powerful ocean swell and storm waves arrive through windows between the offshore islands, coastal rock formations are often poorly consolidated and, owing to floodcontrol dams and channelization along most of the rivers, fluvial sediment supply has diminished in recent decades, depleting protective beaches. Cliff erosion is particularly rapid in front of Camp Pendleton, south of Oceanside harbour, near Encinitas and Del Mar, and at Sunset Cliffs, San Diego. Groundwater discharge weakens soft cliffs, causing exudations and triggering slumping. Inappropriate clifftop development has generally aggravated the problem, for example, through irrigation and lawn watering, increased impermeable surfaces, and poorly located culverts. At La Jolla Scripps Pier has been used as a transect for measuring fluctuations in the beach profile, notably the cut and fill cycles that accompany alternations of winter storm erosion and summer shoreward drifting. These cycles can result in beach lamination, with layers of fine and coarse sediment deposited during accretion with phases of calm-weather fine sedimentation and rough weather winnowing. Rocky shore outcrops show evidence of weathering by alternations of wetting and drying, producing platforms at the level of permanent saturation, as well as corrosion bysalt weathering. Boomer Beach, near La Jolla, is well known for its seasonal alternations. In winter storm waves scour away

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⊡⊡ Fig. 1.4.7 Palos Verdes Bay, showing raised terrace. (Courtesy Geostudies.)

sand exposing a rocky and bouldery shore, but in summer constructive waves sweep sand back in from the sea floor to bury these beneath a sandy beach. To the south San Diego has an urbanised coast, and the former San Diego River outlet has become an important naval and commercial harbour, bordered by Coronado Strand, a built-over spit. The urbanised coast extends to the Mexican border.

References Atwater BE, Hedel CW, Helley EJ (1977) Late quaternary depositional history, Holocene sea level changes, and vertical crustal movement, Southern San Francisco Bay, California. US Geological Survey Professional Paper 1014, Washington, DC Bascom W (1951) Relationship between sand size and beach face slope. Trans Am Geophys Union 32:866–874 Berquist JR (1978) Depositional history and fault related studies, Bolinas Lagoon, California. US Geol Surv Open-File Report 78–802, Washington, DC Bradley WC, Griggs GB (1976) Form, genesis, and deformation of central California wave-cut platforms. Bull Geol Soc Am 87:433–449 Emery KO (1960) The sea off Southern California. Wiley, New York Gilbert GK (1917) Hydraulic-mining debris in the Sierra Nevada. US Geol Surv Prof Paper 105, Washington, DC

Griggs GB, Johnson RE (1979) Coastline erosion, Santa Cruz County, California. Geology 32:67–76 Hicks SD, Crosby JE (1974) Trends and variability of yearly mean sea level 1893–1972. US National Oceanic and Atmospheric Administration, Technical Memorandum NOS 13, Washington, DC Kern JP (1977) Origin and history of upper Pleistocene marine terraces, San Diego, California. Bull Geol Soc Am 88:1553–1566 Orme AR (1980) Marine terraces and quaternary tectonism, northwest Baja California, Mexico. Phys Geogr 1:138–161 Orme AR (1982) Temporal variability of a summer shore zone. In: Thorn CE (ed) Space and time in geomorphology. George Allen & Unwin, London Orme AR (1990) The instability of Holocene coastal dunes: the case of the Morro Bay dunes, California. In: Nordstrom KF et al (ed) Coastal dunes: form and process. Wiley, New York, pp 315–336 Pestrong R (1972) San Francisco Bay tidelands, California. Geology 25:27–40 Wehmiller JE, Lajoie KR, Kvenvolden KA et al (1977) Correlation and chronology of Pacific coast marine terrace deposits of continental United States by fossil amino acid and ­stereochemistry-technique evaluation, relative ages, kinetic model ages, and geological implications. US Geological Survey Open-File Report, Washington, DC, pp 77–680 Woodring WP, Bramlette MN, Kew WSW (1946) Geology and Paleontology of Palos Verdes Hills, California. US Geological Survey Professional Paper 207, Washington, DC Yeats RS (1978) Neogene acceleration of subsidence rates in southern California. Geology 6:456–460

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1.5 Texas

Robert Morton

1.  Introduction The Texas coast bordering the western Gulf of Mexico (>Fig. 1.5.1) is about 590 km long and consists of barrier islands (80%) and mainland beaches (20%). Of the sandy barrier segments, about 45% are retreating, 35% are progradational features and 20% are aggradational features.

Retreating barrier and beach segments, such as the Rio Grande deltaic headland and adjacent South Padre Island (>Fig. 1.5.2) are narrow, thin and have low topographic profiles that allow repeated overwash by extreme storms. These relatively young landward-migrating barriers are located on the flanks of eroding Holocene deltaic headlands constructed by the Rio Grande and the combined

⊡⊡ Fig. 1.5.1 The Texas coast consists of mainland beaches and barrier islands with associated bays and lagoons.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.5.2 South Padre Island is a narrow, low, retreating barrier that is flooded and overwashed frequently by hurricanes.

⊡⊡ Fig. 1.5.3 Central Padre Island is an aggradational barrier with wide stable or accreting beaches and high, vegetated sand dunes.

Brazos-Colorado rivers. In contrast, the aggradational barrier segments with associated tidal flats (>Fig. 1.5.3) and progradational barriers with associated washover fans (>Figs. 1.5.4 and > 1.5.5) are wide, thick and have high dunes that effectively prevent storm surges from washing over the barrier islands. These Holocene barriers formed within the interdeltaic embayments where sand was supplied by longshore currents. The western Gulf of Mexico was one of the first coastal regions, where systematic research regarding

modern sediment and coastal processes was conducted. These efforts, which began in the late 1940s, were largely driven by the petroleum industry that was interested in developing models for predicting subsurface sedimentary facies and reservoir characteristics. Classic field-based geologic studies of Holocene barrier islands (Fisk 1959; LeBlanc and Hodgson 1959), sea level history (Curray 1960) and sediment distribution (Shepard et al. 1960) used Texas coastal features as a research laboratory.

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⊡⊡ Fig. 1.5.4 Long, shore-parallel beach ridges on San Jose Island are characteristic of progradational barrier islands.

⊡⊡ Fig. 1.5.5 A large inactive hurricane washover fan on San Jose Island. (Courtesy Texas General Land Office.)

2.  Geological Setting Like other barrier coasts of the world, the Texas coast owes its position and shape to sea level fluctuations associated with expansion and contraction of continental ice masses during the late Pleistocene. During the last eustatic cycle, most coastal plain rivers maintained their courses, incising deep valleys during the falling phase and backfilling, as sea level rose to its current position. Bays

oriented perpendicularly to the modern coastline represent unfilled remnants of former incised valleys. Only rivers with large drainage basins, such as the Rio Grande and combined Brazos-Colorado systems, carried enough sediment to aggrade their upper valleys and prograde deltas of moderate size into the Gulf of Mexico. Although the general Holocene sea level curves show a gradual reduction in the rate of rise and relative stability beginning about 5,000 years ago (Curray 1960), there is convincing

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evidence that sea level was about 1–1.5 m higher than present during the past few thousand years (Morton et al. 2000). The beaches and barriers of the Texas coast are arranged in an orderly pattern alongshore (Morton 1979). The Holocene Rio Grande and Brazos-Colorado fluvial-deltaic systems and a headland composed of Pleistocene fluvialdeltaic deposits control the coastline types and their positions. These three headlands formed large promontories in the Gulf of Mexico that focused wave energy and created three cells of littoral drift convergence. The headlands are composed primarily of stiff mud (>Fig. 1.5.6) and the nonbarrier beaches that front the headlands are rapidly eroding (Morton et al. 2004). The longest barrier-island segment formed within the southernmost littoral drift cell, which extends from the Rio Grande delta to the Brazos-Colorado ⊡⊡ Fig. 1.5.6 Shores of the Holocene Brazos-Colorado delta, such as this one at Sargent, Texas, typically are composed of stiff, organic-rich mud overlain by thin washover terrace deposits.

delta. Retreating barriers within this segment, including South Padre Island (>Fig. 1.5.2) and Matagorda Peninsula, are located on the flanks of the deltaic headlands (Morton 1979). The barriers are migrating landward as sand eroded from the Gulf coastline is transported across the barriers through the washover channels and into the adjacent lagoons. Transitional barrier segments separate these migrat­ ing barriers from the progradational and aggradational barriers, including Central Padre Island (>Fig. 1.5.3) and North Padre Island, Mustang Island, San Jose Island, and Matagorda Island, which formed in the center of the southern interdeltaic embayment. Within the middle littoral drift cell, which formed between the Brazos-Colorado deltaic headland and the Pleistocene headland, sand-rich progradational barrier islands Galveston Island and Bolivar Peninsula and a thin transitional barrier (Follets Island) were deposited. The progradational barriers have moderately high beach-ridge elevations (3–4 m) and small sand dunes. Cheniers consisting of marsh mud alternating with narrow beach ridges of shelly sand, accumulated against the Pleistocene headland within the easternmost littoral drift cell. This coastline segment, which is west of Sabine Pass, forms the western extension of the chenier plain of southwestern Louisiana. Texas barriers are entering a new phase of their evolution because of recent increases in the rate of relative sea level rise and decreases in sediment supply (Morton and McGowen 1980). Retreating barriers continue to migrate landward, while most of the other barriers are getting narrower. This is because of simultaneous erosion along both Gulf and lagoon shores. An exception to this trend is Central Padre Island, which remains an actively aggrading barrier. Sand released by updrift beach erosion is transported to this barrier segment by converging littoral currents. The sand supplied to the Central Padre Island beaches, dunes and backbarrier environments (tidal flats) by longshore and aeolian transport, is sufficient to maintain a stable landform even as relative sea level rises.

3.  Climate The Texas coastal climate, which ranges from subhumid in the east to semiarid in the south, is greatly influenced by the systematic westward decrease in precipitation and increase in evapotranspiration. The strong climatic gradient has a dramatic affect on biological assemblages, landforms and physical processes (Morton and McGowen 1980). The average annual rainfall for the eastern coast is 125 cm. As a result, nutrient-rich freshwater inflow to the

Texas

bays supports lush saltwater marshes and large oyster reefs, but the higher rainfall also minimises the influence of aeolian processes. In south Texas, where evapotran­ spiration generally exceeds precipitation, coastal waters commonly have abnormally high salinity and saltwater marshes represent only minor bay-margin environments; moreover, large active dune fields and tidal flats dominate the nearshore landscape.

4.  Wind and Waves The Texas coast is a microtidal, storm-dominated region that is influenced periodically by energetic meteorological events (Hayes 1967). Most of the year, wind blows from the southeast, but during the winter, strong northerly winds accompany the passage of cold fronts that can suddenly lower the temperature, rapidly change wind direction and greatly increase wave heights. These winter storms are high-frequency events that occur several times each year. Highest sustained wind speeds accompany major hurricanes such as Carla (1961) and Allen (1980). These infrequent extreme storms are so large that their organised wind patterns fill the northwestern Gulf of Mexico, matching the scale of the coastline arc. The counterclockwise wind circulation of hurricanes, drives strong nearshore currents and sand transport to the west and southwest in the Gulf of Mexico. Fair-weather waves in water depths of 10 m are normally about 60 cm high and have periods of 6 s. The broad shallow continental shelf bordering the Texas coast causes deep-water swell in the Gulf of Mexico to decompose into these low, short-period waves. Although precise measurements of waves breaking along the South Texas coast have not been made, field observations indicate that wave heights are slightly greater along Central Padre Island where forebeach slopes and offshore gradients are slightly steeper than in adjacent areas. The largest deep-water waves recorded in the Gulf of Mexico were 27 m high and had period of about 16 s. These hurricane-generated waves break far from the shoreline because of the broad shallow shelf. The storm waves reform and repeatedly break, creating a wide energetic surf zone that encompasses as many as nine bands of breakers.

5.  Tides Astronomical tides in the western Gulf are diurnal or mixed and typically have a range of about 0.5 m during a normal tidal cycle. Tide ranges are even lower in bays and

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lagoons because of the small number of inlets and great distances from most inlets to the bay margins. Windinduced changes in bay water levels are commonly more significant than those caused by the astronomical tidal cycle (Kraus 2007). Lowest water levels in the bays and lagoons typically are associated with strong northerly winds associated with winter storms that drive water out of the bays. In contrast, highest water levels are caused by hurricanes. During some hurricanes, tidal inlets cannot maintain water levels between the lagoons and the Gulf. As a result, large hydrostatic differentials can be created across the barriers, and powerful currents erode deep washover channels through the barrier islands (Morton 2002).

6.  Present Sea Level Tide gauge records in Texas show the same general variations in sea level that coincide with droughts and periods of abnormally high rainfall. They also show a relative rise in sea level averaging 3.4 and 6.5 mm/year at Port Isabel and Galveston respectively. Some of the relative rise in sea level is caused by compaction of the basin fill and subsidence of the land surface. The high rates of relative sea level rise measured along the southeastern Texas coast are attributed to subsidence induced by shallow ground-water withdrawal and deep hydrocarbon production (Morton et al. 2006).

7.  Tidal Inlets Both primary and secondary tidal inlets are present along the Texas coast (Kraus 2007). A single primary inlet is located in the southern sector of each bay system because of the predominant wind directions and southwesterly sediment transport. Primary inlets are large permanent channels that serve as the principal conduits between the Gulf and large bay systems. The primary inlets are wide, short and oriented perpendicular to the Gulf coastline. This configuration makes them highly efficient trans­ porters of water and sand. They are also relatively stable, having occupied the same general positions during construction of the adjacent barriers. The secondary inlets are narrow, long and slightly oblique to the Gulf coastline. These moderately efficient but unstable inlets, migrated in the direction of net longshore transport before they were modified. A third category of inlets includes the small, ephemeral features that are occasionally active during hurricanes. They are highly inefficient and typically close a few years after being reopened.

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8. Beach and Barrier Island Characteristics Most Texas beaches are composed of well-sorted, subangular to subrounded, fine to very fine sand regardless of whether they are eroding or accreting. Exceptions are the highly erosional deltaic headlands, where mobile beach sediment forms a thin veneer over stiff mud. Locally, these beach deposits contain high concentrations of gravel-size terrigenous and biogenic detritus (Morton and McGowen 1980). Broken and whole shells constitute as much as 50% of some erosional beach deposits. The shells are of brackish to normal marine fauna (Rangia cuneata, Crassostrea virginica, Mercenaria campechiensis) that are diagnostic species of estuarine and bay environments. The non-biogenic gravel is composed of rock fragments and caliche nodules also eroded from the underlying sediment and concentrated on the beach. The gravel fraction of shells and rock fragments represents a coarse lag remaining, after constant winnowing and continuous erosion of sand-poor relict sediment exposed on the shoreface and inner shelf. High concentrations of shell also occur on Central Padre Island. These shells are of Gulf species that were left as a wind-deflation lag on an accreting beach fed by converging longshore currents. Elevations of Texas back-beaches are only about 1 m because of the small tide range and low wave heights. Storm surges generated by major hurricanes, commonly exceed the elevations of backbeaches by 2–3 m. Average wave characteristics and offshore slope produce an upper shoreface about 600 m wide where three migratory breaker bars are commonly maintained by the interaction of waves and longshore currents. Although much of the littoral drift is transported by traction and suspension in this zone of breaking waves, substantial volumes of sand are stored and transported across the middle and lower shoreface. The shoreface extends seaward of the surf zone to depths of at least 10 m and distances of more than 2 km offshore. During major storms, the surf zone widens, wave heights increase, longshore currents accelerate and sediment transport significantly increases along the shoreface. The Rio Grande serves as the international boundary between the United States and Mexico. It drains a large area that is mostly desert or has low rainfall. Withdrawal of water along its course for irrigation and industrial purposes has so severely reduced the river’s flow that the mouth is commonly closed by a shoal. Beaches near the mouth of the river are eroding because wave energy is moderately high and there is a lack of sediment supply. South Padre Island is a retreating barrier that forms the northern flank of the submerging Rio Grande delta.

The southern tip of South Padre Island is densely developed and the site of high-rise construction. The development on South Padre Island is progressing to the north into a hazardous area, where the island is low and narrow, the beach is eroding, the dunes are discontinuous and broad washover channels extend entirely across the island (>Fig. 1.5.2). Some buildings have already been constructed in washover channels. The semiarid climate of South Texas promotes aeolian sand transport and the migration of low active dunes over the road along South Padre Island. During extreme storms the dunes are eroded and become part of the washover sediment deposited in the backbarrier flat or in Laguna Madre. Padre Island, north of the Mansfield Ship Channel is transitional between the narrow migrating barrier of South Padre Island and the wide aggradational barrier of central Padre Island. Beach erosion and dune migration characterise this segment of the coast because the vegetation is too sparse to stabilise the sand. The most common backbarrier subenvironment along Padre Island is a tidal flat. These broad sand flats are typically barren except for a thin algal mat. Plants cannot survive this harsh environment that are sometimes exposed and dry, and sometimes flooded with marine (lagoon) water that becomes hypersaline as it evaporates. Central Padre Island is an aggradational barrier with stable or accreting beaches that receive sand delivered by longshore currents and aeolian transport. It is in the zone of convergence between currents flowing northward from the Rio Grande delta and southwestward from the central Texas coast. This barrier segment is characterised by high, continuous and densely vegetated dunes with elevations up to 15 m. The beach is composed mostly of shell fragments, which make the foreshore much steeper than the fine sand beaches to the north and the south. The backshore is broad and either flat or dips slightly landward where the berm crest is elevated by wave runup. Dunes on North Padre Island are generally stable except for remnants of dune fields that were active during the 1950s drought and blowouts that originate at the Gulf shore and migrate across the barrier. Beach erosion is slow, but persistent on North Padre Island as demonstrated by this sea wall. When the wall was built in the late 1960s, there was a wide beach separating the wall from the Gulf of Mexico. The progradational barriers of the central Texas coast are compound landforms that preserve backbarrier elements of their former transgressive phase, when they were being overwashed and migrating landward. The curvature of channels on the large washover fan on San Jose Island, (>Fig. 1.5.5) depicts the strong influence of hurricanes

Texas

and the counterclockwise flow of overwash generated by the wind. Similar fans also form the backbarrier on Matagorda Island. All these fan-like features are much larger than the washover fans constructed by extant storms, which suggests that storms during the mid to late Holocene may have been more intense. The progradational barrier islands also are characterised by continuous and densely vegetated dunes and broad and relatively stable beaches. Matagorda Peninsula is a low, narrow, migrating barrier that is located on the southwestern flank of the Brazos/ Colorado deltaic headland. It was devastated by Hurricane Carla (1961) that opened numerous washover channels across the barrier. The channels subsequently filled at their Gulf entrance, but they remain as vulnerable sites for future channel incision. Beaches along the Holocene Brazos/Colorado headland, such as at Sargent, Texas, are composed of stiff, organic-rich mud representing former floodplain marsh environments. These rapidly eroding beaches are overlain by thin washover terrace deposits composed of high concentrations of estuarine shells and rock fragments. Considering their overall morphology and welldeveloped beach ridges, Galveston Island and Bolivar Peninsula are progradational barriers. Although Galveston Island is a progradational landform, most of the beaches along the island are eroding. The highest rates of erosion are immediately downdrift (southwest) of the seawall. In the mid 1960s, a wide sandy beach extended seaward of the seawall and a concrete slab at the western end of the seawall formed a ramp for cars to drive directly onto the beach. The beach at this site has retreated more than 200 m and the erosion threatens nearby condominiums. The eastern third of Galveston Island is protected by a seawall that was constructed after the 1900 hurricane killed thousands of people. A field of groynes and riprap were added to protect the wall from persistent beach erosion. In 1995, a beach nourishment project was used to restore the recreational beach in front of the sea wall. The movement of beaches on Bolivar Peninsula is generally systematic with erosion being prevalent on the northeastern end, where it attaches to the Pleistocene headland, relatively stable in the western third, and accreting near the jetties at the tidal inlet entrance to Galveston harbour. The beach accretion is supplied by erosion of updrift beaches. In 1961, Hurricane Carla overtopped the eastern end of Bolivar Peninsula and deposited a wedge of washover sediment as

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much as 1.2 m thick. Subsequent beach erosion has exposed the washover deposits in a backbeach scarp. The headland east of Bolivar Peninsula is composed of stiff Pleistocene mud overlain by a thin wedge of cohesive Holocene marsh mud. A continuous washover terrace composed of sand, shell and rock fragments forms the backbeach. This beach type extends to Sabine Pass, which is the boundary with Louisiana. Beach erosion and frequent flooding have been so severe that the coastal highway is closed. Washover deposit of Hurricane Gilbert (1989) contains boulder size pieces of asphalt ripped up from the roadbed.

References Curray JR (1960) Sediments and history of Holocene transgression, continental shelf, northwest Gulf of Mexico. In: Shepard FP, Phleger, FB, van Andel TJ (eds) Recent sediments, Northwest Gulf of Mexico, American Association of Petroleum Geologists, Tulsa, OK, pp 221–226 Fisk HN (1959) Padre Island and the Laguna Madre flats, coastal south Texas. In: Proceedings 2nd Coastal Geography Conference, Louisiana State University, Baton Rouge, LA, pp 103–151 Hayes MO (1967) Hurricanes as geological agents: case studies of Hurricanes Carla, 1961, and Cindy, 1963. Report of Investigations 61, University of Texas, Bureau of Economic Geology, Austin, TX Kraus NC (2007) Coastal inlets of Texas. USA. In: Proc. Coastal Sediments’07, American Society of Civil Engineers, New Orleans, LA LeBlanc RJ, Hodgson WD (1959) Origin and development of the Texas shoreline. In: Proceedings 2nd Coastal Geography Conference, Louisiana State University, Baton Rouge, LA, pp 197–220 Morton RA (1979) Temporal and spatial variations in shoreline changes, Texas Gulf Coast. J Sed Petrol 49:1101–1111 Morton RA (2002) Factors controlling storm impacts on coastal barriers and beaches – a preliminary basis for real-time forecasting. J Coastal Res 18:486–501 Morton RA, McGowen JH (1980) Modern depositional environments of the Texas Coast. Guidebook 20, University of Texas, Bureau of Economic Geology, Austin, TX Morton RA, Paine JG, Blum MD (2000) Responses of stable bay-margin and barrier-island systems to Holocene sea level highstands, western Gulf of Mexico. J Sed Res 70:478–490 Morton RA, Miller TL, Moore LJ (2004) National assessment of shoreline change: part 1: historical shoreline changes and associated coastal land loss along the U.S. Gulf of Mexico. U.S. Geological Survey Openfile Report 2004–1043, Washington, DC Morton RA, Bernier JC, Barras JA (2006) Evidence of regional subsidence and associated interior wetland loss induced by hydrocarbon production, Gulf coast region, USA. Env Geol 50:261–274 Shepard FP, Phleger FB, van Andel TJ (1960) Recent sediments. Northwest Gulf of Mexico. American Association of Petroleum Geologists, Tulsa, OK

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1.6 Louisiana

Ioannis Georgiou · Mark Kulp · Michael Miner · Duncan FitzGerald

1.  Introduction The Louisiana coastal plain is one of America’s most unique and important areas in terms of regional coastal ecosystems, natural resources, human infrastructure and cultural heritage. Louisiana is located at the terminus of the Mississippi River, North America’s largest drainage basin stretching from the Appalachian Mountains westward to the foothills of the Rocky Mountains, covering a region of more than 3.3 million sq km (>Fig. 1.6.1). The Mississippi River discharges 6.2 × 1011 kg to the Gulf of Mexico annually (Coleman 1988) building a large delta plain (<30,000 sq km) that com-

prises more than one-third of southern Louisiana. Located to the west of the Mississippi Delta Plain, the Louisiana Chenier plain consists of a series of transgressive and regressive sandy ridges that are separated and underlain by prograded mud-rich deposits. Coastal Louisiana has been a long-term depository of sediment; much of the delta plain sits above a 10–15 km thick pile of Mesozoic and Cenozoic sediment. These strata overlie a highly attenuated crystalline basement that was stretched during Late Palaeozoic fragmentation of the Pangea landmass. The Mississippi River has been active since at least the Late Jurassic as a primary system contributing to the northern Gulf margin geology.

⊡⊡ Fig. 1.6.1 Louisiana coast showing the Mississippi River delta plain along the northcentral Gulf of Mexico.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The present day coast of Louisiana is undergoing unprecedented change. The morphology and integrity of the Louisiana coastline are directly related to the supply of sediment contributed to the coast from the Mississippi River and the physical processes acting on the coast. These processes include wave-generated longshore sediment transport, storm-driven sediment transport, bay and estuarine tidal exchange and circulation, inlet hydraulics, bay-wetland interaction, subsidence and eustatic sea level change. Historically, these processes have remained fairly constant with the exception of eustatic sea level, subsidence and sediment yield from the Mississippi River. Day-to-day coastal processes responsible for reworking the lower delta plain and modifying the coast are modest due to low wave energy (mean annual significant wave height of 0.50–0.95 m for the Chandeleur Islands and Chenier plain) and a microtidal regime (about 0.30 m mean annual tidal range). Tides along the Louisiana coastline change in a systematic manner with an overall decrease in tidal range from the western Chenier plain eastward towards Mississippi Sound (Georgiou et al. 2005). In southwestern Louisiana the tides are mixed having a strong diurnal component, whereas in the central region near the modern delta, tides are strongly diurnal. Tidal variations throughout the coast range from 15 cm during the neap tidal cycle to 100 cm during the spring cycle. Meteorological tides often have a profound effect on water levels in the area, especially during the fall and winter months when 10–30 frontal passages per year elevate water levels by 0.3–0.9 m (Boyd and Penland 1981). Associated with the passage of these frontal systems are southeast onshore winds that can generate waves with heights of 2.0–3.0 m. Subsequent to the passage of a frontal system, strong northerly winds are generated. Tropical cyclones, with a recurrence interval of 1.6 years, generate storm surges of 1.0–5.0 m, and deepwater wave heights can be as much as 5.0–10.0 m. Subsidence of the Mississippi River delta is related to natural processes such as sediment compaction, faulting and isostatic adjustment to regional crustal loading (Kolb and Van Lopik 1958; Törn­qvist et al. 2006) as well as anthropogenic effects that include subsurface fluid withdrawal and draining of wetlands for urban development (Morton et al. 2005). Historic sea level curves for the west and central coasts since the 1940s suggest relative sea level rates of the order of 1.03 cm/year. Recent data acquired from the NOAA tide gauge at Grand Isle show a slowing of relative sea level rise in recent years with a long-term (1947–2006) rate of 0.92 cm/year. This is the highest rate along the contiguous United States, which helps explain the rapid coastal evolution in Louisiana. The lack of

s­ ediment supply to the lower delta plain has led to erosion and drowning of marsh and other wetlands and the transgressive nature of most barriers in Louisiana. The historical decrease in size and segmentation of Louisiana’s barrier systems as well as the loss of wetlands has greatly increased the vulnerability of mainland areas to major hurricanes.

2.  Holocene Mississippi Delta Plain The Holocene Mississippi delta plain is the product of deposition by multiple, spatially and temporally offset, fluvio-deltaic deposition fed by distributary systems of the lower Mississippi River drainage basin. Chronological alterations in the location of these depositional centres arise from upstream avulsions of the Mississippi River and its distributaries. With each avulsion event, a new distributary network and attendant delta complex is formed. In total, the Holocene delta plain consists of six delta complexes: Maringouin (7,500–5,000 years bp), Teche (5,500– 3,800 years bp), St. Bernard (4,000–2,000 years bp), Lafourche (2,500–400 years bp), Balize (1,000 years bp – present) and Atchafalaya (400 years bp – present) (Roberts 1997). A delta complex represents the amalgamation of smaller delta lobes that are tied to a common distributary. On a temporal scale of a delta-complex construction, regressive deposition generally lasts about 1,000–2,000 years (Roberts 1997). Regressive deposition is recognised as an important process contributing to the vertically stacked and laterally offset deltaic depositional centres preserved within the shallow Holocene stratigraphic framework of the Louisiana delta plain. During delta progradation, wetlands and distributaries expand laterally (>Fig.  1.6.2), creating a landscape dominated by fluvial pathways, wetlands and bays between active distributary networks. Today, the delta plain consists of large expanses of coastal wetlands within a geomorphologic framework of lakes, estuaries and natural levee systems associated with active and abandoned distributaries. Delta plain formation is the result of the delta cycle, which starts with stream avulsion and delta building and ends with abandonment and delta reworking and deterioration by marine processes and subsidence (Colema­net al. 1998). The delta cycle can be divided into two phases,the fluvial-dominated regressive phase and the marine-­ dominated transgressive phase (Roberts 1997). Products of this autocyclic process of delta lobe progradation and subsequent abandonment are the transgressive components of the delta plain that include barrier island/tidal inlet systems, inner-shelf sand shoals, tidal channels and interdistributary bays (Roberts 1997).

Louisiana

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⊡⊡ Fig. 1.6.2 Reedswamp spreading at the mouth of a Mississippi River distributary. (Courtesy Geostudies.)

3.  Active Delta Plain The constructional phase of active delta growth is characterised by rapid infilling of lakes and bays followed by seaward progradation of the local coastline and the formation of a thick wedge of terrigenous clastic sediment on the shelf. Modern examples include the lacustrine deltas of the Atchafalaya Basin, the Atchafalaya/Wax Lake bayhead deltas (Roberts 1997) and the Balize “birdfoot” delta lobe. Progradational events begin with fluvial avulsion; stream capture from an active distributary. Major avulsion events are a product of decreasing gradient and hydraulic efficiency as the main channel lengthens by means of deltaic progradation. While many gradient advantages exist along the lower Mississippi floodplain, erodible substrate (e.g., abandoned channel sand) and favourable floodplain topography (e.g., active and abandoned floodplain channels) are also critical factors influencing avulsions (Aslan et al. 2005). If these criteria are met, a major river flood or hurricane may trigger avulsion, which results in stream capture by a previously less significant distributary and a delta-switching event, causing a gradual abandonment of the formerly active delta lobe in favour of the more efficient route. Subsequent to deltaic abandonment, previously active delta lobes become erosional headlands, and subsidence and marine reworking results in the landward migration of the shoreline (>Fig. 1.6.3). Sediment within the headland is reworked laterally by waves to form barrier islands and ultimately, inner-shelf shoals. The sandy barrier island

systems along coastal Louisiana constitute an important component of the delta-plain ecosystem because of the habitat they provide, their storm-surge buffering capabilities and their role in maintaining marine and estuarine gradients.

4.  Barataria Bay Barataria Bay is an interdistributary bay that has formed between the Lafourche and modern Placquemines/Balize delta lobe. The wetlands and fronting barrier system comprising this region illustrate well the processes of subsidence dominating abandoned delta plain regions. The Barataria barrier system is experiencing some of the highest relative sea level rise rates in the continental USA. During the past half-century rapid relative sea level rise (0.92 cm/year, Grand Isle tide gauge: 1947–2006) and erosional processes within Barataria Bay have led to substantial wetland loss, converting more than 1,100 sq km of wetlands to open water (16.9 sq km/year) since 1935. Conversion of wetlands to intertidal and subtidal environments is a product of several linked processes including subsidence, marsh front erosion and catastrophic scour during large magnitude hurricanes (FitzGerald et al. 2007). Other factors are also important in wetland loss, such as fluid withdrawal from nearby oil and gas fields (Morton et al. 2005), marsh edge erosion caused by local wind-generated wave erosion and marsh excavation during hurricane impacts (Barras 2006).

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Louisiana

⊡⊡ Fig. 1.6.3 Reworking of sediment by waves on the shore of the Mississippi delta. (Courtesy Geostudies.)

Long-term conversion of wetlands to open water (>125 years) has steadily increased tidal exchange between Barataria Bay and the Gulf of Mexico, resulting in larger inlet tidal prism. Two direct consequences of the increasing tidal discharge are the enlarging tidal inlet geometry and growth of ebb-tidal delta shoals (FitzGerald et al. 2004). During the 1880–2006 period of record the ­cumulative inlet cross section had quadrupled in size coincident with the formation of a new inlet (Pass Abel) and drastic increases in size of Caminada and Quatre Bayou Inlets. It has been shown that wetland loss, bay expansion and increased tidal prisms have strengthened tidal discharges through the tidal inlets and increased the competency of the tidal currents. The inlets have deepened, widened and increasingly transported sand-sized sediment further offshore, which together with relative sea level rise explains the loss of sand from the barrier system and a gradual deterioration of the barrier chain (FitzGerald et al. 2004).

5.  Delta Plain Stratigraphy The stratigraphy of the Holocene Mississippi River delta consists of generally fine-grained sediment deposited within a variety of fluvial, deltaic and coastal depositional environments. These sedimentary bodies formed in res­ ponse to deltaic progradation and abandonment, resulting in an assemblage of overlapping, stacked regressive and transgressive units that consist of unconsolidated sands and muddy sediment (>Fig.  1.6.4). During the waning phases of Holocene sea level rise the early and late Holocene delta plains were deposited as shelf-phase deltas. Trans­sgressive salt marsh, interdistributary bay, lagoon and transgressive  sands (>Fig.  1.6.4, green pattern) overlie regressive deltaic muds, peats and distributary deposits (>Fig. 1.6.4, grey pattern). As the shoreline migrates landward during the transgressive phase, there is a landward shift in paralic environments (brackish marsh, inner-distributary bay, saline marsh and lagoonal barrier). The marine flooding surface of a transgressive

Louisiana

    1.6

⊡⊡ Fig. 1.6.4 Idealised dip-oriented stratigraphic cross section of the Holocene Mississippi River delta plain. (With permission from ASCE.)

event is marked by a transition from freshwater marsh, swamp or fluvio-deltaic depositional environments to overlying paralic environments through open marine. The marine flooding surfaces ­separate parasequences within the transgressive systems tract. The Teche Ravine­ ment Surface coincides with the maximum flooding surface and separates the transgressive (early and late Holocene delta plains) from the highstand (modern delta plain) systems tracts. During the present highstand of eustatic sea level, delta switching led to the development of progradational parasequences separated by transgressive ravinement surfaces. The St. Bernard delta complex, which began to prograde about 4,000 year bp, is the first of the high-stand delta complexes (Boyd et al. 1989). Except for the modern Balize (Birdfoot) delta, which is deposited at the shelf edge, the Holocene deltas were restricted to relatively shallow inner-shelf waters. They produced thin (10–30-m thick) deltas of widespread coverage (as much as 15,000 sq km). By contrast, the Balize shelf-edge delta is small in extent but consists of a regressive sedimentary interval nearly 120-m thick (Coleman 1988; Kulp et al. 2002).

6.  Chenier plain The Chenier plain (French chêne) of SW Louisiana extends from Sabine Pass, Texas, to Southwest Point Louisiana, with maximum elevations of about 2–6 m (>Fig. 1.6.5). It consists of multiple, shore-parallel, sand-rich ridges that are perched on and physically separated from one another by relatively finer-grained, clay-rich sediment (Penland and Suter 1989). The Chenier plain evolved during the Holocene as a series of progradational mudflats that were intermittently reworked into sandy or shelly ridges to form the modern Chenier-plain physiography. Numerous cycles of deposition and erosion created alternating ridges separated by marshlands. The mudflat component of the plain has ­been built from fine-grained sediment discharged by the Mississippi River and transported westward. Reworking of the mudflat coincided with periods of non-deposition due to an eastward shift of the Mississippi River deposition centres. During these periods coarse-grained sediment and shells are concentrated along the shore forming Chenier ridges. Ridges become isolated when a new ­supply

65

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Louisiana

⊡⊡ Fig. 1.6.5 The Chenier plain in southwestern Lousiana (top) showing the distribution of cheniers (dark lines) along the coastal plain, and (bottom) the processes of Chenier plain formation. (Cheniers (dark lines) are from Gould and McFarlan, 1959)

of sediment from the Mississippi River distributaries renews mudflat progradation. Repeated seaward growth and retreat along the Chenier plain is thus a consequence of deltaic deposition farther east and the periodic cessation of sediment supply to the Chenier plain as the deltaic depositional centres become abandoned. Currently, the Atchafalaya River is supplying the Chenier plain with fine sediment and mud and parts of the mudflat coast are presently prograding.

7.  Impact of Rising Sea Level and Storms Currently, southern Louisiana is the site of the highest rates of coastal erosion and wetland loss in the United States, including as much as 75 sq km/year of interior land loss (Barras et al. 2003) and up to 23.1 m/year of coastline erosion for select time intervals. Clearly, southern Louisianal is an ecosystem in crisis. Recent estimates of global sea level rise suggest an average rise of between 0.18 and 0.59 m by the end of this century (IPCC 2007). Added to this projected rise is the high rate of local sea level rise due to subsidence, which undoubtedly will

be disastrous for this low lying area of the northern Gulf of Mexico. Between 2000 and 2008, the Mississippi River delta plain has been the landfall of more than seven hurricanes. Four of these storms brought category 3 and higher wind speeds (>178 km/h) and storm surges (>5 m) to the delta plain. These landfalls have been devastating to this already highly ­fragile system producing extensive marine inundation, widespread barrier island erosion, removal of interior marshlands to create open water areas and degradation of the many flood protection systems that protect the heavily populated delta plain. These effects, in concert with decades of anthropogenically created land loss, have created widespread concerns for the future productivity of the delta-plain ecosystems and the ability of the continually diminishing coastal area to provide the resources necessary to the culture and economy that so heavily rely upon the uniqueness of this region. In response to the predicted drastic changes of the delta plain during the next century, numerous land-loss mitigation and management projects have been implemented and many more are proposed in an effort to protect and even revitalise the system. Successful management

Louisiana

of this transgressive coastal zone requires careful attention to how sediment resources can most efficiently be used so  as to replenish depleted shorelines and how riverine diversions can create localised delta plain growth. Approaches undertaken in the emerging concept of Transgressive Management for the Mississippi River delta plain may establish the guiding principles for dealing with coastal zones undergoing similar changes.

References Aslan A, Autin WJ, Blum MD (2005) Causes of River Avulsion: insights from the late Holocene avulsion history of the Mississippi River, U.S.A. J Sed Res 75:650–664 Barras JA (2006) Land area change in Coastal Louisiana after the 2005 hurricanes: a series of three maps. U.S. Geological Survey Open-File Report 2006–1274 Barras J, Beville S, Britsch D et al (2003) Historical and projected coastal Louisiana land changes: 1978–2050: United States Geological Survey, Open File Report, 03-334, 39p (Revised January 2004) Boyd R, Penland S (1981) Washover of deltaic barriers on the Louisiana Coast. Gulf Coast Assoc Geol Soc Trans 31:243–248 Boyd R, Suter J, Penland S (1989) Relation of sequence stratigraphy to modern sedimentary environments. Geology 17:926–929 Coleman JM (1988) Dynamic changes and processes in the Mississippi River Delta. Geol Soc Am Bull 100:999–1015 Coleman JM, Roberts HH, Stone GW (1998) Mississippi River Delta, an overview. J Coastal Res 14:698–716 FitzGerald DM, Kulp MA, Penland S, Flocks J, Kindinger J (2004) Morphologic and stratigraphic evolution of muddy Ebb-Tidal deltas along a subsiding coast: Barataria Bay, Mississippi River Delta. Sedimentology 51:1157–1178

    1.6

FitzGerald DM, Kulp MA, Huges Z, Georgiou I, Miner MD, Penland S, Howes N (2007) Impacts of rising sea level to backbarrier wetlands, tidal inlets, and barrier islands: Barataria Coast, Louisiana: Coastal Sediments ’07. Proceedings of the 6th international symposium on coastal engineering and science of coastal sediment processes, New Orleans, LA, 1179–1192 May 13–17 Georgiou IY, FitzGerald DM, Stone GW (2005) The Impact of physical processes along the Louisiana Coast. J Coastal Res, Special Issue 44:72–89 Gould HR, McFarlan E Jr (1959) Geologic history of the chenier plain, southwestern Louisiana. Trans Gulf Coast Assoc Geol Soc 9: 261–270 IPCC (2007) Climate change 2007: the physical science basis, summary for policymakers. contributions of working group I to the fourth assessment report of intergovermental panel on climate change Cambridge, UK Kolb CR, Van Lopik JR (1958) Geology of the Mississippi River deltaic plain, southeastern Louisiana, US Army Corps of Engineers Waterways Experiment station, Vicksburg, MS, Technical report No. 3-483 Kulp MA, Howell P, Adiau S, Penland S, Kindinger JL, Williams SJ (2002)  Latest quaternary stratigraphic framework of the Mis­ sissippi  River delta region. Gulf Coast Assoc Geol Soc Trans 52:573–82 Morton RA, Bernier JC, Barras JA, Ferina NF (2005) Rapid subsidence and historical wetland loss in the Mississippi delta plain: likely causes and future implications: U.S. Geological Survey Open-File Report 2005–1216, 115p Penland S, Suter JR (1989) The geomorphology of the Mississippi River chenier plain. Marine Geol 90:231–258 Roberts HH (1997) Dynamic changes of the Holocene Mississippi River Delta Plain: the Delta Cycle. J Coastal Res 13:605–627 Törnqvist TE, Bick SJ, van der Borg K, de Jong AFM (2006) How stable is the Mississippi delta? Geology 34:697–700

67

1.7 Mississippi

Ervin Otvos

1.  Introduction

basin has an area of 22,470 sq km that yields an average discharge of 282 m3/s. The Pascagoula-Leaf watershed, 24,600 sq km in area, has 354 m3/s discharge. Small estuaries occur in the mouths of the Biloxi, Jourdan, and Wolf Rivers, Graveline Bayou, and numerous smaller tidal creeks with minor freshwater discharge. Straddling the Alabama state line on the east are rapidly decaying marshlands that formed when the abandoned Escatawpa River delta filled the head of Grand Bay. With its freshwater source lost, the bay ceased to be a separate estuary. Mississippi Sound, the largest estuarine system, is an unusually wide lagoon between a barrier island chain

The mainland coastline of Mississippi is slightly curved and indented by embayments with small bayhead deltas (Otvos 1981). The hinterland has a series of emerged Quaternary terraces. Late Holocene marshlands and beaches fringe the estuarine shores of the mainland, and an extensive marshland on the SW coast surrounds remants of stranded late Holocene relict islands sev­ eral metres high (Otvos 1998). The largest estuaries are in the lower Pearl and Pascagoula valleys, as well as St. Louis Bay and Biloxi Back Bay. The Pearl drainage

sca

Pa

R.

an

R.

lf

o

Pascagoula

Gulfport

Bay St. Louis

Biloxi

SIPPI MISSIS 4m

Pea rl R.

Cat Island

SOUND

Ship Island

Horn Island

4m

LOUISIAN

R.

W

Escatawpa

Jourd

30 Km

B il o x i R .

20

goula R.

10

0

MISSISSIPPI ALABAMA

⊡⊡ Fig. 1.7.1 Major features on the Mississippi coast.

A

4m

Petit Bois Island

m

10

4m

20m

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.7.2 Generalised geological map of the Mississippi coastal plain. N MISSISSIPPI

LA

Big

Harrison Co.

Hancock Co.

arp Sc e Jackson Co. dg

Ri

AL

te-10

Intersta

Biloxi

Belle Fontaine Point

Slidell

Gautier Pascagoula

Holocene Pleistocene Port Bienville

Gulfport Fm. Prairie Fm.

MS

Pre-Sangamonian

LA 0 0

10 10

20 ml 20

Pliocene

Citronelle Fm.

30 km

⊡⊡ Fig. 1.7.3 Generalised cross-section of the Mississippi coastal plain from the Citronelle Upland across faulted Pleistocene formations to the Late Pleistocene and Holocene formations of the coastal fringe.

N

S CITRONELLE UPLAND

CITRONELLE SCARP

BIG RIDGE SCARP

BACK BAY

BILOXI

MISSISSIPPI SOUND

20 m 10 m 0 –10 m

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PRE-SANGAMON PLEISTOCENE CITRONELLE FM. (PLIOCENE)

GULFPORT FM.

LOWER PLIOCENE

PRAIRIE FM.

3 km

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BILOXI FM.

LATE PLEISTOCENE

Mississippi

and the mainland. Within Mississippi, the barrier island chain includes Cat, West and East Ship, Horn and Petit Bois islands (>Figs. 1.7.1 and > 1.7.2).

2.  Coastal Terraces Rising as high as 120 m south of Jackson in southern Mississippi, the deeply dissected late Pliocene Citronelle upland surface dominates the coastal plain. Near its southern margin along the mainland coast it stands at 15–25 m (>Fig. 1.7.3). Bordering the Citronelle upland Gulfward are two sizable remnants of the late Pleistocene Montgomery terrace occur at 12–15 m elevation in the central and eastern Mississippi coastal plain (Otvos 1998).

3.  Holocene Coastal Landforms An initially rapid, subsequently decelerating, sea level rise determined coastal development on the northern Gulf (Otvos 2004). Indications of higher than present Holocene

    1.7

sea levels have not yet been reliably identified in Missi­ ssippi. Inundation of the mainland was followed by the gradual infilling of river and creek valleys and widespread wetland development. Marshlands formed in areas of low salinity up to 20 km upstream from the present Pearl River delta front. Brackish marshes were established on the ancient banks of the Biloxi River as the drowning valley became the Biloxi Back Bay. Meandering marsh islands outline the former river channel in the bay centre. The former mainland coast is indicated by the narrow 2–4 m high Magnolia Ridge, composed of dunes landward of the southern Hancock County marsh in SW Mississippi (>Fig. 1.7.4). A bifurcated dune-covered 4 km long narrow sand spit prograded about 2,500 year bp from the Gulfport shore of Belle Fontaine Beach in the southeast.

4. The Barrier Island Chain and Mississippi Sound Westward littoral sand transport from Dauphin Island created a line of elongated narrow sandy shoals, on which

⊡⊡ Fig. 1.7.4 South Hancock Marshland. Point Clear Island is a relict barrier island stranded within marshland. (Courtesy NASA Landsat 7.)

Prairie Surface

Magnolia Ridge

Magnolia Ridge

Lakeshore Bayou Caddy

Pt. Clear Island Campbell Island

SOUTH HANCOCK MARSHLAND

Heron Bay SJP

Mississippi Sound

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Mississippi

⊡⊡ Fig. 1.7.5 Lake Borgne and the relict St Bernard delta lobe, showing the South Hancock marshlands and shoals (SHS). (Courtesy NASA Landsat 7.)

Pleistocene upland

St. Louis Bay

SHS

Pearl River Estuary Lake Pontchartrain

buried barrier islands

Stranded Barrier Islands

south Hancock marshlands

Lake Borgne

ancestors of the present barrier islands aggraded, prograding both westward and gulfward since about 4,600–4,400 years ago (Otvos and Giardino 2004). Two relict islands, Point Clear and Campbell, now partially buried, mark the previous westward continuation of the barrier chain in the present SW Mississippi marshland. Prograding from below Magnolia Ridge, they formed soon after 4,400 years bp (>Figs. 1.7.4 and >1.7.5). Before being stranded and partly

Mississippi Sound

Relict Mississippi-St. Bernard Delta

buried, the chain extended across SW Missis­sippi well into SE Louisiana (Otvos and Giardino 2004). Shoaling was related to the growth of a new Mississippi River sub-delta to the south, which stopped island development west of Cat Island about 3900–3700 years ago. Emergence of the island chain separated the large Mississippi Sound from the Gulf. Eroded remnants of a late Holocene Mississippi River sub-delta form the SW shore of the Sound. Within

Mississippi

    1.7

⊡⊡ Fig. 1.7.6 Central and East Mississippi coast. (Courtesy NASA Landsat 7.)

Biloxi

Gulfport

Grand Bay Deer Island

Back Bay Biloxi Bay

Mississippi Sound Cat Island

Belle

Pascagoula R. delta

Fontaine

Relict delta

Pt. aux Chenes Bay

Spoil Island

Camille Cut Georges Cut

Horn Island Petit Bois Island

West & East Ship Islands

Gulf of Mexico ⊡⊡ Fig. 1.7.7 Beach ridges on Cat Island have been truncated and bordered by a dune­covered barrier spit trending SW to NE. (Courtesy United States Geological Survey.)

Mississippi, the Sound is about 112 km long, 10–22 km wide (>Fig. 1.7.6), and generally 2.0–4.0 m deep. Water depths exceed 6 m in and near tidal channels between the islands. While sandy deposits fringe the mainland and island shores, muddy bottom sediment dominate the lagoon floor. The barrier islands between SW Alabama and SE Louisiana originated by westward longshore drift of sand from a large ebb-tidal delta south of Mobile Bay. From about 4,600–4,400 years bp, the narrow islands emerged through vertical accretion on elongated shoals (Otvos and Carter 2008). Sand between the islands was transmitted by way of the elongated shore-parallel sand shoals and small

ebb-tidal deltas located off inlets. The islands expanded through westward and seaward strandplain growth. Cat Island grew 3 km seaward, but westward progradation dominated the other barrier islands. Erosion of the eastern ends of islands was driven both by fair-weather processes and occasional storms. Wide breaches that formed across narrow low sectors during storms led to the revival of sandy shoals. These processes have temporarily or permanently segmented and eliminated extensive barrier island sectors during the past three centuries (Otvos and Carter 2008). Details of changing patterns of erosion and accretion are given in the electronic version of this chapter. They

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⊡⊡ Fig. 1.7.8 The changing coastline of Petit Bois and Dauphin Islands, 1848–1974.

MISSISSIPPI SOUND 1974

PETIT BOIS PASS 1974

1974

PASS IN 1848

1848

DAUPHIN IS.

1848 PETIT BOIS IS. 0

5 km

0

3 ml

GULF OF MEXICO

⊡⊡ Fig. 1.7.9 Belle Fontaine Beach area. (Courtesy NASA.)

Mississippi

    1.7

⊡⊡ Fig. 1.7.10 Erosion of East Belle Fontaine Bluff during a high spring tide in 1987. The dark material on the shore is humate from the Gulfport Formation.

include truncated beach ridge systems and successive ­outgrowths (>Figs. 1.7.7–1.7.10), the growth and destruction of deltas, and a summary of the historical development of the ­barrier islands.

5.  Hurricane Impacts With all hurricane categories considered, the average recurrence interval within 155 km of the Mississippi shoreline averaged 9–10 years. Hurricanes of Category 3 and above recurred once every 32–36 years on average. About 48 hurricanes and tropical storms have had at least a minor impact on the Mississippi coast since 1838. They included Category 3 cyclones Camille (1969) and Katrina (2005). Camille produced a maximum surge of 6.8 m at landfall in the west, while Katrina generated high flood levels in the western islands and on the adjacent mainland coast, reaching unprecedented maxima

of 10–12 and 7–8 m, respectively. They caused rapid erosion and overwash, notably on Cat and Horn islands (Otvos and Carter 2008). Beach erosion along the mainland coast has resulted in periodic beach nourishment.

References Otvos EG (1981) Barrier island formation through nearshore ­aggradation – stratigraphic and field evidence. Mar Geol 43:195–243 Otvos EG (1998) Citronelle Formation, northeastern Gulf coastal plain: Pliocene stratigraphic framework and age issues. Trans Gulf Coast Assoc Geol Soc 48:321–333 Otvos EG (2004) Holocene Gulf levels: recognition issues and an updated sea-level curve. J Coastal Res 20:680–699 Otvos EG, Carter GA (2008) Hurricane degradation-barrier development cycles, northeastern Gulf of Mexico: landform evolution and island chain history. J Coastal Res 24:463–478 Otvos EG, Giardino MM (2004) Interlinked barrier chain and delta lobe development, northern Gulf of Mexico. Sed Geol 169:47–73

75

1.8 Alabama

Ervin Otvos

1.  Introduction Alabama has an embayed mainland coast with broad lagoons, including Mobile Bay and Perdido Bay, behind a sandy barrier island (Dauphin Island) that continues eastward as the Morgan Peninsula, extending to the Florida border (>Fig. 1.8.1). Much of the coastal plain is occupied by the dissected Upland Surface, underlain by late Pliocene deposits. Bluff-bounded, deeply incised wide valleys in the Mobile-Tensaw River system are associated with extensive valley-floor wetlands and a large marshy delta at the head of Mobile Bay. The narrow Holocene coastal plain east of Mobile Bay includes a beach ridge complex, interspersed with narrow lagoons, inlets, paralic lakes and marshes. The outer coast is bordered by a sandy barrier beach, 20 km long and 2–4 km wide, with dunes on the Morgan Peninsula. A large compound barrier spit at the mouth of the bay has narrowed the entrance to Mobile Bay.

In addition to Mobile Bay, there are several smaller estuaries, including Perdido Bay, Fowl River, Fish RiverWeeks Bay and Bayou la Batre. On the inner side of the entrance to Mobile Bay is a large cuspate sandy ebb-tidal delta. To the west, an extension of Mississippi Sound lies between the Dauphin barrier island and the mainland coast.

2.  Coastal Plain The late Pliocene Citronelle depositional upland surface is the oldest and most dominant element of the coastal plain. It is primarily alluvial, but portions were deposited under estuarine conditions (Otvos 2004). It is thought to have been raised by post-Pliocene tectonic uplift. Hundreds of small circular and elliptical depressions, probably of aeolian origin, dot this upland surface.

⊡⊡ Fig. 1.8.1 Major features on the Alabama coast.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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An erosional scarp forms the boundary between the Citronelle upland and the low, gently seaward sloping Prairie alluvial surface. The Prairie plain widens along the western side of Mobile Bay. It includes remnants of Pleistocene barrier islands.

3.  Holocene Evolution Dauphin Island and the adjacent Mississippi barrier chain formed, when the rising sea encroached on and surrounded an isolated elevated remnant of the Gulfport barrier SW of Mobile Bay. Large sand volumes supplied by westward littoral transport from the adjacent large ebb tidal delta built an apron around this barrier remnant.

A line of shoals formed in the nearshore zone and barrier islands were built on them by vertical aggradation (Otvos and Carter 2008). The rising late Quaternary marine transgression gradually drowned the Mobile Basin and smaller valley mouths to the east. Barrier islands then formed (Otvos 2005), enclosing wide lagoons. Morgan Peninsula consists of two obliquely aligned beach ridge, sets overlain by a substantial dune cover (Rodriguez and Meyer 2006). Gradually, constricting the wide entrance to the Bay in the late Holocene (>Fig. 1.8.2), the Morgan barrier spit has slowly extended westward. It has also prograded, forming a series of lobate forelands bearing successively formed beach ridges that have been truncated by subsequent marine erosion (>Fig. 1.8.3).

⊡⊡ Fig. 1.8.2 Mobile Bay and adjacent areas. (Reproduced with permission from Landsat 7.)

Alabama

1.8

⊡⊡ Fig. 1.8.3 Holocene beach ridges in the Morgan-Perdido area.

4.  Mobile Bay The Mobile Bay drainage basin (115,513 sq km) contains the fourth largest river system in the United States. Fluvial discharge from the Mobile is 2,246 m3/s. Bell-shaped Mobile Bay has a surface area of 1,070 sq km and is 50 km long, 12 km wide in the north, 38 km wide in the south. On the northeast, the bay is flanked by 20–30 m high Pliocene bluffs capped by silty sands of the Citronelle Formation that forms the upland surface. The late Pleistocene Prairie terrace alluvium and Holocene deposits flank the bay shore on the SW, W, and SE. Estuarine and fresh water deposits filled a 64 km long stretch of the incised valley, north of the present Mobile River delta. Oyster reefs were abundant. Salt content ranges between 0.9 and 2.8% in the Mobile Bay, and the eastern extension of Mississippi Sound. Microtidal ranges of 0.4–0.7 m prevail, and ebb-tide currents in the deep, narrow inlet channel attain 5.5 km/h. Construction of numerous large reservoirs on tributary streams reduced fluvial sediment transport to the Bay in the twentieth century. 4.8 million tons of suspended and bottom sediment now reach the delta yearly. About 45% of the fluvial sediment discharged, passes through the Delta, but remains in Mobile Bay, while 30% enters the Gulf of Mexico by the deep main tidal pass and Mississippi Sound by way of Grants (Heron) Pass. The 5.4 m deep outlet channel in the southwestern Bay deepens to 17.5 m in the main tidal pass toward the Gulf. Extensive dredging and deposition of sediment since the second half of the nineteenth century has led to radical changes in bottom topography. There have been many changes in and around the entrance to Mobile Bay and on the associated ebb-tide delta. During Hurricane Frederic (1979) small islands were drastically reduced or completely eliminated. Re-emerging from sand shoals, the small islands expand quickly in storm-free years (>Fig. 1.8.4). Details of historical changes on Dauphin Island are given in the electronic version of this chapter. There have been erosive episodes of variable duration on the southern beaches (>Fig.  1.8.5), while two small cuspate forelands prograded also on the sheltered southeastern island shore (>Fig. 1.8.6). The wave energy shadow created by the huge Mobile Pass ebb-tidal delta, affords the shore sector with limited protection. Sets of multiple sand bars parallel the Gulf shoreline in the nearshore zone. Perdido Bay, a drowned valley, is a 7 km long, maximum 6.5 km wide estuary on the Florida state border. Its average depth is 2.4 m. Clayey silt and silty clay cover the bottom areas, and well-sorted clean sands fringe the bay shores.

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Alabama

⊡⊡ Fig. 1.8.4 Changes in islands and shoal configuration on the Mobile Bay ebb-tidal delta between 1908 (top) and 1936 (bottom).

Alabama

1.8

⊡⊡ Fig. 1.8.5 Localised erosion at Fishing Pier, east Dauphin Island. Erosion due to changes in wave sheltering and sand transport, influenced by shifting positions of Sand Island, just offshore and on the western flank of the Mobile ebb-tidal delta. Remnants of pine forest previously buried by dunes, then exposed by 1993–1994 coastline retreat. The boardwalk in the foreground now leads nowhere.

⊡⊡ Fig. 1.8.6 Two small cuspate forelands on the southeastern coast of Dauphin Island, prograded in the wave-shadow of an adjacent ebb-tidal delta. (Courtesy Scott L. Douglass.)

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References Otvos EG (2004) Lithofacies and depositional environments of the Pliocene Citronelle formation, Gulf of Mexico coastal plain. Southeast Geol 20:1–20 Otvos EG (2005) Coastal barriers, Gulf of Mexico: Holocene evolution and chronology. J Coastal Res SI-42:141–163

Otvos EG, Carter GA (2008) Hurricane degradation-barrier development cycles, northeastern Gulf of Mexico: Landform evolution and island chain history. J Coastal Res 24:463–478 Rodriguez AB, Meyer CT (2006) Sea-level variations during the Holocene deduced from the morphologic and stratigraphic evolution of Morgan Peninsula, Alabama, U.S.A. J Sed Res 76:257–269

1.9  Florida

Charles Finkl

1.  Introduction The nature of the Florida coastline is about 3,663 km long. Its landforms are closely tied to conditions on the continental shelves off the Gulf and Atlantic and coasts. These can be divided into six shelf zones, each associated with energy segments and sediment associations that influence coastal configuration and the character of the shore. Tanner (1985) noted that wave energy was generally low on the Gulf coast (mean breaker height <4 cm, sometimes called zero wave energy) and higher on the Atlantic coast (mean breaker height about 1 m). Tide range is typically less than a metre. He considered the nature of Florida beaches, rock control of shoreline configuration, and the distribution of coastal sediment both alongshore and cross-shore (>Fig. 1.9.1). The coastline of Florida is well known for its exten­­ sive  sandy beaches, which have a total length of almost 1,300  km. These beaches provide protection from storm surges along developed shores, present valuable habitat and contribute to the economic prosperity of tourism. The beach-dune system is thus a critical component of state infrastructure that requires periodic maintenance through beach renourishment. Florida beaches range from small pocket beaches in the Florida Keys to long unbroken segments that stretch for many kilometres. Navarre Beach is the longest continuous stretch of protected beach (13 km in the Gulf Islands National Seashore) in northwest Florida. Panama City Beach and Miami Beach are among the longest renourished beaches in Florida. According to the morphodynamic beach classification system Florida Atlantic beaches exemplify dissipative and intermediate beach states. Reflective beach states occur immediately downdrift (southern margins) of stabilised inlets, but are of limited extent, rarely spanning more than a kilometre or two of coastline. Gulf coast beaches tend to be dissipative, except along the Panhandle where more intermediate beach states occur. Florida beaches show compositional variation from pure quartz to carbonate with a complete range of intergrades between these. In general, beaches north of the 28th parallel on both coasts tend to be quartzose while

those to the south are of carbonate sand. Shells and shell hash are common to all beaches in Florida, regardless of the sand-size composition. Berm grain sizes tend to range from 0.25 to about 0.35 mm on Atlantic coast beaches (>Fig.  1.9.2) whereas Gulf coast beaches show both coarser and finer grain sizes depending on the amount of contained shells. Fine sand beaches with particle size distributions commonly falling in the range of 0.15–0.20 mm are often referred to as sugar sand beaches. These find sand beaches are considered excellent for sand sculpting. Along some parts of the beach, usually on the back berm near the dunes, the sand is so clean that when dry it squeaks when walked upon. Coarser particle sizes may occur under specialised conditions where there are many shells or where there is a concentration of lag deposits on eroded remains of beaches when the finer grains have been washed away. Berms on these coarser-grained beaches often show grain sizes in the range of ­0.35–4.0 mm. Most Florida beaches were backed by low dunes, but these have been removed along developed shores. Beach colour ranges from light grey, due to mixtures of shell fragments with varying contents of quartz and carbonate grains, to pure white, as in the case of the west coast sugar sand beaches, and black and dark grey sands near Venice on the southwest coast to yellow in Flagler County. White sand is due to quartz grains whereas the black sand is due to the incorporation of fragmented fossils in the beach. Yellow sand is related to the adsorption of iron oxides by coquina shells that become mixed with the quartz sand. Darker colours can also be associated with organic materials mixed with sand on eroded beaches where the barrier has retreated over marsh sediment, as sometimes occurs along the Florida Panhandle. Beach erosion has been a problem: about 225 km of Florida beaches were classified as critical erosion areas in 1985 (FNDR 1985), while 370 km were critically eroded in 1993 (Clark 1993), the difference being related to increased coastal development, changing environmental conditions, and better assessment techniques. Florida has successful beach nourishment programmes on both Gulf and Atlantic coasts. The Atlantic coast has about 50 nourished areas that together have received

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.9.1 The coast of Florida, showing breaker zone wave energy levels (see text) and marine terraces at 80, 50 and 35 m above sea level. (Courtesy W.F. Tanner.)

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rocky coasts (wave-cut cliff, wave-cut platform) large offshore shoal, measured in kilometers scattered reefs and patches of vermetid gastropods living coral reef spectacular exposures of sabellariid worm reefs in the intertidal zone scattered sabellariid worm reefs well-developed beach ridge plains of late Holocene age – from west to east: St. Vincent Island, Dog Island, Sanibel Island, Marco Island (in good part destroyed by urbanization), and Cape Canaveral. dune fields the +1.5-m mid-Holocene high-sea-level shoreline growing deltas Pre-Holocene delta, perhaps Sangamon in age point from which the wave-cut platform radiates outward in three directions (E, S, W). beach-ridge plains, pre-Holocene in age, probably Sangamon beach ridge plains, perhaps Sangamon but perhaps older Pleistocene old coalesced sets of beach ridges, poorly preserved, perhaps Pliocene in age

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⊡⊡ Fig. 1.9.2 Trend of composite beach grain sizes for the east coast of Florida, showing a general increase in grain size from north (Volusia County) to south (Miami-Dade County). Moving average trends smooth out station variations and the red linear regression line shows the overall longshore trend.

Florida

up  to  65 million m3 of sediment since the mid 1940s. Miami Beach, built from 1978–982, was the largest single construction event in the history of beach nourishment on the U.S. East Coast with about 9.2 million m3 of sediment dredged from several different borrow areas located in inter-reefal sediment troughs. The success of the Miami Beach project may be attributed to the long extent of the nourished area (17 km) that reduced fill spreading rates, relatively low wave energy, and the fact that longshore drifting ends at a very long downdrift jetty where the sand accumulates. Some coastal segments such as in the Florida Keys, are predominantly rocky with small pocket beaches that are only occasionally renourished with sand from upland sources. Still other coastal areas such as the Big Bend (Taylor, Dixie, Levy counties) are mostly marshlands or tropical mangroves as along the southwest shore of mainland Monroe County on the southern tip of the peninsula. Rocky outcrops are rare, but have a strong influence on coastline configuration. This is more obvious on the Atlantic coast. where barrier islands are related to bedrock highs, where the Cape Canaveral cuspate foreland is determined by bedrock shoals offshore, where coral reef tracts follow bedrock lineaments, and where beaches are perched on older marine benches and platforms that were formed during higher Pleistocene sea levels. In northern Palm Beach County and throughout Martin County the Anastasia Formation outcrops as low sea cliffs and shore platforms. Beach profiles are strongly influenced by bedrock outcrops on the seafloor. The shape of surveyed beach profiles along Broward County is determined by the position of the first rock outcrop offshore. Between Miami and Palm Beach Robertson et al. (2008) found that rock outcrops and structural sand flats, as mapped by Finkl et al. (2005), influenced the point beyond which there was no significant sediment transport. Other examples of rock control are associated with reef systems (colonies of worm casts) built by sabellariid worms in surf and subtidal environments along the Atlantic coast.

2.  The Coastline The Gulf coast of Florida between the Alabama state line and Levy County (about 520 km) consists of beaches backed by sandy beach and dune ridges. The western part of this coastal segment is referred to as the Florida Panhandle and the eastern part as the Big Bend. This zone is noted for its white sand barrier island beaches, and stands behind a shelf sector that bears sand ridges, buried channels, and shoreface sand sheets, with numerous

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ebb-tidal deltas and offshore shoals. The Apalachicola delta is bordered by barrier islands with multiple beach ridges indicating stages in growth and truncation. St. Vincent Island, for example, has several sets of beach ridges truncated on its eastern shore. Intermittent progradation has been related to sea level alternations during the past 4,000 years, but has given place to predominant erosion on these shores (Tanner and Stapor 1972). To the east is Apalachee Bay, where wave energy is generally very low because prevailing winds blow from land to sea, the concavity of the coastline provides for divergence of wave orthogonals and hence energy reduction, the sea floor is unusually shallow and wide, so that essentially all deep water wave energy is dissipated by frictional processes in the passage across the shelf. Between the southern margin of the Big Bend in Dixie County and the northwest flank of Florida Bay the coast (about 485 km) encompasses sand ridges and nearshore sand flats. Incident wave energy increases southward from low to moderate (mean breaker height rising to about 50 cm). This coastal segment has barrier islands and lagoons, and offshore there are drowned karstified limestones surmounted by nearshore sand sheets. Ebb-tidal deltas are associated with major estuaries (such as Tampa Bay and Caloosahatchee River). Sand ridges tend to be obliquely oriented to the coast, although shore-parallel and shore transverse ridges occur in restricted locations. Charlotte Harbour is a large embayment protected by a chain of ­barrier islands that extends south to a landward curve at Sanibel Island. To the south the coast is beach-fringed behind Estero Bay, but beyond Naples it becomes irregular, fringed by the Thousand Islands. This is a soft-sediment coastline with mangrove forests that grade inland to estuarine areas and finally Everglades marsh. At Cape Sable the coastline turns eastward behind Florida Bay and the long chain of islands (the Florida Keys) that run out westward to Key West. Florida Bay, encompassing about 2,100 km2, is a shallow (about 1–2 m deep) inner shelf lagoon-like environment, with more than 200 mangrove islands, interconnected basins, and grassy mud banks. The Florida Keys have a rocky Atlantic coastline with low bluffs (up to 2 m high) and pocket beaches. Sand flats and sea grass meadows occur in the channel between the Keys and the Florida Reef Tract offshore. Mangroves occupy sheltered environments, but north of Biscayne Bay the Atlantic coast is fringed by a sandy beach. In this 220 km sector the continental shelf is narrow, being only about 1.5 km wide off central Broward County. The outer edge has a zone of coral reefs that extends northward from

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the Florida Keys. The most prominent geological and geomorphological features along the southeast Florida coast include sand flats, nearshore sand bars and rocky reefs. North of Palm Beach is a 170 km sector of sandy beaches, interrupted by tidal inlets. Some of these have shoals offshore (>Fig. 1.9.3). Southward longshore drifting is indicated by accretion against obstacles (>Fig. 1.9.4) and on the northern side of jetties (>Fig. 1.9.5).

The large Cape Canaveral cuspate foreland has numerous beach ridges indicating stages in its evolution (CraigSmith 2005). Prominent transverse bars and sand ridges round the cape offshore. From Daytona Beach to the Georgian border (about 235 km) sandy beaches are backed by dunes. The continental shelf has large shore-parallel sand ridges, sand waves, sand banks and shoals, with an extensive nearshore ⊡⊡ Fig. 1.9.3 Sebastian Inlet, on the Atlantic coast of Florida. (Courtesy J. Morelock.)

⊡⊡ Fig. 1.9.4 Accretion of southwarddrifting sand alongside a grounded ship at Rivera Beach. (Courtesy J. Morelock.)

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⊡⊡ Fig. 1.9.5 Accretion of southwarddrifting sand alongside a jetty at Port Canaveral. The cuspate foreland Cape Canaveral is in the background. (Courtesy J. Morelock.)

sand sheet. Sand is carried onshore to the beach by constructive swell in fine weather, and withdrawn to the sea floor during storms.

References Clark RR (1993) Beach conditions in Florida: a statewide inventory and  identification of the beach erosion problem areas in Florida.  Tallahassee: Florida Department of Environmental Pro­ tection, Beaches and shores technical and design memorandum, 202, pp 89–91

Craig-Smith SJ (2005) Cuspate forelands. In: Schwartz ML (ed) The Encyclopedia of Coastal Science, Dordrecht, Kluwer, pp 354–355 FDNR (Florida Department of Natural Resources) (1985) Beach restoration. A state initiative. Tallahassee, Division of beaches and shores Finkl CW, Benedet L, Andrews L (2005) Submarine geomorphology of the continental shelf off southeast Florida based on interpretation of airborne laser bathymetry. J Coastal Res 21:1178–1190 Robertson W, Zhang K, Finkl CW, Whitman D (2008) Hydrodynamic and geologic influence of event-dependent depth of closure along the south Florida Atlantic coast. Mar Geol 252:156–165 Tanner WF (1985) Florida. In: Bird ECF, Schwartz ML (eds) The World’s Coastline, Van Nostrand Reinhold, New York, pp 163–167 Tanner WF, Stapor FW (1972) Accelerating crisis in beach erosion. Int Geogr 2:1020–1021

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1.10 Georgia

Miles Hayes

1.  Introduction The coastline of Georgia borders a wide plain on the trailing edge of the North American plate, which is generally tectonically stable to slightly downwarped. The coastline is low-lying and depositional, with Pleistocene and Holo­ cene barrier island, bay, lagoonal and estuarine deposits blanketing an irregular topography eroded during low stands of sea level in the Pleistocene epoch (Hayes 1994). The state borders the western flank of the Georgia Embayment, a large indentation in the coastline of southeastern United States. The outer coastline is about 160 km long, but the intricate tidal shorelines behind the barrier islands are several times that length. Wave heights are generally less than 0.8 m, and mean spring tides are between 1.2 and 2.8 m. The Georgia coast is generally tide-dominated, in contrast with the more wave-dominated coast in Florida to the south, and the strongest currents are generated through tidal inlets. The climate is humid subtropical, with abundant rainfall throughout the year. Hurricanes frequent the area in  late summer and early fall, and extra-tropical cyclonic depressions are common during the winter months. Except during the passage of these cyclonic storms, the winds blow prevailingly offshore, making this a lee coast.

2.  The Coastline The coast of Georgia is dominated by a chain of sandy barrier islands interrupted by sounds and backed by marshy estuarine embayments (>Fig. 1.10.1). Several rivers cross the coastal plain and widen into estuaries that converge behind the barrier islands. The rivers with watersheds restricted to the coastal plain, typically have large estuarine systems at their mouths (e.g., Satillo River), whereas the two rivers that drain the piedmont and mountain regions (Savannah and Altamaha) have created modestsized river deltas. The barrier islands show numerous beach and dune ridges running generally parallel to the Atlantic coastline,

but locally in transverse patterns that have been truncated. The ridges curve back beside inlets past and present. Four of the larger barrier island complexes have younger (Holocene) barrier islands on their seaward sides that are backed by remnants of earlier (Pleistocene) barriers and extensive wetlands. The stratigraphy of the barrier islands shows evidence of phases of progradation and retrogradation: inlets have formed, migrated and been filled (Hubbard et al. 1979). A rising Holocene sea level invaded valleys and lowlands to form the estuarine area, while barrier islands formed from sand either washed in from the continental shelf (having been previously deposited there by rivers) or brought to the shore by the two piedmont rivers. Some of the outer barrier islands have shown intermittent landward migration, notably during hurricane overwash. Transgressive dunes are spilling landward on Cum­ berland Island in the south of the state. Blackbeard Island has a complex pattern of intersecting beach ridges, representing several stages of accretion and erosion, with overall progradation. A marshy corridor, threaded by Blackbeard Creek, links it to Sapelo Island, which has a Pleistocene core, presently without beach ridges. The Holocene outer portion of Ossabow Island, to the north, has grown southward as several recurves. Tybee Island contains marsh with tidal creeks and an outerfringe of beach ridges. Attempts have been made to establish sequences of erosion and accretion in beach ridge systems. While there may be long-term recession of the coastline in response to the rising sea level in some areas, the survival of beach ridges indicates that progradation has occurred within the past 5,000 years elsewhere. With the largest tides in the SE United States, and the flattest coastal plain, the coast of Georgia is essentially inundated twice daily by the tides. The large estuarine embayments are flanked by extensive salt marshes. Howard and Frey (1980) classified the estuaries into two main categories: salt marsh and riverine. Salt marsh estuaries, the more abundant type, are dominated by marine waters. They contain sandy point bars, which are distinguished from fluvial point bars by their more abundant biogenic

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.10, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.10.1 The coast of Georgia (based on a mapby Howard and Frey 1980).

structures and mud content. Riverine estuaries are fed by piedmont or coastal plain rivers. Sediment carried by the piedmont rivers is more immature mineralogically than those of coastal plain rivers, which have passed through more than one cycle of erosion and deposition.

Relationships between organisms and sedimentology are well illustrated by salt marsh zonations, as on Sapelo Island (Edwards and Frey 1977) (>Fig. 1.10.2). Tidal inlets have large ebb-tidal deltas of the type illustrated in (>Fig. 1.10.3).

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⊡⊡ Fig. 1.10.2 Salt marsh zonation on Sapelo Island (after Edwards and Frey 1977).

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References Edwards JM, Frey RW (1977) Substrate characteristics within a Holocene marsh, Sapelo Island, Georgia. Senckenberg. Marit 9:215–259 Hayes MO (1994) Georgia Bight. In: Davis RA Jr (ed) Geol­­o­­gy of the  Holocene barrier island system. Springer-Verlag, Berlin, Chap 7, pp 233–304 Howard JD, Frey RW (1980) Holocene depositional environments of the Georgia coast and continental shelf. Guidebook 20: Excursions

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in Southeastern geology; The Archaeology-Geology of the Georgia Coast. Annual Meeting of the Geological Society of America 66–134 Hubbard DK, Oertel G, Nummedal D (1979) The role of wave and tidal currents in the development of tidal-inlet sedimentary structures and sand body geometry: examples from North Carolina, South Carolina, and Georgia. J Sed Petrol 49:1073–1092 Oertel GF (1973) Examination of textures and structures of mud in layered sediments at the entrance of a Georgia tidal inlet. J Sed Petrol 43:33–41

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1.11 South Carolina

Miles Hayes

1.  Introduction The coastline of South Carolina, like that of Georgia, flanks  a wide depositional coastal plain, which includes successive Quaternary formations, mainly sandy barrier islands and swamps, blanketing an irregular topography eroded during low stands of sea level in the Pleistocene epoch (Hayes 1985, 1994). It occupies the NW flank of the Georgia Embayment, a large indentation in the coastline (>Fig. 1.11.1). Coastal barrier islands are intersected and  backed by estuaries and bays, passing landward to freshwater swamps. The outer coastline is about 320 km long. Wave heights are generally between 0.8 m and about 1.2 m at the North Carolina border. Mean spring tides are between 1.5 and 2.8 m, and the morphology of the South Carolina coast is tide dominated, though with important wave effects; it was referred to as a mixed-energy coast by Hayes (1994). The climate is humid subtropical, with abundant rainfall throughout the year. Hurricanes occur, particularly in late summer and early fall, and extra tropical cyclonic depressions cross the coast with some regularity. The prevailing winds are westerly and so offshore.

2.  The Coastline The South Carolina coast is a depositional coastal-plain with barrier islands located between major estuarine systems. Two principle types of barrier islands are present on this coast, those that consistently migrate landward (transgressive) and those that build seaward (prograding). On some parts of the coast the landward-migrating barrier islands may move landward tens of m/year. However, those that build seaward usually grow more slowly, at rates of less than 3 or 4 m per decade. Which of these two types of barrier islands are present at a particular location along the coast depends upon the rate of sea level change relative to the sand supply (Moslow and Heron 1978; Hayes 1994). Diminished, or low, sand supply and/or rapid sea level rise promotes the development of transgressive ­barrier islands,

and the opposite for prograding islands. Which type of barrier island is present is a site-specific issue, with the availability of sand, which comes mostly from both offshore and alongshore sources, being the controlling factor in most cases. There are 36 km of transgressive barrier islands and 120 km of prograding barrier islands on the South Carolina coast. Transgressive barrier islands (>Fig. 1.11.2) are composed of coalescing wash over fans, or a wash over terrace, which is overtopped at high tides, usually several times a year. In the process of migration, the entire wash over terrace complex moves landward, leaving an eroded near shore zone in its wake (>Fig. 1.11.3). As a result of this type of migration, in three dimensions the entire complex consists of a relatively thin (<2–4 m) wedge of sand and shells of the wash over terrace that overlies muddy sediment originally deposited in the estuaries behind the islands. Prograding barrier islands (>Fig. 1.11.4) are typically composed of multiple, relatively parallel linear ridges of sand topped by fore dunes. As a result of this type of seaward growth, in three dimensions they consist of a wedge of sand 7–10 m thick that has built seaward over offshore mud. Because of the mesotidal setting, the barrier islands are relatively short (<15 km) and they are intersected by numerous tidal inlets. Examples of South Carolina tidal inlets, in which the ebb-tidal deltas dominate, are seen in >Figs.  1.11.5 and > 1.11.6. The ebb-tidal deltas, which contain 77% of the sand in the coastal zone of South Carolina, are giant sediment traps, which markedly influence erosional and depositional patterns on the adjacent barrier islands. They cause updrift progradation and downdrift erosion in some cases (Nummedal et al. 1977). The most notable example of this is where the main channel of a large ebb-tidal switched to the north abruptly in the mid-1850s, trapping the 150,000 m3 of sand/yr that typically moves southward along this coast in a natural groyne effect (Hayes and Kana 1976). At inlets where sand does readily by-pass the inlet, the coastline may build out on the downdrift side, as shown in >Fig. 1.11.6. Detailed studies of the tidal inlets in South Carolina by Hubbard et al. (1979) and FitzGerald (1976) showed that

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.11.1 Location map, South Carolina coast.

the character of the inlets changes from deep and seaward extended (tide dominated) in the south, as at Fripp Inlet (>Fig. 1.11.6), to complex and sand choked (wave dominated) toward the north. In summary, the dominant morphology of the central South Carolina coastline is a mesotidal barrier island complex intersected by numerous tidal inlets and backed by extensive salt marshes and tidal flats. The barrier islands that are prograding have multiple beach ridges, some ­surmounted by dunes, in patterns that indicate complex histories of erosion and accretion, and their persistence

showing overall accretion. Most of the larger, prograding barrier islands have drumstick-shaped configurations, as on Kiawah Island (>Fig. 1.11.7). Charleston stands beside a large estuary with a harbour entrance bordered by spits. To the northeast, Cape Romain is a cuspate foreland built by southward drifting sand, with prograding spits on both sides. The Santee/Pee Dee Delta is the largest river delta on the east coast of the United States, and Winyah Bay is now a large estuary with an entrance deflected southward by a longshore spit. This marks a change to a smoother, arcuate wave-dominated

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⊡⊡ Fig. 1.11.2 Morphology and subsurface three-dimensional configuration (stratigraphy) of a transgressive barrier island, with a cross-section (A-A’).

⊡⊡ Fig. 1.11.3 A transgressive barrier islands at Edingsville Beach, South Carolina. This coastline has retreated almost a mile since the 1850s. Photograph taken in 1974. (Courtesy M.F. Stephen.)

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⊡⊡ Fig. 1.11.4 Morphology and subsurface three-dimensional configuration (stratigraphy) of a prograding barrier island, with a cross-­sections (A-A’).

⊡⊡ Fig. 1.11.5 Price Inlet, showing ebb tide delta.

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⊡⊡ Fig. 1.11.6 Fripp Inlet, showing ebb tide delta. The coast to the south (left) has prograded further than that to the north. (Left; downdrift side of inlet).

⊡⊡ Fig. 1.11.7 An oblique infrared view from NE of Kiawah Island, South Carolina, which clearly illustrates the drumstick configuration of this 8.5-mile-long, prograding barrier island. Note the presence of linear ridges of sand vegetated by maritimeforest vegetation (arrow), which indicate the positions of the foredunes just back of the beach at earlier stages in the growth of the island. Photograph taken in 1976: the island is now highly developed with residences. (Courtesy D. K. Hubbard.)

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coastline that extends past Murriells Inlet and Myrtle Beach. The coastline then curves eastward into North Carolina towards Cape Fear. The present coastline shows both eroding and accreting sectors. Continuing changes are evident when recent Google Earth imagery is compared with air photographs published by Shepard and Wanless (1971). Erosion problems have arisen for coastal developers who have built structures too close to unstable and shifting tidal inlets. Captain Sam’s Inlet, located between Kiawah and Sea­ brook Islands, has been migrating southward at the rate of 60–70 m/year (Sexton and Hayes 1983).

References FitzGerald DM (1976) Ebb-tidal delta of Price Inlet, South Carolina. In: Hayes MO, Kana TW (eds) Terrigenous clastic depositional environments; some modern examples, Coastal Research Division, University of South Carolina, Technical Report 11-CRD:143–157

Hayes M  (1985) Atlantic USA – south. In: Bird ECF, Schwartz ML (eds) The World’s Coastline, Van Nostrand Reinhold, New york, pp 207–211 Hayes MO (1994) Georgia bight. In: Davis RA Jr (ed) Geology of the Holocene barrier island system, Springer-Verlag, Berlin, pp 233–304 Hayes MO, Kana TW (1976) Terrigenous clastic depositional environments, some modern examples. Coastal Research Division, University of South Carolina, Technical Report 11-CRD Hubbard DK, Oertel G, Nummedal D (1979) The role of wave and tidal currents in the development of tidal-inlet sedimentary structures and sand body geometry: examples from North Carolina, South Carolina, and Georgia. J Sediment Petrol 49:1073–1092 Moslow TR, Heron SD (1978) Relict inlets: preservation and occurrence in the Holocene stratigraphy of southern Core Banks, North Carolina. J Sediment Petrol 48:1275–1286 Nummedal D, Oertel GF, Hubbard DR, Hine AC (1977) Tidal inlet ­variability – Cape Hatteras to Cape Canaveral. In: Proceedings of American Society of Civil Engineers, Coastal Sediments 77, Charleston, SC, pp 543–562 Sexton WJ, Hayes MO (1983) Natural bar-bypassing of sand at a tidal inlet. In: Proceedings of the 18th International Conference on Coastal Engineering, Cape Town, South Africa Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw Hill, New York

1.12 North Carolina

Stanley R. Riggs

1.  Introduction The North Carolina coastline is a very large and complex system that consists of about 560 km of ocean coast, 23 inlets, over 8,000 km of estuarine shore, with over 8,000 sq km of brackish-water estuaries. North Carolina receives about 1,200 mm of rainfall throughout the year, resulting in a subtropical climate dominated by heavy vegetation. One drainage basin (Roanoke River) flows off the front side of the Appalachian Mountains, four others (Chowan, Tar, Neuse, and Cape Fear rivers) originate within the Piedmont Province, while six others (North, Pasquotank, Perquimins, Alligator, White Oak, and Waccamaw rivers) are small, black-water streams that originate within the Coastal Plain. The coastal system forms the interface between the Atlantic Ocean and a wide coastal plain composed of Cre­ taceous through Holocene sediments. The coastal plain has been formed on the trailing edge of the North American plate, a generally stable tectonic setting with only minor tectonic uplift associated with the mid-­Carolina Platform High (the former Cape Fear Arch) in the southernmost portion of the State. Also, the Albemarle Embayment in the northern portion of North Carolina experienced slight subsidence through the Quaternary. The general geological map of the North Carolina Coastal Plain shown in >Fig. 1.12.1 suggests major differences between the northern and southern coastal regions that reflect their geological heritage – the underlying geological framework. A line drawn from Raleigh through Cape Lookout separates the coastal system into two coastal provinces, each having a unique geological framework that results in distinctive types of continental shelves, barrier islands, inlets, and estuaries. The coastal system in the Southern Province, from Cape Lookout to the South Carolina border, is primarily underlain by rocks of Upper Cretaceous through Pliocene ages. In this region, there is only a thin and highly variable skin of surficial sand and mud of Quaternary age. The older units are generally composed of hard rocks including mudstone, sandstone, and limestone associated with the Carolina Platform, a major geological structure that

underlies the region between Myrtle Beach in South Carolina and Cape Fear in North Carolina. This structural platform rises close to the earth’s surface causing the older and harder rocks to be eroded and truncated by the coastline. This erosional topography results in relatively steep slopes with common exposures of the older rock units that control the coastal geometry. In contrast, the coastal system in the Northern Province, from Cape Lookout north to the Virginia border, is underlain primarily by younger sediment of Quater­ nary age. These generally consist of slightly unconsolidated mud, muddy sand, sand, and peat that thicken northward to fill the subsiding Albemarle Embayment with up to 70 m of sediment. The generally soft sediment were deposited during sea level fluctuations resulting from multiple glaciations and deglaciations of the Quaternary ice ages (Riggs et al. 1992). Consequently, the depositional topography is dominated by young sediment with low slopes and the older rock units are deeply buried beneath these surficial sediments. The different geological frameworks of the Southern and Northern Provinces produce dissimilar sediment supplies and land slopes. The Southern Province has an average subaerial slope landward of the coast of 0.6 m/km. In contrast, the Northern Province has an average subaerial slope of about 0.03 m/km landward of the coast. Thus, the rising sea level floods the disparate slopes producing different kinds of barrier island-inlet systems and associated estuaries. The steeper slopes of the Southern Province produce short, stubby barrier islands with eighteen inlets and narrow back-barrier estuaries. The gentle slopes of the Northern Province produce long barrier islands with only four inlets, and an extensive sequence of drowned-river estuaries that constitute the vast Albemarle-Pamlico estuarine system. These northern barrier islands project seaward, forming Cape Hatteras and the famous Outer Banks: a sand dam that semi-­ isolates the Albemarle-Pamlico estuarine system from the ocean. In addition, the continental shelf morphology within the two coastal provinces is dictated by the underlying geological frameworks. In the Southern Province, Onslow

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.12, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.12.1 Generalized geological map of the North Carolina Coastal Plain showing the two coastal provinces and the four geomorphological compartments of the North Carolina coastal system. The cuspate embayments are defined by the Carolina capes and associated cross-shelf, sand shoals. Due to the different geological framework and spatial geometry, the coastal system within each province and compartment is significantly different from the others. Geological map is modified from the NC Geological Survey State Map (1985).

and Long Bays have broad, shallow shelves dominated by hard bottoms due to the extensive rock-floored character of eroded Cretaceous to Pliocene strata cropping out in a broad sweeping arc around the Carolina Platform. In contrast, the shelf morphology within the two coastal compartments of the Northern Province is narrower, slightly steeper, and dominated by soft Quaternary sediment units that fill the Albemarle Embayment (Riggs et al. 1992). The resulting difference in shelf geometry is a prime factor in determining the type and magnitude of physical energy (waves, currents, and tides) expended on the shoreline, as  well as the geological and biological coastal system response. The North Carolina coastal system consists of four geomorphic compartments, each with its own physical

and chemical dynamics and resulting biological and ­geological components. These compartments, or cuspate embayments, are formed by the capes and associated shoals. Each cape contains a shore-perpendicular, shallow sand shoal that extends seaward across the continental shelf: Diamond Shoals off Cape Hatteras, Lookout Shoals off Cape Lookout and Frying Pan Shoals off Cape Fear. The geographic orientation of each compartment and the variable continental shelf geometry determine wave and current dynamics, astronomical and storm tide characteristics, and the type and response to specific storm systems including the extra-tropical storms (nor’easters) and tropical storms (hurricanes). Consequently the coastal systems within the Northern and Southern Provinces of North Carolina are different.

North Carolina

In the Northern Province, coastal systems consist of very high energy, wave-dominated barrier islands with few inlets  into deeply embayed, shallow, irregularly flooded; wind-tide dominated micro- to nano-tidal estuaries. Whereas, coastal systems of the Southern Province consist of moderate energy, wave-dominated barrier islands with abundant inlets to narrow, regularly flooded, micro- to mesotidal dominated estuaries. Consideration of these morphological differences has significant consequences upon both the deposition and preservation of the Holocene record. In the Southern Province, the shallow depth of the shoreface to inner-shelf ravinement surface has truncated and removed only the tops of the estuarine channel infills leaving a vast network of channels exposed on the continental shelf (Riggs et al. 1995). Within the narrow, back-barrier estuarine systems, the preserved Holocene record generally represents the last 3,000 years and only locally are older (Pleistocene) sediments preserved. Generally poor preservation of the Pleistocene formations is attributed to re-excavation of previously deposited sediment during sea level fluctu­ations. The embayed estuaries of northeastern North Carolina are found within the slightly subsiding Albemarle Embay­ ment, a large Quaternary basin on the mid-Atlantic continental margin that extends southward from the Virginia line to Cape Lookout. This basin contains an enigmatic stratigraphic record displaying an embayment history dominated by multiple episodes of fluvial incisement during glacial events and depositional sequences that infill the incised valleys during subsequent interglacials (Riggs et al. 1992). The infill record reflects general transgressions and regressions with sediment sequences that theoretically grade from fluvial to estuarine, barrier island, and open marine sediment or some combination thereof. Barrier islands show much variation. Many barriers are sediment starved while some have adequate supplies to maintain a healthy beach system (>Fig. 1.12.3). Each barrier segment is dependent upon its inheritance from the gene pool that determines its characteristics and health. The gene pool is a complex interaction between the barriers location relative to potential sediment sources, paleotopography of the underlying geological framework, and physical dynamics of the coastal system (Riggs et al. 1995). This inheritance determines sand supplies, shoreface geometry, barrier morphology, inlet stability and geometry, and shoreline recession rates. Ultimately, sediment availability determines the evolutionary history of each portion of the coastal system. Some barriers have been sediment-rich at various stages in their Holocene evolutionary development if they

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occurred in specific geological settings with significant ­sediment supplies. Three major potential sources of sand exist along the North Carolina coast: (a) Extensive sand and gravel occurs in fluvial channels and deltaic structures deposited by Piedmont-draining trunk rivers on the continental shelf during sea level low stands; (b) Tremen­ dous volumes of sand are potentially available to coastal segments adjacent to cape-shoal structures (Diamond, Lookout, and Frying Pan Shoals); (c) Locally, sand-rich stratigraphic units crop out on the inner continental shelf and shoreface that are being eroded by the ravinement surface. Ultimately, sand availability dictates the character of the barrier island that forms and determines how it subsequently responds to both ocean dynamics on short-term scales and sea level changes on the long-term scale. This in turn dictates the degree of isolation, the physical and chemical characteristics, and ecosystem composition of the sedimentological regime within the back-barrier estuarine basins. Sediment-starved coastal segments develop simple overwash barriers that are eroding landward (>Fig. 1.12.2). Landward migration of a coastline produces a transgressive barrier island that results from either lack of a sediment supply or a rise in sea level. The absence of sand supplies is commonly associated with submarine headlands with shoreface and inner shelf stratigraphy dominated by indurated lithologies or with paleovalleys dominated by estuarine mud infill. Both situations result in negligible production of ‘new’ sand and high shoreline recession rates. Barrier segments with minimal sand supplies are characterised by low and narrow islands dominated by storm-tide overwash processes and frequent opening and closing of shallow and highly migratory inletoutlet systems. Much of the sand is incorporated into storm surge overwash fans and associated flood-tide delta deposits that are often dominated by extensive back-­barrier salt marshes. The resulting islands contain only local and ephemeral dunes with little or no maritime forest or extensive vegetative cover due to the dominance of annual saltwater overwash events. Simple overwash barrier islands are abundant along the North Carolina coast. Many are within the Cape Look­ out National Seashore (including North and South Core Banks) and Cape Hatteras National Seashores (including portions of Pea, Hatteras, and Ocracoke Islands). About 40 km of the latter segments are in a severe state of erosion causing greater difficulty in holding the coastal highway on collapsing barrier island segments. A few overwash islands are in other hands such as Masonboro Island State Park. However, some of these islands are in private ownership and represent areas of severe and ongoing conflicts

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⊡⊡ Fig. 1.12.2 Simple overwash barrier islands form in sand-limited settings. Minimal and ephemeral sand occurs on an eroding shoreface that contributes limited amounts of ‘new’ sediment to the coastal system. These islands are dominated by overwash fans that form during major storm-tide events and produce wide and shallow sand and marsh habitats extending well into the back-barrier estuaries. These shallow flats are quickly colonized by fringing marshes that continue to trap sediment as long as the overwash processes continue unhindered by either natural changes or human development such as building barrier dune ridges, roads, and extensive walls of buildings. If the latter happens, then the back-barrier estuarine shoreline may shift from one dominated by constructive processes to one dominated by loss of fringing marsh habitat and net coastline recession.

between continuing development and natural dynamics. These include Topsail and Figure Eight Islands, Ocean Ise,  Wrightsville and Holden Beaches, and the towns of Rodanthe and Kure Beach. Sediment-rich coastal segments develop complex barrier islands consisting of two or more stacked sand sequences (>Fig. 1.12.3). Multiple accumulations of sediment occurred during previous portions of the Holocene history where and when significant sand supplies were available (>Fig. 1.12.4). If large sediment supplies were available sand accumulated in one or more storage bins that included barrier beach dunes, backbarrier dune fields, or regressive beach ridges (>Fig. 1.12.5). Seaward progradation of a shoreline produces a regressive barrier island that results from either introduction of large sand supplies or a  drop in sea level. If inlet systems develop, sand is stored in ebb- and flood-tide deltas and migrating channel deposits. If an inlet closes down, the ebb-tide delta collapses and supplies sand locally to adjacent barriers. Such

local influxes of ‘new’ sand can result in the development of small-scale backbarrier dune fields. Complex barrier islands are characterised by unique coastal features and are common on the barrier islands. Nags Head Woods, Jockeys Ridge, and Shackleford Banks are characterised by back-barrier dune fields with dunes up to 15 m high (>Fig. 1.12.4), whereas Kitty Hawk Woods, Buxton Woods, and Bogue Banks are characterised by accretionary beach-ridge structures and associated maritime forests (>Fig. 1.12.5). In consort with development of barrier width and height comes the growth of maritime forests and prehistoric and historic human occupation. Complex barriers may also be characterised by large estuarine sand shoal complexes such as Colington Shoals (>Fig. 1.12.4) or the filled portions of back-barrier estuaries such as portions of Currituck, Roanoke, and Core Sounds. If sand supplies decrease through time, barrier segments will grade from progradational to stable to recessional.

North Carolina

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⊡⊡ Fig. 1.12.3 Complex barrier islands form in coastal areas with large and available sand sources. Fluctuations in sea level and ocean dynamics result in accreted beach ridges, welded sand prisms, and development of back-barrier dune fields. These islands are high and wide with extensive sand deposits that prevent storm tides from washing over the top of the island. Thus back-barrier estuarine shorelines have no direct connection with oceanic processes and result in characteristics and processes similar to mainland estuarine shorelines.

Associated beaches respond similarly and grade from those characterised by long-term to short-term accretion and to low annual recession rates. The barrier islands have been much modified by human activites. In northeastern North Carolina the natural processes were forever altered in the late 1930s with construction of barrier dune ridges from the Virginia line south to Ocracoke Inlet. Continued construction and maintenance of the barrier dune ridges for the past seven decades has changed the dynamics of all barrier islands, in particular simple overwash barriers (>Fig. 1.12.6). The barrier dune ridges acted as walls that prevented most overwash processes associated with average storm events. Mini­mizing overwash resulted in little to no sediment delivery to the inner shores of the island long-term maintenance of the bordering marshes. This led directly to increased rates of sound-side ­shoreline erosion (Riggs 2001; Riggs and Ames 2003) and

extensive growth of fresh-water vegetation that characterises the northern barrier islands today. Only large storm events or multiple and sequential smaller storms lead to severe dune ridge erosion that finally breaches the dune, causing development of small overwash fans. The semi-fixed dune ridge, in concert with a natural sand deficiency, and net ocean coastline recession, has caused the ocean beach profile to steepen resulting in ­relatively higher rates of coastline recession. In addition, the Outer Banks that are backed by the vast expanses of highly energetic Pamlico and Albemarle Sounds are experiencing increased rates of back-barrier shoreline recession, causing an overall narrowing of the simple overwash ­barrier segments. Like walls of a wooden fort, the artificial dune ridges significantly altered the dynamic barrier island ­processes. In addition, it provided a false sense of security that encouraged development on the overwash barrier segments and fueled the ever-increasing conflicts

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⊡⊡ Fig. 1.12.4 Schematic map and geological cross section of the Kitty Hawk to Nags Head portion of the NC Outer Banks showing preliminary geological interpretations of a complex barrier island system. Interpretations are based upon high resolution seismic and ground-penetrating radar surveys in consort with limited drilling and radiocarbon age dates. The oldest Holocene deposit is the vast submarine sand shoal called Colington Shoals. It prograded southward over the RoanokeAlbemarle channel complex and encompassed extensive Pleistocene paleotopographic highs. The first sand prism (Kitty Hawk Woods-Colington Island-Nags Head Woods) was deposited by rising sea level on top of Colington Shoals. GPR data suggest that at least one and possibly two younger sand prisms were subsequently welded on to the barrier island.

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North Carolina

⊡⊡ Fig. 1.12.5 Interpretation of beach ridges and beach ridge sets on the Kitty Hawk Woods complex barrier island (from Fisher 1967). The map also indicates the back-barrier dune field that is overriding the beach ridges, and the modern beach prism.

Long poin I B

CURRITUCK SOUND

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⊡⊡ Fig. 1.12.6 The 1998 geo-referenced, infrared aerial photograph of a simple overwash- and inlet-dominated barrier island in the area between Avon and Buxton is an example of ongoing “barrier-island narrowing” since the 1852, topographic survey. Since 1852, the ocean shoreline has receded up to 750 m for a net loss of 75% island width in 146 years (average annual erosion rate = −5 m/yr). Also, the coastal highway, built in 1955, has been moved westward four times and now is located adjacent to the back-barrier shoreline and protected from the ocean by a double row of constructed barrier dune-ridges.

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between human and natural dynamics on these geologically ephemeral islands.

Hawk, with a pier along which measurements are made of nearshore profile variations in relation to incident wave and current processes.

2.  The Coastline of North Carolina

References

From the South Carolina border to the Virginia line the coast is entirely low-lying and sandy, dominated by beaches, barriers and backing lagoons and swamps. The chief sites of scientific interest are the cuspate forelands at Cape Fear, Cape Lookout and Cape Hatteras. Cemented shell beds off these capes have modified incident wave patterns and so the shaping of the coastline. There are numerous and variable tidal inlets which have tended to migrate southward: historical maps and charts show how they have changed in position, form and dimensions, some having been sealed off while others have been opened, usually during hurricane surges. There is a major research facility at Duck, to the north of Kitty

Riggs SR (2001) Shoreline erosion in North Carolina Estuaries. North Carolina Sea Grant College Program, Raleigh, Publication UNCSG-O1–11, 69p Riggs SR, Ames DV (2003) Drowning the North Carolina Coast: SeaLevel Rise and Estuarine Dynamics. North Carolina Sea Grant College Program, Raleigh. Publication No. UNC-SG-03-04, 152p Riggs SR, York LL, Wehmiller JF, Snyder SW (1992) Depositional patterns resulting from high-frequency Quaternary sea-level fluctuations in northeastern North Carolina. In: Fletcher CH, Wehmiller JF (eds) Quaternary coasts of the United States: marine and lacustrine systems: SEPM (Society of Sedimentary Geology). Special Publication 48:141–153 Riggs SR, Cleary WJ, Snyder SW (1995) Influence of inherited geologic framework on barrier shoreface morphology and dynamics. Marine Geol 126:213–234

1.13 Atlantic Coast Central (USA)

(Virginia, Maryland, Delaware and New Jersey)

1.  Introduction The coastlines of the Mid-Atlantic Bight in the eastern United States are backed by a coastal plain and fronted by a  broad continental shelf, which is structurally a major ­continental margin geosyncline (the Baltimore Canyon Trough). This shelf is underlain by more than 10,000 m of sediment, the basal deposits dating back to the Triassic. The position of the present coastline in relation to the continental shelf varies from Virginia, where it stands near the outer edge of the continental shelf to northern New Jersey, where it stands close to the hard rock basement. The reason for this variation is probably glaciotectonic. The northern end of the New Jersey coast is very close to the southernmost limit of the North American continental glacial ice sheet in Quaternary times. Many believe that a continental bulge formed, with isostatic upwarping in the area peripheral to the ice sheet at its maximum extent, and that with the waning of the ice sheet and its disappearance from New England, this peripheral bulge subsided rapidly. As the ice sheet melted in Late Pleistocene to early Holocene times, sea level rose toward its present level, penetrating further in the northern area of subsidence, notably into the Hudson River valley, and transgressing less far in Maryland and Virginia. Coastal submergence produced the major estuaries of Chesapeake Bay and Delaware Bay, and minor embayments, fronted by chains of barrier islands. The coast between Virginia and New Jersey is on the track of major storms that pass from south to north: hurricanes, which occur roughly once every 16 years, leaving marker horizons in marsh stratigraphy (Donnelly et al. 2001) and more frequent NE (cyclonic) storms, which occur at random, sometimes as many as five or six a year, and produce wind-driven waves from the northeast. Coastline retreat is sometimes very rapid during ­hurricanes and northeasters, but SE waves arriving across the long fetch of

the Atlantic Ocean generate shoreward ­drifting of sediment into Chesapeake Bay, Delaware Bay and Sandy Hook Bay and seaward drifting when storm waves withdraw sediment from the shore and deposit it on the inner shelf. Tide ranges are small. Mean spring tide range at Cape Henry is 1.0 m, at Atlantic City 1.5 m and Sandy Hook, northern New Jersey, 1.8 m.

2.  The Coastline The coast from the Carolinas to New Jersey is dominated by northward longshore drifting, which has built projecting spits at Cape Henry in Virginia, Cape Henlopen in Delaware, and Sandy Hook in New Jersey. South of each spit was a zone of erosion, which provided sediment, some of which was carried northward by longshore drifting. In Virginia, Maryland and Delaware, the coast intersects uplands of Pleistocene sediment resting on 30–50 m over late Tertiary sediment, while in New Jersey they are composed of Quaternary sediment overlying Cretaceous and Tertiary formations. Supplied with northward-drifting sand, the spits have been shaped by southeasterly waves, refracted around their northern ends. Cape Henry is a lobate foreland with many recurved beach ridges, enclosing a lagoon and marsh area on the southern side of the entrance to Chesapeake Bay. Chesapeake Bay is a large branching ria, formed by submergence of the valleys of the lower Susquehanna River and its tributaries. These were river valleys at the time of the low sea level that accompanied the Wisconsin glaciation, and they became submerged during the late Pleistocene– early Holocene marine transgression. Chesa­peake Bay, the largest inlet on the coastal plain of the Atlantic seaboard, is an estuary 311 km long and up to 40 km wide, opening to the sea between Cape Henry and Cape Charles. On its west coast are the long estua­rine inlets at the mouths of the

Edited version of a chapter by John C. Kraft in The World’s Coastline (1985:213–219). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.13, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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James, Pamunkey, Rappahan­nock, Potomac and Patuxent Rivers, the intervening promontories being locally cliffed, while the east coast is low-lying and marshy, with inlets at the mouths of smaller river valleys and some sandy beaches (> Fig. 1.13.1). The cliffs show erosion by frost action as well as wave attack. Within Chesapeake Bay, there are estuarine meander channels shaped by ebb and flood tidal currents (Ahnert 1960). There is a salinity gradient from sea water (35‰) between Cape Henry and Cape Charles to less than 15‰ at Annapolis and fresh water in the rivers.

Variations occur in response to weather conditions, with freshening by heavy rain and river floods and increasing brackishness in dry seasons. Most of the sediment deposited in Chesapeake Bay have come from the Susquehanna and other inflowing rivers. Deposition of fluvial sediment has increased as the result of soil erosion in the catchment, and has led to the advance of salt marshes, followed by woodland, around river mouths (Froomer 1980). On the other hand, continuing submergence, caused partly by land subsidence, ⊡⊡ Fig. 1.13.1 Sandy beach at Hooper Island, north coast of Chesapeake Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 1.13.2 Waterlogging in subsiding marshland, Hooper Island, north coast of Chesapeake Bay. (Courtesy Geostudies.)

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey)   

has resulted in waterlogging within marshes (> Fig. 1.13.2) and erosion of low-lying islands, some of which have shrunk in size while others have disappeared completely (Hunter 1914). Sharp’s Island in Maryland had an area of 283 ha in the late seventeenth century, reduced by erosion to about 242 ha by 1850. Even in the first decade of the twentieth century, it had farms and a hotel. By 1950, there was just room for half a dozen men to stand on a grassy islet. Now only the Sharp’s Point Light shows where it was (Douglas et al. 2000). The low-lying east coast of Chesapeake Bay borders the Delmarva peninsula, so called because it lies partly in Delaware, partly in Maryland and partly in Virginia. On its Atlantic coast, beaches and barrier islands have formed where sand supplied by longshore and onshore drifting has been delivered to the shore. The beaches are backed by dunes, then a corridor of lagoons and salt marshes, bordering the coastal plain. A scarp lying inland along the southern end of the Delmarva Peninsula may be a relic of an earlier Holocene coastline, formed 3,000–4,000 years ago, but this is difficult to reconcile with a generally accepted sea level history of continuing Holocene marine transgression. Great Machipongo Inlet, Little Machipongo Inlet, Wachapreague Inlet, and Metomkin Inlet are gaps between the barrier islands that have varied in width, depth, and position in response to the contest between tidal currents, which tend to maintain them, and waves, which tend to drift sand onshore and alongshore to divert and sometimes seal them. Several have intertidal and subtidal sandy ⊡⊡ Fig. 1.13.3 The east coast of Cape Henlopen, showing the eroding dunes. A fence has been built in the hope of trapping wind-blown sand to restore the profile. (Courtesy Geostudies.)

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deltas formed by inflowing currents (flood deltas), and more rarely outflowing currents (ebb deltas). At Chin­ coteague, the barrier islands are set forward in front of Chincoteague Bay, a coastal lagoon separated from the sea by elongated Assateague Island, a broad barrier island, which ends in a recurved spit at its southern limit. This is an anomaly on a coast with predominant northward longshore drifting, but probably results from occasional southward drifting during northeasters. At Ocean City, the barrier has been built over, and the beach is backed by a boardwalk. North of Ocean City, the barrier beach abuts the coastal plain, enclosing valley-mouth lagoons, and Rehoboth Beach culminates in the large recurved spit at Cape Henlopen. These barrier islands have been transgressive, moving landward as the result of washovers and blowouts during the continuing slow sea level rise. This is indicated by the stratigraphy of the coastal formations, the barrier sands overlying marsh and lagoon deposits in a vertical sequence that matches the horizontal landward succession. The landward advance of the barriers is accompanied by a landward spread of marshes and lagoons on to the backing coastal plain. Before the arrival of Europeans, Cape Henlopen was similar to Cape Henry in Virginia, a cuspate foreland bordering a marshy lagoon area behind recurved spits that began to form more than 2,000 years ago. After an offshore breakwater was built, the pattern of longshore drifting changed, and a spit began to grow northward in response to SE wave action. Erosion on the east coast (> Fig. 1.13.3)

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is matched by accretion on the north coast. The spit extended by almost a kilometre northward between 1842 and 1977, even though its northward growth has been impeded by the effects of tidal ebb and flow currents. This spit is tending to curve to the northwest, and is likely to develop into a cuspate foreland in the next 300–500 years (Kraft et al. 1980). Delaware Bay is a smaller version of Chesapeake Bay, formed by marine submergence of the lower Delaware valley. Tributary valleys are small, and the bay is simpler in outline, funnel-shaped, 84 km long and widening to 20 km at the entrance between Cape Henlopen and Cape May. The estuarine shores are generally low-lying and marshy, and a channel 10 m deep and 360 m wide has been dredged in the Delaware estuary to give ships access to the port of Philadelphia. From Cape May, a sandy shore curves round to the northeast and becomes another series of barrier islands with intervening gaps, notably Cold Spring Inlet, Hereford Inlet, Townsend Inlet, Carsons Inlet and Great Egg Harbour Inlet, beyond which the sandy barrier widens at Atlantic City. To the north are barrier islands interrupted by Absecon Inlet, Brigantine Inlet, Little Egg Inlet, and Beach Haven Inlet. The barrier then becomes more continuous at Long Beach Island, separated by Barnegat Inlet from Island Beach. As in Virginia, Maryland and Delaware, these barrier islands are backed by an elongated corridor of lagoons with extensive salt marshes, as at Brigantine

(> Fig.  1.13.4), while the inlets have had a history of migration, shallowing, closure, reopening, and deepening in relation to interacting wave and tidal current processes. The sandy coastline shows alternations of erosion and accretion, with major changes in and around tidal inlets (Dolan et al. 1979). Island Beach converges with the mainland coast from Point Pleasant northward to Highlands, where the Sandy Hook spit projects NNW. This is a simple recurved spit, bearing dunes and ridges that indicate stages in its growth. The east coast is exposed to Atlantic waves (> Fig. 1.13.5) and shows northward longshore drifting, while the west coast is generally more sheltered, but exposed to westerly storms. There has been erosion on the eastern shore, countered by armouring (> Fig. 1.13.6) and beach replenishment (Nordstrom and Allen 1980), matched by accretion at the northern end. Human impacts have been extensive on this coast. There are cities such as Rehoboth Beach in Delaware, built on high sectors, and in New Jersey most of the barrier islands are occupied by urban and suburban ­developments, mainly for recreational purposes. Ocean City Maryland, built on a low-lying barrier, is a large city that includes an extensive chain of condominiums. Poor planning and building practices have led to development and building over dunes and on to the backshore in many coastal areas, and as a result northeaster storms and hurricanes have wreaked havoc on this coast. In 1962, a condominium at

⊡⊡ Fig. 1.13.4 Salt marsh at Brigantine. (Courtesy Geostudies.)

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey)   

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⊡⊡ Fig. 1.13.5 The beach at Sea Bright, on the east coast of Sandy Hook. (Courtesy Geostudies.)

⊡⊡ Fig. 1.13.6 Western shore of Sandy Hook spit, north of Fort Hancock. (Courtesy Geostudies.)

Ocean City rolled over on its side in a northeaster. Yet, building in this coastal zone is continuing rapidly. Planners are attempting to legislate setback lines to limit the cost of storm destruction, which in the past has come to many billions of dollars, but although the nature of the coastal sedimentary environments, the rising sea level and the

hazards of inevitable storm surges are well understood, there have been few links between planners, engineers, and geologists. This is mainly because the development has been conducted by the private sector, and only recently have people recognised that a serious problem has developed (Cooper et al. 2008).

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Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey)

References Ahnert F (1960) Estuarine meanders in the Chesapeake Bay area. Geogr Rev 50:390–401 Cooper MJP, Beevers MD, Oppenheimer M (2008) The potential impacts of sea level rise on the coastal region of New Jersey, USA. Climatic Change 90:475–492 Dolan R, Hayden B, Rea C, Heywood J (1979) Shoreline erosion rates along the middle Atlantic coast of the United States. Geology 7:602–606 Donnelly JP, Roll S, Wengren M (2001) Sedimentary evidence of intense hurricane strikes from New Jersey. Geology 29:615–618 Douglas BC, Kearney MS, Leatherman SP (2000) Sea level rise: history

and consequences. Academic Press, New York Froomer N (1980) Morphologic changes in some Chesapeake Bay tidal mar­ ­shes. Zeitschrift für Geomorphologie, Supplementband 34:242–254 Hunter JF (1914) Erosion and sedimentation in Chesapeake Bay. United States Geological Survey Professional Paper 40:7–15 Kraft JC, Allen EA, Belknap DE, John CJ, Maurmeyer EM (1980) Processes and morphologic evolution of an estuarine and coastal barrier system. In: Leatherman SP (ed) Barrier Islands, from the Gulf of St Lawrence to the Gulf of Mexico. Academic Press, New York, pp 149–183 Nordstrom KF, Allen JR (1980) Geomorphologically compatible solutions to beach erosion. Zeitschrift für Geomorphologie, Supplementband 34:142–154

1.14 New York and New England

Henry Bokuniewicz

Introduction The Atlantic Ocean coast of the United States north of 40°30ʹ N is crossed by two important physiographic boundaries. The first is the limit of the coastal plain, a thick sequence of unconsolidated or semi-consolidated Tertiary formations that form a seaward thickening wedge down the eastern seaboard. Long Island (New York State) is the northern limit of this feature, the boundary running south of Staten Island, through The Narrows, up the East River along the east side of Manhattan, down Long Island Sound and out under Block Island, Martha’s Vineyard and Nantucket Island. Some Pliocene sediment, mostly glacially reworked, underlie Cape Cod. Hardly any of this material outcrops, being covered with glacial sediment. Staten Island, to the north of this boundary, is underlain by Triassic and Palaeozoic formations. North of Long Island, New York State, the coastline is principally composed of consolidated bedrock primarily of Cambrian and Ordovician age in the southern reaches (Connecticut) and Devonian in the north (Maine). Notable exceptions are Staten Island, which contains Triassic formations dating back to the formation of grabens at the opening of the Atlantic Ocean, and New Haven Harbour in Connecticut, where the less resistant Triassic sandstones intersect the coast. These rock formations mark the eastern margins of a Triassic graben. The second boundary is the southernmost limit of the  last continental glaciation. The Laurentide Ice Sheet reached these latitudes during the Wisconsin Stage of the Pleistocene Epoch, achieving its maximum extent about 23,200 years bp (Balco et al. 2002). The entire coast has been glaciated, and a thick blanket of glacial drift covers Long Island and Cape Cod. Recession of the ice left a terminal moraine from about 18,000 years bp that crosses Staten Island. It is breached at The Narrows and continues along to divide on either side of Peconic Bay on Long Island. At Lake Success a younger recessional moraine continues along the north shore of Long Island to the ‘north fork’ (Orient Point), and from there to the ocean coast of western Rhode Island and the north shore of Cape

Cod. The older recessional moraine crosses to the south fork (Montauk Point) and its eastern fragments form parts of Block Island, of Nantucket and, of Martha’s Vineyard. Deposits on Martha’s Vineyard have been dated 25,000 and 22,000 years bp. A younger moraine along the shore of Buzzards Bay, Massachusetts was laid down 18,800 years ago (Balco et al. 2002). The south coast of Long Island is fashioned in outwash sands, while the more northern parts exhibit a variety of eskers, drumlins, moraine segments and ground moraines. During the last century sea level has been rising at an average rate of about 3 mm a year along the southern coasts of this region and 2 mm a year in the north, as documented from tide gauge measurements (Emery and Aubrey 1991); radiocarbon evidence suggests that the rate of sea level rise has increased over the past few hundred years (Donnelly 2006). The tides are semi-diurnal. The coastline is microtidal in the south, where ocean tides are less then 1 m along the southern Long Island coast. Tidal ranges are 30 cm or less in restricted embayments, although a resonant standing tide, with a range in excess of 2 m, forms in western Long Island Sound. Tide ranges increase northward to 3 m at Boston and 6.5 m in northeastern Maine. Except for restricted embayments, notably Long Island Sound, Narragansett Bay and Cape Cod Bay, the coastline is exposed to open ocean waves propagated over a wide continental shelf. Significant wave heights can exceed 1 m. The region is subject to hurricanes, tropical storms and nor’easters, which are a combination of high, atmospheric pressure over land paired with low pressure offshore. Nor’easters produce strong winds from the northeast, which drive up water levels along the coast by Ekman transport. So-called storm tides (storm surges) that reach 1 m recur annually, while events with recurrence intervals of a century approach an elevation of nearly 2 m. The south coast of Long Island is divided into three stretches (Taney 1961). The western (ocean) coastline is a series of barrier islands backed by large, shallow bays (lagoons). Tidal wetlands and marsh islands are prevalent, especially in the western bays. All the tidal inlets have been

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stabilised. Near the eastern end of the Island the beach is cut directly into a glacial outwash plain in front of the terminal moraine. This section contains several small coastal ponds, which appear to be drowned proglacial stream valleys that have been isolated from the sea by beaches built and maintained by the prevalent, westward longshore drifting. Beaches are typically 30–60 m wide and ­composed of fine to medium quartz sand (>Fig. 1.14.1). The extreme eastern point of the Long Island ocean coastline, Montauk Point, is a high coastal bluff of unconsolidated rock capped by the terminal mo­­raine (Ronkonkoma moraine). The north coast of Long Island on Long Island Sound can be divided into two stretches. The western part is highly embayed as a result of the drowning of an irregular, ground moraine draped over antecedent topography, while the eastern section is a gently curving, uninterrupted line of high, unconsolidated coastal bluffs. The coastal plain is absent from the Connecticut coast, which has been formed by the partial submergence of the Coastal Slope peneplain, cut into Palaeozoic and ­Pre-Cambrian bedrock. The Coastal Slope has relatively low relief, and intersects sea level at a slope of about 1:100.

It is indented, harbouring numerous, small, isolated salt marshes. The Connecticut coast is a mosaic of consolidated bedrock (>Fig.  1.14.2; e.g. between Groton and Ston­ington parts of the Thimble Islands), pocket beaches, tombolos (e.g. Long Beach in Stratford, Con­­necticut) or spits (e.g. Griswold Point), all composed of reworked, glacial sands (>Fig. 1.14.3). Sandy beaches are ­isolated and usually less than 3 km long. Some are cut into sandy outwash heads like the Lordship Outwash at the mouth of the Housatonic River. Other stretches of the Connecticut coastline is composed of moraine fragments (e.g. the Norwalk Islands), and some, coastal bluffs. The western coast of Rhode Island is a series of baymouth barriers between sectors of coastline cut into mo­­ rainic sediment (Fisher 1985). Westward longshore drifting has formed beaches across the mouths of drowned, proglacial outwash stream valleys. The eastern, Rhode Island shore is deeply embayed, as at Narragansett Bay, which lies in a structural basin formed in less resistant Carboniferous slates and schists. Bedrock islands are mantled with glacial drift and sandy sediment is sufficient to form only minor pocket beaches. The recessional moraine is absent across

⊡⊡ Fig. 1.14.1 Southampton, New York. The steep incline of the foreshore is due to the relatively coarse nature of the sand. The foreshore becomes less inclined to the west as the sand becomes finer. (Courtesy Bret Bennington.)

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⊡⊡ Fig. 1.14.2 Bedrock coastline at Bluff Point, Connecticut. (Courtesy Nancy McHone and Ralph Lewis.)

⊡⊡ Fig. 1.14.3 Bluffs and beach at Bluff Point, Connecticut. (Courtesy Nancy McHone and Ralph Lewis.)

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the mouth of Narragansett Bay, but reappears as the Elizabeth Islands, which are islands of low relief composed of glacial drift and sheltering Buzzards Bay. Boulders from the moraine form a natural protection for sectors of the coastline. The low-energy coast of Buzzards Bay is cut into till and glaciofluvial deposits. The moraine continues eastward on the north side of southern Cape Cod. The large outer islands - Block Island, Martha’s Vineyard and Nan­ tucket Island are situated between 10 and 40 km from the mainland. Each has a section of coast cut into fragments of the terminal moraine that forms the south fork of Long Island. On Block Island the erosion of the moraine and its associated ground moraine results in high coastal cliffs. Sandy beaches have been formed on the east and west shores by longshore drifting (Fisher 1985). The moraine stretches along the northern parts of Martha’s Vineyard and Nantucket Island, while their southern shores are cut into glacial outwash plains. Cliff recession of up to 100 m year was measured on Wasque Point by Kaye (1973). Spits and bay-mouth barrier beaches were formed from sand eroded from the coastal cliffs. Notable on the north shore of Nantucket Island are a series of six cusps formed on the inner side of the spit enclosing Nantucket Harbour. The formation of these cusps has been ­attributed to reversals in wave direction (Rosen 1975), although other mechanisms such as tidal eddies or seiching have been proposed. The ocean coast of Cape Cod can be divided into four sections. The southern coast is cut into a sandy, pitted outwash plain south of the moraine. Submergence of the

proglacial topography has produced a series of coastal ponds and baymouth barriers. Cape Cod then extends northward, bordering Cape Cod Bay. This stretch of coast marks the boundary between two lobes of the Laurentide ice sheet, and is divided into the other three sections. The middle stretch is one of unconsolidated coastal cliffs, in excess of 50 m in elevation, cut into the glacial sediment (>Fig. 1.14.4). Longshore drifting of sand in front of these bluffs has built large spits both to the north (Provincetown) and south (Nauset and Monomoy Islands). The largest coastal sand dunes in the region are parabolic dunes found at Provincetown along the northern spit, which approach 60 m in elevation. The southern spit appears to undergo a cycle of breaching and inlet migration southward with a period estimated to be about 150 years. A series of post-glacial relict shoreline features have been identified submerged in Cape Cod Bay. The present coastline of Cape Cod Bay has relatively small coastal cliffs along its eastern margin (>Fig. 1.14.5) shaped by storms from the west (Uchupi et al. 2005). The embayed coastline has been reworked into spits and barriers. Boston Harbour sits in a Pre-Cambrian basin, and partly submerged drumlins appear as a scattering of islands around the Harbour, reworked at their margins into spits and tombolos. North of Boston rocky headlands dominate the coastline to Cape Ann, and north of this headland Plum Island (8 km long) is one of the northernmost large barrier islands. The New Hampshire coastline is of low relief. ­Cres­cent  beaches are found between rocky headlands (Fisher 1985; >Fig. 1.14.6). Post-glacial submergence has

⊡⊡ Fig. 1.14.4 Southern limit of unconsolidated coastal bluffs on the Cape Cod coastline. Sand spits extend southward. (Courtesy Jim O’Connell.)

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⊡⊡ Fig. 1.14.5 Beach at East Sandwich on Cape Cod Bay, Massachusetts. (Courtesy Jim O’Connell.)

⊡⊡ Fig. 1.14.6 Shore of weathered, metamorphic rock at Pemaguid Point, Maine. (Courtesy Maine Geological Survey.)

produced estuaries with fairly restricted access to the sea on rocky, glaciated terrain. Great Bay, at the New Hampshire-Maine border, for example, is a large, shallow estuary set in crystalline metamorphic and igneous bedrock covered irregularly with till and glaciofluvial sediment. It is connected to the sea through a narrow channel (the Piscatagua River) running about 15 km to the coast.

About half of the tidal coastline of Maine is composed of coastal cliffs cut into unconsolidated glacial drift. On the remainder, gently sloping granite descends to irregular coastlines, as on Baker Island, often with a scarp where the blanket of glacial drift has been cut back to expose the  crystalline basement rock. Glacial striations are seen on the granite surface, in places extending

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down through the zone of breaking waves. Cliffs are rare on the granite. Where there are steep coastal slopes, they are often inherited from older landscapes rather than of marine origin. The imprint of glaciation is obvious on Mount Desert Island, which is a high ridge crossed by a N-S glacial trough.

The Maine coastline can be sub-divided into four sections (Kelly 1987) dominated by the bedrock ­structure (>Fig. 1.14.7) and geomorphology. The southern most stretch is one of arcurate bays, where most of the sandy beaches are found. This section is composed of northeaststriking metasedimentary rocks blanketed by glacial sands ⊡⊡ Fig. 1.14.7 Jenness Beach, Rye, New Hampshire. (Courtesy Robert Roseen, University of New Hampshire.)

⊡⊡ Fig. 1.14.8 Gravel beach between rocky headlands at Machiasport, Maine. (Courtesy Marine Geological Survey.)

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and mud. Some of Maine’s largest salt marshes have formed behind these beaches (Tanner et al. 2006). Further north the coastline is indented by long, narrow, fjord-like bays that trend perpendicular to the coast, a signature of tightly folded metasedimentary bedrock. The inlets are rias rather than true fiords. Mudflats and salt marshes occupy sheltered embayments. Stretches of both em-bayed coast and straight, rocky cliffs can be found in the ­northernmost section (Tanner et al. 2006). The bays here are broader than those found further south. Beaches tend to be sandy in the southern section of the Maine coast, but shingle pocket beaches can be found to the north (>Fig. 1.14.8).

References Balco G, Stone JOH, Porter SC, Caffee MW (2002) Cosmogenic-nuclide ages for New England coastal moraines, Martha’s Vineyard and Cape Cod. Massachusetts USA. Quat Sci Rev 21:2127–2135 Donnelly JP (2006) A revised Late Holocene sea-level record for northern Massachusetts USA. J Coast Res 22:1051–1061

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Emery KO, Aubrey DG (1991) Sea-level, land-level and tide gauges. Springer-Verlag, New York, 237p Fisher JJ (1985) Atlantic USA – North. In: Bird ECF, Schwartz ML (eds)  The World’s Coastline, Van Nostrand Reinhold, New York, pp 223–234 Kaye C (1973) Map showing changes on shoreline of Martha’s Vineyard during the past 200 years. US Geological Survey Field Studies Map, MF-534 Kelly JT (1987) An inventory of environments and classification of Maine’s estuarine coastline. In: Rosen P, FitzGerald D (eds) A treatise on glaciated coastlines, Academic Press, San Diego, CA, pp 151–176 Rosen PS (1975) Origin and processes of cuspate spit shorelines. In: Cronin LE (ed) Estuarine Research, Vol 2, Academic Press, New York, pp 77–92 Taney NE (1961) Geomorphology of the south shore of Long Island NY. Department of the Army, Beach Erosion Board, Corps of Engineers, Technical Memorandum 128:51 Tanner BR, Perfect E, Kelley JT (2006) Fractal Analysis of Maine’s glaciated shoreline tests established coastal classification scheme. J Coast Res 22:1300–1304 Uchupi E, Giese G, Driscoll N, Aubrey DG (2005) Postglacial geomorphic evolution of a segment of Cape Cod Bay and adjacent Cape Cod, Massachusetts, USA. J Coast Res 21:1085–1106

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1.15 Great Lakes (USA)

Mary-Louise Byrne

1.  Introduction The Great Lakes consist of five interconnected freshwater lakes: Superior, Michigan, Huron, Erie and Ontario (Carter and Haras 1985). Their coastlines, including islands and connecting channels, total about 17,000 km. The lakes occupy tectonic basins modified by glacial erosion and are underlain, and in places bordered by solid rock. Pre-Cambrian igneous and metamorphic rocks of the Canadian Shield underlie Lake Superior and the northern shore of Georgian Bay, whereas Palaeozoic dolomites, limestones, shales, and sandstones underlie the other lakes (Hough 1958). Pleistocene drift deposits, mostly till and lake-related glacial deposits, mantle the rocks over much of the basin. Lake shore rock outcrops and drift deposits are commonly fronted by narrow (less than 15 m wide) sand or shingle beaches. Although, the overall physical characteristics of the basin were probably formed before the Pleistocene glaciation – there are earlier shales and limestones beneath the basins of Lakes Michigan, Huron, Erie and Ontario – the Pleistocene epoch left the most visible legacy in the form of the unconsolidated drift deposits on the floors of the lakes and around their shores. The lakes had various forms during the melting and retreat of the Wisconsin ice sheet after 18,000 years bp, at first draining southward to the Mississippi, but as the ice completely melted they took up their present positions. Postglacial isostatic rebound also had a profound effect on the basin, notably with respect to former lake elevations. The coastlines of Lakes Huron and Ontario have been rising relative to those of Lake Erie and  Lake Michigan by about 0.3 m/century. Larson and Schaetzl (2001) summarized the geological evolution of the Great Lakes, stating that issues with the modern Great Lakes are concerned with the shoreline configurations and changes in water level. Because the Great Lakes are located in the interior of the continent, between the source regions of the polar and tropical air masses, the basin has complex, rapidly

changing weather patterns. However, the lakes exert a modifying effect on temperature and humidity by acting as heat sources or sinks. Mean annual precipitation ranges from 746 mm (559 mm rain and 187 mm as snow) in Thunder Bay in the northeast of Lake Superior to 893 mm (804 mm rain and 89 mm as snow) in Kingsville on the south of Lake Erie and 964 mm (791 mm rain and 183 mm as snow) in Kingston on the southeast of Lake Ontario. The prevailing winds come from the west, and are the main cause of surface waves and currents. They can also produce short-term fluctuations in lake level during and after intense storms; fluctuations as great as 5 m on Lake Erie. Because the lakes are small, astronomic tides are limited, usually less than 5 cm in range; but seasonal fluctuations, related largely to variations in precipitation and evaporation, cause a seasonal variation of about 35 cm. Prolonged periods (several years) of abnormal precipitation or evaporation result in long-term oscillations of up to 2 m. Erosion and accretion on the lake shores is related primarily to these fluctuations (U.S. Army Engineer Division 1971; Larsen 1973). The lakes were a well-regulated natural system because of their small connecting channels, but two of them, Superior and Ontario, are now artificially regulated. There has been extensive urban and industrial development along the shores of the lakes, particularly near Cleveland and Chicago. Lake Superior is the largest and the most elevated (183 m above sea level) of the Great Lakes. Its regulated outflow drops 7 m in its journey of about 60 km down the St. Mary’s River before spilling into Lake Huron. Lake Huron is connected to Lake Michigan by the wide Straits of Mackinac and the two lakes are a hydrological unit, both being 176 m above sea level. The St. Clair River flows southward to Lake St. Clair and the Detroit River continues into Lake Erie, the fall in water level being only 2.5 m over a distance of about 120 km. At the eastern end of Lake Erie the Niagara River flows northward to Lake Ontario. It is  just over 50 km long and has an average discharge

A revised version of a chapter by C.H. Carter and W.S. Haras (1985).

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⊡⊡ Fig. 1.15.1 The cliff at Niagara Falls. (Courtesy Geostudies.)

of 57,000 m3/s over the Niagara Escarpment (>Fig. 1.15.1), falling about 50 m to create the spectacular Niagara Falls. The remaining 50 m difference between the levels of the two lakes is spread over 10 km of a much ­narrower river channel below the falls, where a ­thundering current rushes through a deep gorge before it dies out in Lake Ontario, 74  m above sea level. St. Lawrence River transports the outflow of Lake Ontario some 800 km downstream to the Atlantic Ocean.

2.  Lake Superior The American coastline of Lake Superior is about 2,990 km long, including islands. The southern coast consists mostly of Pre-Cambrian rock, but the relief is lower (less than 60 m) and more irregular. In addition to the rock are clay banks, sand and gravel banks, and dunes and marshlands. Sand and cobble beaches alternate with rocky stretches without beaches. Sites of special interest along the southern shore include Pictured Rocks National Lakeshore and the Grand Sable sand dunes, Quaternary sands that rise to about 60 m above the lake. The Apostle Islands National Lakeshore lies near the western end of the lake. It is now thought that Lake Superior separated from Lake Michigan about 1,200 years ago (Johnston et al. 2007), on the basis of the evidence of regional tilting indicated by water level gauges (Mainville and Craymer 2005).

3.  Lake Michigan Lake Michigan, the only lake to lie entirely in the United States, has a coastline length of about 2,635 km. The NW coast, which includes Green Bay, consists of wetlands, limestone bluffs up to 60 m high, and coastal slopes made up of Pleistocene drift, fronted by narrow sand and gravel beaches derived from the eroding glacial drift. Further south, along the western coast, the overall relief of the till cliffs increases to 40 m high, again with narrow sand and gravel beaches. Towards the southern end of the lake, the coast is largely artificial from Chicago, to Gary, Indiana where beaches are angled between groynes but the beaches widen along the southeastern shore. Sites of scientific interest include the Indiana Dunes National Lakeshore and the Sleeping Bear Dunes National Lakeshore. North­ward along the east coast, sand beaches are backed by dunes up to 135 m high, separated by bluffs and slopes in glacial drift. A complex history of lake level changes is indicated by the various types of dunes on the Michigan coastal plain (Arbogast et al. 2002; Timmons et al. 2007). There have been episodes of intermittent beach progradation, marked by the evolution of parallel dune and beach ridges, during phases of falling lake level (Olson 1958), but these are cut back when the lake level rises (Larsen 1973): an illustration of the Bruun Rule (Schwartz 1967). Baedke and Thompson (2000) analysed the relationship between isostasy and lake level and found that there have been ­different rates of vertical movement

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caused by isostasy and/or tectonism around the lake even though there are similar lake-level variations. The NE coast, including the branching inlets of Grand Traverse Bay, consists in part of marshes, sand and clay banks in addition to bluffs of ­glacial drift.

into Lake St. Clair. This small lake, which acts as part of the connecting channels between Huron and Erie, has a coastline of about 272 km; Lake St. Clair is bordered by wetlands. Anthropogenic structures dominate the western side of the lake, beside the city of Detroit.

4.  Lake Huron

6.  Lake Erie

Lake Huron has a coastline length on the American side of about 4,111 km, including 932 km on islands, which are more numerous than those in the other Great Lakes. Saginaw Bay is inset between Au Sable Point and Point aux Barques on the west coast. Between Mackinac and Saginaw Bay, the west coast of the lake is generally low-lying, with marshes and low slopes of glacial drift or bedrock. South from Point aux Barques there are cliffs and bluffs in glacial drift up to 12  m high. Except for the marshes, most of the shore is fronted by narrow sand beaches that increase in width toward the southern end of the lake.

Lake Erie, the shallowest and southernmost of the Great Lakes, has an American coastline of about 808 km, including 69 km on islands. Along the south coast the relief is low, and the shore is made up of sandy beaches and barriers alternating with cliffs and bluffs in glacial drift, bedrock, and some marshy sectors. The coast has been modified by port and industrial development at Toledo and Cleveland. There are some low limestone islands north of Sandusky, near the western end of the lake. The long, fairly straight coast east of Cleveland has cliffs and bluffs of drift and shale up to 30 m high, and eastward longshore drifting has built a large recurved sand spit that runs out at Presque Ile to shelter the port of Erie. Erosion along its western shore has been countered by beach nourishment and the building of a chain of nearshore breakwaters (>Fig.  1.15.2). To the east, the coast is made up largely of shale bluffs commonly capped by glacial drift.

5.  Lake St. Clair At its southern end, Lake Huron flows into St. Clair River in a low-lying corridor that opens through a marshy delta

⊡⊡ Fig. 1.15.2 The west coast of Presque Ile spit, Lake Erie, showing nearshore breakwaters. The sand lobe in the foreground is in the lee of a breakwater. (Courtesy Geostudies.)

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7.  Lake Ontario The American coastline of Lake Ontario is about 608 km long, with 88 km on islands. It has an average depth of about 86 m, with a maximum depth of 244 m. Along the south coast of the lake, from the western end are cliffs and bluffs up to 18 m high composed of glacial drift and outwash deposits. Further east, the bluffs decrease in size, and sandy beaches and barriers appear with some marshes. At the eastern end of the south coast, there are glacial drift cliffs up to 21 m high, sand barriers, dunes and marshes. The beaches are generally narrow and composed of sand  and gravel. A curved sandy beach forms the east coast, and to the north are numerous islands and bays in the approach to the Thousand Islands Channel that leads into St. Lawrence River.

References Arbogast A, Hansen E, Van Oort M (2002) Reconstructing the geomorphic evolution of large coastal dunes along the southeastern shore of Lake Michigan. Geomorphology 46:241–255

Baedke S, Thompson T (2000) A 4,700 year lake level and isostasy for Lake Michigan. J Great Lakes Res 26:416–426 Carter CH, Haras WS (1985) Great Lakes. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 253–260 Hough JL (1958) Geology of the Great Lakes. University of Illinois Press, Urbana, IL Johnston JW,   Thompson TA, Wilcox DA, Baedke SJ (2007) Geomor­ phic and sedimentologic evidence for the separation of Lake Superior from Lake Michigan and Huron. J Paleolimnol 37:349–364. DOI 10.1007/s10933-006-9052-3 Larsen C (1973) Variation in bluff recession in relation to lake level fluctuations. Illinois Institute of Environmental Quality, Chicago, IL Larson G,   Schaetzl R (2001) Origin and evolution of the Great Lakes. J Great Lakes Res 27(4):518–546 Mainville A, Craymer M (2005) Present-day tilting of the Great Lakes region based on water level gauges. Geol Soc Am Bull 117: 1070–1080 Olson JS (1958) Lake Michigan dune development. J Geol 66:254–263, 345–351 and 473–483 Schwartz ML (1967) The Bruun theory of sea level rise as a cause of shore erosion. J Geol 75:76–92 Timmons EA, Fisher TG, Hansen EC, Eisaman E, Daly, T, Kashgarian M (2007) Elucidating aeolian dune history from lacustrine sand records in the Lake Michigan Coastal Zone. The Holocene 17:789–801 U.S. Army Engineer Division, North Central (1971) National Shoreline Study, Great Lakes Region Inventory Report. United States 93rd Congress, 1st Session, House Document 93–121, 5:1–221

1.16 Hawaii

Charles H. Fletcher · Eden J. Feirstein

1.  Introduction The state of Hawaii consists of eight main islands: Hawaii, Maui, Kahoolawe, Lanai, Molokai, Oahu, Kauai, Niihau, and 124 small volcanic and carbonate islets offshore of the main islands, and to the northwest (> Fig. 1.16.1). The climate is sub tropical oceanic. Honolulu has a mean monthly temperature of 21.7°C in January, rising to 25.6°C in July, and an average annual rainfall of 802 mm. The Hawaiian archipelago lies in the zone of northeast­­erly  trade winds, which bring high rainfall to the windward mountain slopes while the leeward coasts are in rain shadow. Hawaii’s preeminent example of this orographic rain is Kauai’s Mount Waialeale, which receives an average 1168 cm of rain a year. In contrast, the dry air descending down Kauai’s leeward

side creates local semi-arid conditions at Polihale Beach on the west side of the island which receives on average a mere 20 cm of rain a year. In addition to northeasterly waves generated by the trade winds there are ocean swells from all directions, but notably from the north in the winter and the south in the summer months. Tide ranges are small, Honolulu having a mean spring tide range of 0.6–0.8 m. Geologically the islands are a series of volcanoes. The main Hawaiian Islands are great shield volcanoes built by successive flows of pahoehoe and aa basalt lavas. Some have had prolonged phases of erosion, followed by renewed eruptions. Hawaiian beaches include creamy white calcareous sand, derived from the tests of micro-organisms, weathered coral, calcareous marine algae, lithic fragments

⊡⊡ Fig. 1.16.1 The Hawaii archipelago consists of eight main islands and numerous smaller volcanic and carbonate islets.

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(typically of Pleistocene skeletal limestone), mollusc and echinoderm fragments (Harney et al. 2000), and black and green sand derived from volcanic material. Ferromagnesian olivine and other basaltic minerals are relatively unstable in the tropical climate of Hawaii, and are reduced by weathering. Calcareous beaches are dominant on all the older Hawaiian Islands where significant coral reef communities have been able to develop. The vast majority of Hawaii’s reefs are of the fringing variety (> Fig. 1.16.2). When an island is young and still volcanically active, lava entering the ocean prevents reef accretion. But moving up the west side of the Big Island where the seafloor is not swept by high swell, the beginnings of fringing reefs attached to the land can be seen and beautiful coral gardens are found. North of the Big Island on the Kihei coast of Maui and extending out from the north Maui shoreline are broad fringing reefs, demonstrating that reef development has a firm foothold on stabilised volcanic coasts. Fringing reefs generally grow in size and become commonplace among the islands north of the Big Island. But among the northwest Hawaiian Islands fringing reefs give way to submerged pinnacles, drowned platforms, and atolls as the volcanic shield structure subsides beneath the waves and the reefs struggle to stay near the surface. Two organisms serve as principal architects of Hawaiian reefs: scleractinian (stony or hard) corals, and coralline and

calcareous algae. There are over fifty species of coral found in the Hawaiian Islands but only a few are common. These grow in a range of forms designed to generally maximize the collection of sunlight and food, and minimize their exposure to stresses made by large waves. Stout branching, delicate branching, platy, encrusting, doming, mounding, these and other terms describe the numerous growth forms assumed by corals as they make the most of their environment. The more abundant Hawaiian genera include rice corals (Montipora species), lobe and finger corals (Porites species), cauliflower or moosehorn corals (Pocillopora species), and false brain corals (Pavona species). Coralline algae and calcareous algae are members of a marine plant group on the reef that deposits calcium carbonate in its tissue. When the algae dies, it leaves a fossil skeleton behind that is hard, whitish, and essentially the same chemistry as the coral. A few species of calcareous algae, such as the Halimeda, are especially abundant in Hawaii and important reef components. Hard plant debris builds up as piles of sediment in reef environments and are important sources of beach sand, making up over half the grains on many Hawaiian beaches. The coralline algae look like coral and grow in a binding and encrusting form on the reef, competing for space with corals. Most coralline algae are red, but there are some exceptions. A visit to any intertidal rocky coast in Hawaii will reveal the ⊡⊡ Fig. 1.16.2 Fringing reefs dominate coastal geomorphology and sedimentary processes on shores not influenced by active volcanism (Larsens Beach, Kauai). (Courtesy Coastal Geology Group, University of Hawaii.)

Hawaii

encrusting coralline community coloring the rocks a brilliant hue in between the rise and fall of the waves. The coastal plains of most Hawaiian Islands hold major calcareous aeolian, littoral, and marine sand deposits formed during and following late Holocene high sea levels  and persistent recent aeolian deposition under seasonal  winds (Fletcher and Jones 1996; Grossman et al. 1998). Sand is also stored on the reef flat in shore-normal reef channels and shallow Pleistocene karst depressions (Fletcher et al. 2008). Longshore transport dominates ­sediment movement on the coast in distinct littoral cells. The central Pacific location of the Hawaiian Islands exposes them to wind and ocean swells from all directions. Sectors of coastline may have rain, wind and wave shadows, and are either protected from, or vulnerable to, wind or wave impact. The four dominant regimes responsible for large waves in Hawaii are: north Pacific swell, trade wind swell, south swell, and Kona storms. The regions of influence of these regimes, outlined by Moberly and Chamberlain (1964), are shown in (> Fig. 1.16.3); a wave rose depicting annual swell heights and directions has been added to their original graphic (Vitousek and Fletcher 2008). Inter-annual and decadal cycles including El Niño Southern Oscillation (ENSO) occurring about every three to four years, and Pacific Decadal Oscillation (PDO) ⊡⊡ Fig. 1.16.3 Hawaii dominant swell regimes after Moberly and Chamberlain (1964), and wave monitoring buoy locations.

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occurring around 20–30 years, influence the ­variability of the Hawaiian wave climate. These large-scale oceanic and atmospheric phenomena are thought to control the magnitude and frequency of extreme swell events. For example, times of strong ENSO may result in larger and more frequent swell. Understanding the magnitude and frequency of extreme wave events is important as they may control processes such as coral development, sediment supply, and beach morphology. In the winter, Hawaii receives large ocean swell from extra-tropical storms that track predominantly eastward from origins in the northwest Pacific. These storms produce waves that travel for thousands of kilometres until reaching the shores of Hawaii. North swell have annually recurring maximum deep-water significant wave heights of 7.7 m with peak periods of 14–18 s. However, the size and number of swell events each year is highly variable – varying by a factor of 2. The annual maximum wave height ranges from about 6.8 m (in 1994, 1997, 2001) to 12.3 m (1988). Occurring about 75% of the year, the trade winds arrive from the east and northeast with an average speed of 25 km/hr and direction 73°. In winter months, the north Pacific high generating these winds flattens and moves closer to the islands decreasing trade wind persistence. Although the number of windy days in summer months

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increases, the mean trade-wind speed in summer and winter months remains similar. The trades generate choppy seas with average wave heights of 2 m (1σ = 0.5 m) and peak periods of 9 s (1σ = 2.5 s) from the northeast. Although these represent nominal conditions, trade-wind swell can exceed 5 m in height and have periods of 15–20 s. Southern swell arriving in Hawaii is typically generated farther away than north Pacific swell. These are usually produced by storms south of the equator near Australia, New Zealand and as far as the Southern Ocean and propagate to Hawaii with little attenuation outside the storm-generated region. South swell occur in summer months (southern hemisphere winter) and reach Hawaii with an annual significant wave height of 2.5–3 m and peak periods of 14–22 s, which are slightly longer than north Pacific swell. Hawaii has occasionally received tsunami generated by circum-Pacific earthquakes, as during the Alaskan earthquake of 1946. Kona storms are low-pressure areas (cyclones) of subtropical origin that usually develop northwest of Hawaii in winter and move slowly eastward, accompanied by southerly winds from whose direction the storm derives its name, and by the clouds and rain that have made these storms synonymous with bad weather in Hawaii. Strong Kona storms generate wave heights of 3–4 m and periods of 8–11 s, along with wind and rain, and can cause extensive damage to south and west facing shores. While minor Kona storms occur practically every year in Hawaii, major Kona storms producing strong winds, large wave heights and resulting shoreline change tend to occur every 5–10 years during the 20–30 year negative PDO cycle. Consequently, positive (warm) PDO, and El Niño phases tend to suppress Kona storm activity. Local relative sea level in Hawaii is not only dependent on global eustatic trends (about 3 mm/yr), but is also affected by subsidence of oceanic lithosphere, which responds elastically to volcanic loading over the hot spot. It is estimated that half of the upward building of Hawaiian volcanoes is reduced by subsidence and that most of the volcanoes have subsided 2–4 km since emerging above sea level (Moore 1987). The main Hawaiian Islands span about 5 million years in age. They are at differing distances from the hot spot, and thus at different stages of subsidence, and so have different relative sea levels. These relative differences in sea levels are demonstrated by modern tide gauge rates and support the view that subsidence is active over the hot spot (Moore 1987). Submerged wave cut notches and benches, raised coral reef, subaerial marine terraces, and alluviated river valleys can be found at various elevations and locations around

Hawaii, indicating oscillations of land and sea level (Fletcher et al. 2008). A fossiliferous marine limestone (known as the Waimanalo Formation) 3–6 m above sea level on Oahu is a typical example of a past stand of the sea. Holocene sea level has been influenced both by eustatic postglacial meltwater as well as equatorial oceanic siphoning associated with the changing postglacial geoid (Mitrovica and Milne 2002). These led to a high sea level (about 2 m) about 3,000 years bp followed by a sea level fall. Tide gauges record a sea level rise since 1900 in Hawaii. There is widespread but variable coastal erosion in the Hawaiian Islands in response to human interference with sand availability and the inferred influence of eustatic rise. For instance, researchers at the University of Hawaii measured the historical rate of shoreline change on every beach on the island of Kauai. Their data reveal that 72 percent of the beaches on Kauai are eroding and the average rate of erosion is 0.3 m/yr. On 22 percent of the eroding beaches, the rate of erosion is accelerating. In pristine coastal areas calcareous sand stored on the low-lying coastal plain is released to the beach as sea level rises, allowing a wide sandy shore to be maintained even as the beach migrates landward. However, the threat to coastal property has led to extensive armoring of the coastline. Artificial hardening is a form of coastal land protection that occurs at the expense of the beach, preventing waves from accessing the sandy reservoirs impounded behind the seawalls and revetments. Thus, efforts to mitigate coastal erosion have created a serious sand deficiency on armored beaches leading to widespread beach loss, particularly on the most populated and developed islands. The need to address this issue is acknowledged by the state and local communities, and the hope is that a broadly scoped management plan will keep a balance between the natural coastal morphology and human resource demands. The islands will be considered in sequence from WNW to ESE (which is also their age sequence): Niihau, Kauai, Oahu, Molokai, Lanai, Kahoolawe, Maui and Hawaii (The Big Island).

2.  Niihau Niihau is a small elongated island (29 km × 10 km) that stretches from SW to NE with a 145 km coastline. It lies at the far NW end of the main Hawaiian islands and is the low lying subaerial remains of a shield volcano built about 4.89 million years ago (Clague and Dalrymple 1987). The horseshoe shaped Lehua Island off the north coast is the product of rejuvenated volcanism and has been sculpted by marine erosion. The steep cliff on the south

Hawaii

side of this relatively small island has been notched with caves at the base. The island tapers to low-lying points that border a wide mouthed bay that opens to the north. The coast of Niihau has a variable morphology. Por­ tions are composed of low volcanic cliffs that have been weathered and shaped by marine erosion, boulder beaches, embayed sandy beach systems backed by 30 m high dunes, erosional sandy coasts where massive slabs of beach rock have been excavated by large swell and stacked on the shore, and small littoral cells with sandy coves.

3.  Kauai Kaui is known for the variety of microclimates that exist throughout the island including temperate regions, dry sand dune complexes and lush river valleys. Volcanic rocks form cliffs and the backshores of beaches, and there are segments of fringing reef cut by paleostream channels, coarse calcareous sandy beaches are separated by rocky points and interspersed with small stretches of boulder coast. The north coast beaches receive large winter surf that may create strong currents. Waterfalls cascade down deep­ ­ly eroded valleys, feeding streams that flow across the beach into the ocean. On the northwest coast one finds the steeply dipping knife-edge ridges and deep erosional V-shaped valleys that have been truncated into spectacular cliffs. The windward and south coasts are characterised by fringing reefs and inter mittent rocky headlands separating distinct littoral cells usually containing calcareous sand beaches. Persistent trade winds and the waves they generate, summer swell from the south, and winter swell from the north all refract around various portions of this coast and influence sedimentary processes. For instance, the Mana Plain on the west shore is a broad accretionary strand plain that originated with falling late Holocene sea level, and converging longshore sediment transport where north swell (refracting from the northwest) and trade wind waves (arriving from the east) deliver sand from two directions at various times of the year.

4.  Oahu Oahu has 180 km of irregular coastline that was greatly influenced by massive landslides that removed about a third of the eastern side of the island and about half the western side. The shape of the island is related to two mountain ranges, the eroded remnants of separate shield volcanoes truncated by the landslides.

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The north shore has massive winter surf (waves 15 m high and more), long sandy beaches, rocky points, and patches of beach rock exposed particularly in winter. The south coast of Oahu has been modified by urban and industrial development (> Fig. 1.16.4). Pearl Harbour formed as the island subsided towards the end of the main shield ­volcano building phase, drowning the river valleys that drain central Oahu. It contains almost 50 km of coastline backed by extensive wetlands through which highly sedimented waters enter the harbour. Along the shoreline, outcrops of 125,000 year-old skeletal limestone reveal fossil corals (> Fig. 1.16.5) in growth position and other paleoreef features that accreted under the higher seas and warmer temperatures of the last interglacial period.

5.  Molokai The island of Molokai has a 142 km coastline, and was formed by at least three shield volcanoes (MacDonald et  al. 1986). Maunaloa (420 m), the West Molokai volcano, is a dry flat-topped shield volcano partly protected from east and NE trade winds by East Molokai (1,514 m), which is wetter and has been cut into deep spectacular valleys and high steep ridges that are lush and vibrantly green. Most of the north coast is cliffed, with prominent ridges and valleys that have been carved by stream ­erosion. There are sectors of vertical cliff cut in volcanic formations.

6.  Lanai Lanai is a single shield volcano that formed from summit eruptions and along three rift zones between 1.2 and 1.46 million years ago, a classic example of a Hawaiian shield volcano with a gently sloping profile. The small sub-circular island has a 76 km coastline and a dry ­climate with minimal stream activity. As on Molokai overgrazing of domestic and feral animals in the nineteenth century and widespread deforestation have drastically changed the stability of the soil. The vegetation has never fully recovered and there is con­siderable wind erosion on the island (Macdonald et al. 1986). Along the northern coast windy conditions have led to the development of a series of low sand dunes behind beaches that are fronted offshore by a narrow fringing reef. The beaches, known collectively as Shipwreck Beach are composed mostly of calcareous sand punctuated by ­outcrops of lithified beach rock. Narrow Polihua Beach has a fringing reef.

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⊡⊡ Fig. 1.16.4 The south coast of Oahu is characterised by intense urban development. World famous Waikiki Beach is shown here in a view taken from the summit of Diamond Head, an extinct volcano on the coast southeast of Waikiki.

⊡⊡ Fig. 1.16.5 The rock shores of O’ahu (photo near Ka’ena Pt.) preserve a record of high sea levels and warmer seas from the last interglacial period (Eemian Interglacial). The yellow line highlights a large coral head that is about 125,000 years old. (Courtesy M. Dyer.)

7.  Kahoolawe Kahoolawe is the smallest of the main Hawaiian Islands, with just 47 km of coastline around a single shield volcano

that formed 1.03 million years ago. It has had a ­complicated history of human exploitation that has affected its terrain and environment, leaving it dry and barren. In the nine­­ teenth and early twentieth centuries, tens of ­thousands of

Hawaii

sheep and goats were raised on the island. Consequent grazing removed most of the vegetation, leaving the soil exposed to high winds, which eventually removed much of it. The coast has lava headlands separated by sandy beaches supplied with sediment from fringing reefs, with pocket beaches of detrital volcanic sand that lie at the mouths of stream gulches that descend from the centre of the island.

8.  Maui The island of Maui is composed of two large volcanoes separated by a low-lying isthmus. The West Maui Volcano (formed 1.6 million years ago) at an elevation of 1,764 m lies to the west of the massive Haleakala ­volcano (formed 0.8 million years ago) with a sum­­mit that reaches 3,055 m. Maui has 193 km of coastline that fringes the two main shield volcanoes and the isthmus. Again the north coast has large ocean waves, and cliffed headlands of volcanic rock alternate with sandy beaches.

9.  Hawaii The island of Hawaii lies over or just north of the Hawaiian hot spot and is composed of five volcanoes and one active seamount: Kohala, Hualalai, Mauna Loa, Kilauea, Mauna Kea and Loihi (located offshore). Of these, only Mauna Loa, Kilauea and Loihi are considered active, while Haulalai is dormant with its most recent eruption ending around 1800–1801.

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The island has 428 km of coastline and is so large relative to the other Hawaiian islands that it is known locally and abroad as the Big Island. Well-developed black and green sand beaches signify the reworking by waves and currents of recent lavas. These beaches are relatively limited along the rough volcanic coastline, and white calcareous sands are restricted because of poor coral reef development due to recent volcanic activity. There is lush vegetation in the NE, where the annual rainfall is 1,500–4,000 mm. The west coast of the Big Island has narrow white sand beaches and lush coral growth at several locations. Locals and visitors alike find that the shoreline has numerous spectacular examples of Hawaiian coastal environments. Kilauea is a large active volcano on the SE flanks of Mauna Loa. On 12 May 2002 the Mothers Day Flow commenced on the south flank of Kilauea and travelled as molten lava through pre-existing lava tubes down to the coast, where it formed a broad lava lobe. When the lava flows into the sea the molten rock (which may exceed 1,100°C) creates steam plumes When pāhoehoe lava enters the ocean for extended periods of time, new land is created in the form of a fan-shaped platform known as a lava bench. Lava pouring into the ocean from either surface flows or lava tubes cools rapidly, usually shattering in the cold water into sand- to block-size fragments. These accumulate along the shore to form a loose foundation that can support overlying lava flows which build a bench above sea level. However, the bench is deceptive; while it looks like solid land, it will in fact eventually collapse into the sea (> Fig. 1.16.6) sweeping unwary visitors to their death. National Park Service personnel usually rope these

⊡⊡ Fig. 1.16.6 A lava bench at East Lae’apuki on the south shore of the Big Island (left) collapsed during the night of 13 August 2007 (right). (Courtesy T. Orr.)

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­ angerous areas off to protect the public. Nonetheless, d foolhardy hikers are seen on these dangerous features, and have been lost when they visited the wrong place at the wrong time.

References Clague DA, Dalrymple GB (1987) The Hawaiian-Emperor volcanic chain: Part 1. Geologic evolution. U.S. Geological Survey Pro­fessional Paper 1350, 5–54 Fletcher CH, Bochicchio C, Conger CL, Engels M, Feirstein EJ, Gross­ man EE, Grigg R, Harney JN, Rooney JJ, Sherman CE, Vitousek S, Rubin K, Murray-Wallace CV (2008) Geology of Hawaii reefs, Chap. 11. In: Coral reefs of the U.S.A. Springer, Berlin, p 435–488 Fletcher CH, Jones AT (1996) Sea-level high stand recorded in Holocene shoreline deposits on Oahu, Hawaii. J Sediment Res 66 (3):632–641

Grossman EE, Fletcher CH, Richmond BM (1998) The Holocene sealevel high stand in the equatorial Pacific: analysis of the insular paleo sea-level database. Coral Reefs 17:309–327 Harney JN, Grossman EE, Richmond BM, Fletcher CH (2000) Age and composition of carbonate shoreface sediments, Kailua Bay, Oahu, Hawaii. Coral Reefs 19:141–154 Macdonald GA, Abbott AT, Peterson FL (1986) Volcanoes in the Sea: the  Geology of Hawaii, 2nd edn. University of Hawaii Press, Honolulu, HI Mitrovica JX, Milne GA (2002) On the origin of late Holocene sea-level high stands within equatorial ocean basins. Quat Sci Rev 21 (20–22):2179–2190 Moberly R, Chamberlain T (1964) Hawaiian Beach Systems. Hawaii Institute of Geophysics Moore JG (1987) Subsidence of the Hawaiian ridge. In: Decker RW, Wright TL, Stauffer PH (eds) Volcanism in Hawaii, US Geological Survey Professional Paper 1350:85–100 Vitousek S, Fletcher CH (2008) Maximum annually recurring wave heights in Hawaii. Pac Sci 62:4:541–553

2.0  Canada – Editorial Introduction

The major geomorphological divisions of Canada are the  Mountain Ranges in British Columbia, the Prairie Provinces, the low-lying Canadian Shield, the Arctic islands and the Maritime Provinces in the east. The Rocky Mountains consist of strongly folded Cainozoic formations, the Prairie Provinces gently-dipping strata, the Canadian Shield Pre-Cambrian crystalline and sedimentary rocks, the Arctic islands a mixture of granitic gneiss mountains and sedimentary lowlands and the Maritime Provinces folded and faulted Palaeozoic and Mesozoic rocks, an extension of the northern Appalachians. Canada has a very long and diverse coastline (McCann 1980; Trenhaile 1990). Its diversity reflects not only the vast size of the country and its geological complexity, but also its range of climate and tides. It has 243,935 km of coastline fronting on the Pacific, Atlantic and Arctic Oceans, of which roughly 172,950 km or 70.9% lies within the Arctic. Pleistocene glaciation was extensive in Canada. The mountains of British Columbia were strongly glaciated, there are residual mountain glaciers, and the coast has numerous fiords. Glacial drift is widespread in the Prairie Provinces, around the Great Lakes and in the Maritime Provinces, and the Canadian Shield includes large areas of ice-scoured rock. The Arctic islands have residual glaciers, notably on Ellesmere, Devon and Baffin Island. Much of Canada has winter snow cover, permafrost is widespread in northern regions, periglacial processes are active in mountain and Arctic regions, and in sea and shore, ice occur at least in winter, along much of the northern and eastern coastline.

1.  Pacific Coast The effects of glaciation and geological structure are evident along the steep coasts of > British Columbia, and in the many fiords. Active cliffs are of limited extent, many steep coast sectors plunging into deep nearshore water. Beaches are also of limited extent, and dunes very rare. South of Vancouver is the large Fraser River delta, forming a lowland that contrasts with the generally steep coast.

2.  Atlantic Coast > New

Brunswick, Nova Scotia and Cape Breton Island consist mainly of gently folded Mesozoic sandstones and shales interspersed with granite intrusions. The New Brunswick coast runs into the Bay of Fundy, and tide ranges increase rapidly to more than 15 m. Wave energy diminishes, and the intertidal zone widens, and becomes muddy and marshy in the Minas Basin and Chignecto Bay  to the north. The southern (Atlantic) coast of Nova Scotia is more exposed. Shore ice is present in winter, but storms generate high wave energy and the tide range is  small (<2 m). There are cliffed sectors and beaches of  gravel and sand, partly derived from the erosion of ­glacial drift, notably drumlins. The northern (Gulf of St. Lawrence) coasts of New Brunswick and Nova Scotia are relatively sheltered, but have cliffed sectors as well as sand and gravel beaches. The > Gulf of St Lawrence is a large marine inlet that narrows into the estuary of the St. Lawrence River. Wave action is generally moderate, with occasional storms. Shore ice is present in winter, and the tide range in­­ creases from about 2 m around the Gulf to more than 5 m in the St. Lawrence estuary. Prince Edward Island and then east coast of New Brunswick are fringed by sandy barriers fronting lagoons and marshes, the sand having come from glacial drift eroded in cliffs, or reworked on the sea floor. West of the cliffy Gaspe peninsula, the shore is marshy and boulder-strewn, while the northern coast bordering Labrador, is rocky with numerous small valleymouth rias. > St Pierre et Miquelon lie south of the large island of > Newfoundland, which constricts the eastern end of the Gulf. To the west are the Magdalen Islands, ­consisting of sandy ridges and dunes derived from the sea floor.

3.  Arctic Coast > Northern

Canada is defined broadly to include the Labrador coast is similar, with shore ice in winter. Deep

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fiords trench the Pre-Cambrian shield rocks north from Battle Harbour. There are frequent storms, when large waves break on rocky shores, and the tide range is relatively small (< 2 m). Further north, sea ice is extensive for much of the year, shore ice persisting into early summer when storms produce strong wave action in Davis Strait and Baffin Bay. The mountainous east coasts of Ellesmere Island and Baffin Island are intricately dissected by fiords, and the glaciers that persist on Ellesmere and Devon Islands, locally descend into the fiords. The megatidal coasts of Ungava Bay (tide range more than 6 m) and the macrotidal east coast of Hudson Bay (tide range about 4 m) have extensive tidal flats and marshes strewn with boulders deposited from glaciers. Emerged features, including beach ridges, result from postglacial isostatic uplift accompanying sea level variations. These are extensive around Hudson Bay, a large and relatively shallow (<100 m) marine depression in the Canadian Shield. Hudson Bay passes northward into the Foxe Basin, bordered to the west by the Melville Peninsula. Whereas Ellesmere Island and Baffin Island are mountainous, the northern coast with Prince of Wales Island, Victoria Island and Banks Island and the outlying Parry Archipelago are all generally low-lying. Treeless tundra landscapes

predominate. Permafrost is extensive and active periglaciation produces frost-shattered scree, cryoturbation features and thermokarst depressions. The coastline shows the effects of prior glaciation, with many fiards on coastal lowlands and fiords where the coast is steep. Glacial drift has been cut back in sea cliffs, and beaches of sand and gravel are supplied with sediment from cliff erosion, frost shattering on steep coastal slopes and fluvial deposits, notably from the Mackenzie River to the west. Sea ice is widespread in winter, closing many straits between the islands, and shore ice reduces the effectiveness of waves in the brief summer period. Wave energy is also reduced by short fetch across narrow straits along much of the north coast. The tide range is generally small (<1 m). West of Cape Bathurst, the Arctic lowlands include the large Mackenzie River delta and extensive low hilly country west to the Alaskan border.

References McCann SB (1980) The coastline of Canada. Canadian Geological Survey Paper 80–10 Trenhaile A (1990) The geomorphology of Canada. Oxford University Press, Toronto Ontario, Canada

2.1  British Columbia

John J. Clague

1.  Introduction The Pacific coast of Canada is a leading-edge continental margin backed by rugged mountains up to 4,000 m in elevation. It is characterised by a relatively narrow continental shelf and an irregular and mainly rocky coastal zone with countless inlets, fiords and islands (Clague and Bornhold 1980). The coastline has been extensively modified by Pleistocene glaciers and is affected by nearshore currents and waves. Energy levels are highest on exposed shores adjacent to the open North Pacific Ocean and lowest in sheltered inshore waters of the fiords and inlets.

2.  Coast of British Columbia The coastal areas of British Columbia can be divided broadly into three physiographic regions (Holland 1964). The first physiographic region is the Coast Mountains, into which all the fiords of mainland British Columbia extend. The Coastal Trough comprises of the Hecate Depression on the north and the Georgia Depression on the south. The Outer Mountain area includes the mountains of Haida Gwaii (Queen Charlotte Islands) and Vancouver Island. The Coast Mountains consist mainly of granitoid igneous rocks, but metamorphic, volcanic and sedimentary rocks of a range of ages are common. Vancouver Island and southwestern Haida Gwaii are composed mainly of folded and faulted volcanic and sedimentary rock of Mesozoic age, with some plutonic rocks. Northwestern Haida Gwaii is a low-lying, dissected surface underlain by lower Tertiary volcanic rocks. Weakly deformed sedimentary rocks ranging in age from Cretaceous to late Tertiary underlie northeastern Haida Gwaii, eastern Vancouver Island and Fraser Lowland, and also occur as a narrow fringe along parts of western Vancouver Island.

3.  Coast Mountains – “Fiord Land” The Coast Mountains lie east of the Coastal Trough and contain the fiords and inlets that dominate the British

Columbia mainland coast. Fiords range in width from 1 to 5 km, extend over 100 km inland, and are up to 700 m deep. Coastlines along fiords are dominantly steep; beaches are rare, short and gravelly. Small and medium-size deltas are common. Some deltas at fiord heads have extensive mudflats and tidal marshes upon which much of the development on this part of the coast has taken place. South from the Alaskan border the fiords increase in size and complexity, with many inner and transverse straits separating high islands. Mean spring tide ranges are large (5.3 m at Prince Rupert) and strong currents are genera­­ted in the straits. There are segments of low coastal pla­­teau similar to Norwegian strandflats but well vegetated (>Fig. 2.1.1).

4. Hecate Strait and Queen Charlotte Sound Hecate Strait lies between Haida Gwaii and the British Columbia mainland to the east. The bordering coasts consist of shallow open embayments, commonly with rocky intertidal zones, but locally with long stretches of gravelly beach. Small deltas are common along this relatively open coast, and in many areas sediment is swept from river mouths along the coast in the direction of the prevailing littoral currents. Argonaut Plain on the east coast of Haida Gwaii, east and southeast of Masset Sound, is underlain by thick sediment deposited by meltwater streams that flowed from a late Pleistocene ice front. With the exception of three small areas of bedrock, the entire 120 km length of coastline from Masset on the north to south of Tlell consists of sand and gravel beaches, locally backed by impressive sand dunes (Anderson and Walker 2006) (>Fig. 2.1.2). The mountain spine of Haida Gwaii is extremely rugged. Sea cliffs more than 200 m high are common along the west coast and consequently the littoral zone is narrow or absent in most areas. Numerous fiords indent the coast, and a few have short gravelly beaches and small deltas at their heads. To the south, Queen Charlotte Sound is open to the Pacific Ocean and backed by a coast of intricate fiords and islands.

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⊡⊡ Fig. 2.1.1 Forested strandflat near Prince Rupert on the northern mainland.

⊡⊡ Fig. 2.1.2 Broad sandy beach bordering the Argonaut Plain on Haida Gwaii.

5.  Vancouver Island The Strait of Georgia occupies a structural trough that has been shaped by Pleistocene glaciers. The glaciers eroded bedrock and deposited large amounts of till and stratified drift (>Fig.  2.1.3). The strait widens southward towards the U.S. border. The east coast of Vancouver Island has a narrow coastal plain, below 600 m in elevation, less than 30 km

in width, and extending from Johnstone Strait in the north to Victoria in the south (Nanaimo Lowland). The coastline north of Nanaimo consists largely of low-­ gradient broad sand and gravel beaches, derived mainly from erosion of Pleistocene sediment underlying the lowland. More bedrock and mixed sediment-rock intertidal zones occur south of Nanaimo; these are interspersed with local broad beaches. Differential erosion of sedimentary rocks has produced highly indented

British Columbia

2.1

⊡⊡ Fig. 2.1.3 Savary Island consists largely of outwash sands deposited in front of a glacier advancing down the Strait of Georgia about 30,000 years ago.

⊡⊡ Fig. 2.1.4 Goose Spit at Comox on the east coast of Vancouver Island. (Courtesy Geostudies.)

c­ oastlines on southern Vancouver Island and on adjacent islands at the south end of the Strait of Georgia. Only one fiord (Saanich Inlet) occurs on the east coast of Vancouver Island. Abundant sediment is supplied to the littoral zone along the east coast of Vancouver Island, because of erosion of Pleistocene sediment that underlie much of the adjacent lowland. In many areas, these sediment are actively transported along the coast by strong littoral currents and as a result, broad beaches, spits and bars are

common (>Fig. 2.1.4). Rising sea level is expected to increase erosion and sediment transport in these areas (Clague 1989; Shaw et al. 1998). The west coast of Vancouver Island has numerous fiords, many of which extend more than 50 km towards the interior of the island (>Fig.  2.1.5). Much of the shoreline is rugged and rocky, both on the exposed Pacific coast and along the fiords and inlets. Small deltas, marshes and gravelly beaches occur within the fiords. A narrow coastal ­lowland, the Estevan Coastal Plain, extends more than 270 km along

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⊡⊡ Fig. 2.1.5 Fiord on the west coast of Vancouver Island.

⊡⊡ Fig. 2.1.6 Long Beach, western Vancouver Island.

the west coast of Vancouver Island from the mouth of Juan de Fuca Strait to Brooks Peninsula in the north. The only extensive beaches on the west coast of Vancouver Island occur within the Estevan Coastal Plain (>Fig.  2.1.6). Most of the beaches have developed in areas underlain by thick Pleistocene sediment where offshore gradients are low. At Florencia and Wickaninnish bays,

between the towns of Ucluelet and Tofino, wide sandy beaches have developed through erosion of uncon­solidated sediment. Florencia Bay is a crescent-shaped embayment with a small cuspate foreland in the middle and rocky headlands at its northwest and southeast ends. The beach, about 5 km in length, is backed by a steep cliff of Pleistocene sediment.

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⊡⊡ Fig. 2.1.7 Cliff in Pleistocene sediment, Point Roberts south of Vancouver.

6.  Fraser Delta

References

Much coastal geoscience research in British Columbia has been directed to understanding the Fraser River delt south of Vancouver (Clague et al. 1998; Barrie and Currie 2000; Kostaschuk 2002; Mosher and Thomson 2002). The active, western front of the delta is about 23 km long and is fed by a network of distributary channels. Perhaps the most conspicuous features of the delta are sand and mudflats up to 6 km wide, locally crossed by distributary channels and bordered seaward by a foreslope with an average gradient of 1.5°, but with slopes up to 23° along its inner, shallower portions. Large amounts of Fraser River sediment are carried to the river mouth during the summer freshet. The western delta front is separated from the inactive southern part of the delta by Point Roberts Peninsula, a ridge cored by Pleistocene sediment (>Fig.  2.1.7). The upper foreslope of the active delta is primarily muddy north of the main distributary channel and sandy to the south, reflecting the dominant ­northward transport of Fraser River suspended sediment during floods.

Anderson JL, Walker IJ (2006) Airflow and sand transport variations within a backshore-parabolic dune plain complex; NE Graham Island, British Columbia, Canada. Geomorphology 77:17–34 Barrie JV, Currie RG (2000) Human impact on the sedimentary regime of the Fraser River delta, Canada. J Coastal Res 16:747–755 Clague JJ (1989) Sea levels on Canada’s Pacific coast: past and future trends. Episodes 12:29–33 Clague JJ, Bornhold BD (1980) Morphology and littoral processes of the Pacific coast of Canada. In: McCann SB (ed) The coastline of Canada. Geol surv Canada Paper 80–10. pp 339–380 Clague JJ, Luternauer JL, Mosher DC (1998) Geology and natural hazards of the Fraser River delta. Geol Surv Canada Bull 525:270 Holland SS (1964) Landforms of British Columbia, a physiographic outline. Br Columbia Dep Mines Petrol Resour Bull 48:138 Kostaschuk R (2002) Flow and sediment dynamics in migrating salinity intrusions, Fraser River estuary, Canada. Estuaries 25:197–203 Mosher DC, Thomson RE (2002) The Foreslope Hills; large-scale, finegrained sediment waves in the Strait of Georgia, British Columbia. Marine Geol 192:275–295 Shaw J, Taylor RB, Forbes DL, Ruz MH, Solomon S (1998) Sensitivity of the coasts of Canada to sea-level rise. Geol Surv Canada Bull 505:79

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2.2 New Brunswick and Nova Scotia

Eric Bird

1.  Introduction New Brunswick and Nova Scotia are geologically part of the northern section of the Appalachian system, and consist predominantly of deformed Palaeozoic sedimentary rocks trending SW–NE. The region was strongly glaciated in Pleistocene times, notably in the Wisconsin phase. Glaciated valleys and morainic deposits are extensive in lowland areas. The climate is cold in the winter months, when shore ice may form in sheltered bays (Owens and Bowen 1977). Atlantic ocean swell, attenuated across a broad con­ tinental shelf, reaches the coast of southeastern Nova Scotia, and is refracted into the Bay of Fundy and the New Brunswick coast. There is exposure to frequent storm waves generated during the passage of cyclonic storms from west to east across the northern Atlantic. Mean spring tide ranges are commonly less than 2 m on the Atlantic coast of Nova Scotia, but they are larger in the Bay of Fundy.

2.  New Brunswick From the Maine border, where Croix River flows into Oak Bay, the New Brunswick coastline is irregular, with bays and inlets between rocky headlands, as at Point Lepreau. Much of the coast is cut across crystalline rocks, forming a hilly hinterland, but there is a sector of Carboniferous rocks at Quaco. The coastline borders the Bay of Fundy, and exposure to Atlantic Ocean waves diminishes northeastward. At Saint John the mean spring tide range is 7.4 m, and it increases along the coast towards the upper branches of the Bay of Fundy, exceeding 10 m in Chignecto Bay. Coasts with tide ranges exceeding 6 m are classed as megatidal. Stacks cut in Carboniferous conglomerates in Blacks Harbour show unusually deep and tall abrasion notches due to the large tide range (>Fig.  2.2.1). Tidal bores run upstream in the Petitcodiac River, near Moncton at the head of Chignecto Bay. The tide is semidiurnal, with only a small diurnal inequality, and tide range has increased over the last 6,000 years as the dimensions of the bay have changed because of eustatic sea level rise and ­isostatic

recovery. Much of the interest in the tides and their influence on sediment dynamics is related to the possibilities of tidal power generation. The megatidal Cumberland Basin has wide muddy shores exposed at low tide and extensive salt marshes. There has been much land reclamation. A submerged forest at the mouth of the Missaguash River, near Fort Lawrence, is exposed at low spring tides and submerged by up to 10 m of sea water at high spring tides.

3.  Nova Scotia The coastline of Nova Scotia includes the megatidal Bay of Fundy, with maximum tide ranges rising to more than 16 m at the head of the bay, and the Atlantic coast which is generally low lying, mesotidal and dominated by occasional storm waves. Geologically Nova Scotia consists of two tectonic plate segments, the Avalon in the north and the Meguma in the south, separated by the major Cobequid-Chedabucto fault system which runs through Parrsboro and Five Islands ESE to Chedabucto Bay. South of this fault system, an Early Jurassic extrusion of basalt (the North Mountain Formation) forms headlands at Cape D’Or, Partridge Island, Wasson Bluff and Five Islands, and the long ridge of North Mountain and Digby Neck south of the Bay of Fundy. During the Quaternary there were several phases of glaciation, when ice moved from north to south across Nova Scotia. River valleys were deepened into troughs, and when the glaciers melted they left extensive deposits including moraines, drumlins (ice-moulded drift mounds) kames (gravelly, low hummocky mounds deposited on or within the ice), eskers from subglacial streams and glacifluvial outwash plains. The last (Wisconsin) glacial phase ended about 16,000 years ago when the climate became warmer and sea level rose in the Late Quaternary marine transgression. Spring tide range in the Bay of Fundy increases from 5.75 m at Brier Island to 13–14 m in Chignecto Bay and Amherst Harbour has a range of 11.6 m. At Burntcoat

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 2.2.1 Stacks on the New Brunswick shore of the Bay of Fundy, showing unusually tall abrasion notches on a megatidal coast. (Courtesy Gudrun Reichert.)

Head on the south coast of the Minas Basin spring tide range is 16.3 m. These are among the highest tide ranges in the world, possibly equalled or exceeded in Ungava Bay, Labrador. Sea level changes in Nova Scotia have been influenced by glacio-isostatic movements of the land. The Late Quaternary marine transgression was rapid between 8,000 and 2,000 years bp, and then decelerated. Tide gauge records show a continuing rise of about 3.2 mm/ year at Halifax, 4.7 mm/year at Yarmouth and 3.8 mm/ year at North Sydney on Cape Breton Island. Rising sea level in the mid-Holocene between 7,000 and 4,000 years bp amplified the tide range in the Bay of Fundy. Coastal processes are modified in winter by the formation of sea ice, particularly in the Bay of Fundy, where it begins to form in December, attains a maximum in February, and melts in March. Shore ice persists in estuaries into April. Sea ice, often containing boulders, gravel and sand, and darkened with finer sediment, is agitated by wave action and contributes to the erosion of cliffs and shore platforms. It is also rafted on to intertidal flats and salt marshes, contributing sediment as it melts. Alter­natively, sediment trapped in sea ice may be carried away in ice blocks floating offshore, and deposited on the sea floor. The boundary between New Brunswick and Nova Scotia follows the Missaguash River, which flows into the head of the Cumberland Basin. To the south is Lusby Marsh, dissected by meandering tidal creeks, on a coast where the spring tide range is more than 15 m. The wide estuary of the River Hebert comes in from the south, joined by the Maccan River estuary, which is deflected southward

by Amherst Point, a spit that widens southward, backed by salt marshes. Sedimentation has maintained these in a period of rising sea level (Shaw and Ceman 1999). Extensive salt marsh areas around the upper Bay of Fundy have been embanked (dyked) and reclaimed as pastureland. This megatidal coast has very wide shores exposed at low tide, with a tide-dominated morphology of sand bars and ripples of varying dimensions. There is a notable submerged forest near Fort Lawrence, in the Missaguash River estuary, exposed at low spring tides and submerged by up to 10 m of sea water at high spring tides. The megatidal environment results in tidal bores in the Missaguash River and the River Hebert at the head of the Cumberland Basin. Much of the interest in the tides and their influence on sediment dynamics is related to the possibilities for tidal power generation. Between the Cumberland Basin and these estuaries is the broad anticlinal Minudie promontory, marshy at its northern end and becoming hilly southward. At Downing Head the coast becomes cliffy, with exposures of Late Carboniferous (Cumberland Group) rocks that extend southward past Joggins and out to Ragged Point. The cliffs at Joggins, between Hardscrabble Point (Coal Mine Point) to the north and Ragged Reef Point to the south are known internationally as a classic coastal geology site -the Joggins Fossil Cliffs Provincial Park (> Figs. 2.2.2 and > 2.2.3). The vertical cliffs and wide intertidal shore platforms are cut in southward dipping grey Carboniferous sandstones and shales, with occasional coal seams and limestones. The strata exposed in the cliff face show the southward dip, with some faulting, on the flank of the

New Brunswick and Nova Scotia

Minudie Anticline, declining as it descends into the Atholl Syncline to the south. The basal Boss Point Formation outcrops in the cliffs and shore platform north of Little Cove and passes southward beneath the Joggins Formation, which in turn dips under the Springill Mines and Ragged Reef Formations to the south. The shore is accessible at Lower Cove, or down steps into a chine at the mouth of Bell Brook, where two

⊡⊡ Fig. 2.2.2 Gravel-strewn shore platform at Joggins, fronting a cliff in southward-dipping Carboniferous sandstone. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.3 Cliffs near Hardscramble Point. Joggins. (Courtesy Geostudies.)

2.2

streams descend dip planes and converge. At Bell Brook there is a shingle upper beach with grit and sand, some coal pebbles and bright red bricky sandstone fragments sorted into shore-parallel zones. The cliffed coastline cuts across the geological strike, and shows variations related to harder and softer outcrops. The intertidal shore platform shows similar variations, and 2.5° to 3°, spring tide range here being 13 m. The platform is strewn with

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gravel, including coal fragments, particularly north of the piles and timbers which are the remains of the dock at Joggins where coal was loaded. The lower shore is sandy. Ribs of hard sandstone run out across the shore and form minor cuestas, with steep escarpments facing north and gentler dip-slopes southward between lower and flatter zones on softer intervening outcrops. The shore platform has been cut by wave abrasion as the tide rises and falls by the action of driven sea ice action, and frost weathering. Near Hardscramble Point the Upper Carboniferous rocks are overlain by glacial drift. Coal has been mined here, and the coal-bearing sediment include upright casts of tree trunks perpendicular to the dipping strata. The rock formations are fossiliferous, yielding grey coral fossils, crinoid stems, shells and leaf impressions. In 1842 Charles Lyell visited here, and with local experts John William Dawson and Abraham Gesner debated the origin of coal. Dawson and Lyell found vertebrate bones in a lithified fossil tree stump. It has been suggested that the Joggins Fossil Cliffs should become a UNESCO World Heritage Site because of their geological significance. The Carboniferous rocks are capped in several sectors by unconsolidated Pleistocene glacial drift deposits. The cliff section near MacCarrons Brook shows three stages of glacial deposition – an upper yellow Shulie Lake till over middle grey Joggins Till resting upon lower reddish muddy McCarron Brook Till.

The generally cliffed coast continues SW to Cape Capstan, beside Apple River Bay, where The Bar is a midbay barrier spit fronting grassy salt marsh in the estuary of Apple River. The spit passes south into a beach with cliffs behind curving Edgells Beach, which has shore plat­forms cut in grey Carboniferous sandstone. Beyond Pudsey Point the coast rises and steepens southward to the bold headland at Cape Chignecto. At Cape Chignecto the coast turns ENE and becomes very steep, dissected by several deeply-incised river valleys. At McGahey Brook it curves ESE and begins to decline and at West Advocate the sloping cliffs (>Fig.  2.2.4) descend to a shingle beach piled with driftwood at high tide and interrupted by outcropping ledges of pink Triassic sandstone forming a sloping abrasion platform. The beach continues as a barrier spit fronting a former high tide salt marsh enclosed by a dyke and reclaimed as meadowland - then Advocate Harbour – which has a narrow outlet between paired spits. The hinterland is hilly and wooded on the Carboniferous formations. South of Advocate Harbour the land rises to Cape D’Or, a bold cliffy headland of Jurassic basalt, with high partly overhanging cliffs on the western side. This is the Cape d’Or Scenic Area. The headland takes its name from the cliffs stained yellow by copper (not gold) ores. A steep grassy bluff descends to a low promontory with a projecting basalt ridge and southern cliffs with ledges cut across columnar basalts dipping south. The large tides produce a pronounced eddy, the ebb current swinging round to run

⊡⊡ Fig. 2.2.4 Cliffs in Triassic sandstone at West Advocate. (Courtesy Geostudies.)

New Brunswick and Nova Scotia

westward inshore and forming waterfalls over basalt reefs. A similar eddy develops on the rising tide on the western side, and the two meet in the rough waters of the Dory Rips. On the eastern side of Cape d’Or is Horseshoe Cove, where a shingle beach impounds a stream-mouth pond. Cliffs of southward-tilted basalt run out eastward to Cape Spencer – also of basalt (>Fig. 2.2.5). The high forested basalt ridge of Cape d’Or is backed by a low corridor with a fault along which rock formations have been displaced laterally (much like the Great Glen in Scotland). In Greville Bay, to the east, low bluffs of glacial drift and grey Carboniferous shales are interrupted by the mouths of several incised valleys, some with salt marsh segments at high tide level – as at Wards Brook which drops through a gorge to a marshy tidal inlet. At Port Greville a coastal terrace is bordered by a bluff that descends to a shingle beach. Fox River and a series of small streams are incised into a hilly coastal fringe, and descend to marshy inlets, some with sand and shingle barrier beaches across their mouths, as at Fox Point and the mouth of Diligent River. The coast rises to steeply sloping cliffs at Bull Bluff, and at Black Rock a longshore spit of grey shingle has grown eastward to shelter a high tide lagoon. Cape Sharp is another headland of Jurassic basalt. Black Rock is also the name of a small rocky island offshore, round which strong tidal currents swirl. Barrier spits enclose tidal lagoons (>Fig.  2.2.6) and tombolos link nearshore islands. The Minas Channel narrows to about 5 km here, with views

⊡⊡ Fig. 2.2.5 Basalt cliff on promontory near Cape D’Or. (Courtesy Geostudies.)

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across to North Mountain, a high basalt ridge ending westward at Cape Split. The Late Quaternary marine transgression flooded the land to various levels, declining eastward from at least 35 m near Cape Chignecto to 0 m at Truro, as recorded by proglacial deltas. In the Parrsboro region there are terraces of meltwater outwash and postglacial fluvial sediment dissected as the sea dropped back from the Late Glacial marine limit. The major Cobequid-Chedabuto Fault system runs inland behind Parrsboro, parallel to the coastline and marked by a scarp in the Upper Carboniferous rocks on the hard Horton Group beside the softer downfaulted Parrsboro Formation. The strait between Cape Split and Cape Sharp, west of Parrsboro, leads into the Minas Basin, where mean spring tide range is over 15 m. At Burntcoat Head on the south coast of the Minas Basin spring tide range is 16.3 m. These are among the highest tide ranges in the world, possibly equalled or exceeded in Ungava Bay, Labrador. The macrotidal coast has very wide shores exposed at low tide, with a tide-dominated morphology of sand bars and ripples of varying dimensions. The bordering coast has extensive glacial drift deposits and soft sandstones, locally cut into cliffs which are reached by waves only briefly at high spring tide, and have a sloping profile. There are beaches of sand and shingle, derived from the eroding cliffs (>Fig.  2.2.7), ­passing ­seaward into gravel with a thin mud veneer. Amos and Long (1980) estimated a net annual sediment input into

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the Basin of 4.8 million m3, of which 3.09 million m3 is  derived from cliff recession. In addition to beaches, there are longshore spits, shaped by a combination of  high tide waves and currents. Erratic boulders (>Fig.  2.2.8) have been stranded on the shore, and in sheltered bays there are segments of salt marsh in the upper intertidal zone; these become more widespread in the upper bays and estuaries.

East of Cape Sharp is the curving shore of Parrsboro Roads, backed by high sloping bluffs. These curve out to East Bay, where a broad tombolo of sand and gravel links forested, flat-topped Partridge Island – also of dark basalt – to the mainland. East Bay has a curving beach of fine platy shingle, with overwash fans along the crest, bordering the tombolo, a grey gritty plain that includes a tidal creek and sparse marsh, with an outlet on the western

⊡⊡ Fig. 2.2.6 Barrier spit and tidal lagoon at Black Rock. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.7 Sloping cliff of glacial drift near Sand Point on the north coast of the Minas Basin. (Courtesy Geostudies.)

New Brunswick and Nova Scotia

2.2

⊡⊡ Fig. 2.2.8 Erratic boulders on the shore at Five Islands. (Courtesy Geostudies.)

shore close to the island. Across the Minas Channel is Cape Blomidon, with an upper vertical cliff of columnar Jurassic basalt above a steep slope on red Triassic rocks. Parrsboro has a spring tide range of 11.6 m. Farrells River opens into a branching estuary at Parrsboro, and to the east, low cliffs extend to Clarke Head, and on past Wasson Bluff and McKay Head to the Five Islands Provincial Park. The coast east of Wasson Bluff has extensive glacial drift deposits and soft sandstones, locally cut into cliffs which are reached by waves only briefly at high spring tide, and has a sloping profile. There are beaches of sand and shingle, derived from the eroding cliffs, passing seaward into gravel with a thin mud veneer. As well as beaches there are longshore spits, shaped by a combination of high tide waves and currents. Erratic boulders have been stranded on the shore, and in sheltered bays there are segments of salt marsh in the upper intertidal zone. At Five Islands the high islands form a dissected extension of the ridge running west from Economy Mountain. A cliff south of the lighthouse shows the red sand­stones and mudstones of the early Jurassic McCoy Brook sedimen­ tary Formation faulted against dark grey North Moun­tain basalt to the west. On the northern side of the promontory, the red McCoy Brook Formation is capped by glacial drift and outwash, while on the southern side the upfaulted North Mountain basalt caps sloping red cliffs in the Triassic Blomidon Formation. At the end of the mainland cliff is the Old Wife basalt sea stack. The backshore has sectors of Spartina marsh.

The area around the Five Islands drains out as the tide falls to reveal wide sand flats and large sand bars. The sand bars, which are generally composed of medium to coarse sand, are elongated and may be several kilometres in length, with an amplitude of 10–15 m. Their surfaces are covered with large ripples generated by tidal currents, which reach velocities in excess of 1.5 m/s. It has been suggested that the large sand bars are equilibrium forms subjected to a dispersal system of alternating flood and ebb transport. Tidal currents and sediment transport pathways maintain the pattern of sand bars and channels in Cobequid Bay at the head of the Minas Basin (Knight 1980). Present on the sand bars are four kinds of bedform: current ripples (height about 0.05 m, wave length about 0.3 m), two kinds of megaripples (height 0.05  m, wave length 1.0 m and height 0.05 m, wave length 15 m) and sand waves (height 0.2–3.0 m, wave length 15–200 m) (Dalrymple et al. 1978). Most sand waves have megaripples superimposed on them (>Fig. 2.2.9); and ripples are superimposed on all larger bedforms. As a result of these and similar studies, the Minas Basin is a key reference area for models of bar and ripple formation. The combination of coarse grain size and mobile sediment means that the intertidal fauna of the sand bars is sparse, but the muddier tidal flats support a wide variety of organisms. Resuspension of biodeposits contributes significantly to turbidity of the Minas Basin waters. East of Economy Mountain at Carrs Brook are cliffs and stacks in the Triassic Wolfville Formation. Economy

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⊡⊡ Fig. 2.2.9 Sloping cliff of glacial drift near Sand Point on the north coast of the Minas Basin. (Courtesy Geostudies.)

Point is another wooded peninsula with red earthy cliffs on its western coast. The coast to the east is generally low, with the mouths of several rivers draining down from the Cobequid Mountain. Portapique River is gravelly, and Bass River opens into a marshy bay. At Saints Rest an intertidal deposit includes tree stumps dated 8,180 and more than 4,400 years bp, indicating a lower sea level. At Highland Village low cliffs cut in pink clay (Triassic Wolfville Formation) stand behind wide sand and mudflats exposed at low tide, and the coast is low and marshy east of Spencers Point. Cobequid Bay has water area of 306 sq. km at high tide, and intertidal area of 186 sq. km exposed at low tide. It narrows eastward the combined estuary of North River and Salmon River at Truro, and has extensive intertidal sand bars (Dalrymple et al. 1975). Amos and Long (1980) estimated a net annual sediment input into the Minas Basin of 4.8 million m3, of which 3.09 million m3 was derived from cliff recession. The southern coast of Cobequid Bay is also low-lying, and the funnel-shaped estuary of Shubenacadie River comes in from the south. West from Salter Head the coast is embayed, with many small creeks opening through salt marshes. Cobequid Bay widens westward into Minas Basin, which receives the broad estuary of the Avon River. Cornwallis River drains the north-eastern part of the Annapolis Valley into the Minas Basin. Each of these estuaries has slumping banks of brown clay exposed at low tide, when megaripples are exposed in the intertidal zone.

Rising tides form diurnal tidal bores that flow upstream at mid-tide. During spring tides a tidal bore rushes upstream into the rivers that enter Cobequid Bay – particularly the St Croix River near Windsor, the Shubenacadie and the Salmon River at Truro. The river channels are bordered by high artificial levees. There are several historic river ports, including Hantsport on the River Avon and Port Williams on the Cornwallis River. Wolfville Harbour drains out at low tide. The western coast of Minas Basin has sloping cliffs of red Triassic rock, which passes northward beneath Jurassic basalt forming the bold cliffs of the Blomidon Peninsula, which show columnar structures (>Fig. 2.2.10). This marks the beginning of the long ridge of Jurassic basalt (North Mountain) which runs WSW to Brier Island, with marine gaps at Digby Gut, Petit Passage and Grand Passage. It forms a cuesta with an escarpment facing SSE across the Annapolis Valley, the tidal Annapolis Basin, the Digby isthmus and St Marys Bay, and a dip-slope declining towards the Bay of Fundy coast, where seaward dipping layers of basaltic lava outcrop on the shore. At its NE end it curves sharply round Scots Bay so that the escarpment forms the southern coast of the Minas Channel. It ends at Cape Split, with an outlying stack. There are roaring turbulent currents (up to 4 m/s) here on the middle of a rising tide. From Cape Split the Scots Bay coastline consists of seaward-dipping basalt with coves at the mouths of small valleys drained by short streams and wide intertidal sand and mudflats. The streams lengthen as the coast curves

New Brunswick and Nova Scotia

2.2

⊡⊡ Fig. 2.2.10 Cape Blomidon, showing Long Mountain basalt over Triassic sediment. (Courtesy Geostudies.)

southward, then WSW to the long straight section along the south coast of the Bay of Fundy. The view from the hinterland slope is across the Bay of Fundy to outlying Isle Haute, Cape Chignecto, Cape D’Or and Cape Split to the east. Valleys descend to small bays, several of which have small harbours for fishing boats, sheltered by massive, boulder-armoured breakwaters, as at Halls Harbour. Low cliffs and sloping ramps of basalt with scattered boulders are interrupted by curving bays of well-sorted grey shingle derived from basalt and glacial drift, as at Hampton Beach. Parkers Cove has another harbour with stone breakwaters and massive granite boulders on the eastern side. The Annapolis Valley runs ENE-WSW, and the Annap­ olis River has a meandering estuary that opens into the  tidal Annapolis Basin. At Annapolis Royal the tide range is about 7 m, utilised by a hydroelectric station in a  narrow sector of the Annapolis River. The Annapolis Basin widens towards Digby, and Digby Gut is a steepsided channel through the North Mountain coast range, forming an outlet to the Bay of Fundy (>Fig.  2.2.11). The semidiurnal tide has a range of 8 m at Digby, and a broad shore zone is exposed around the Annapolis Basin at low tide. Basalt layers dip NW on either side of Digby Gut, and sloping shore ledges follow the dip below the lighthouse at Point Prim on the western side. The Bay of Fundy coast continues to the WSW with valley-mouth coves between headlands with ramps of dipping basalt. Broad Cove

c­ ontains a harbour enclosed by breakwaters in a bay with a grey cobble beach curving round to lava cliffs. Gullivers Cove is a bay with an upper beach of cobbles and boulders, mainly basalt and also of other rock types from glacial drift, partly held by a backshore stockade of timber, with a timber groyne. The North Mountain cuesta is set forward along a fault line at Gullivers Head, and the next section is known as Digby Neck. There are low cliffs cut in glacial drift and successive lava ledges along the shore. The outermost of these on Gullivers Head is an almost flat shore platform, draped with seaweed. Goose­ berry Cove has a grey boulder and cobble beach with a high cliff of seaward dipping basalt on the SW side. The cliffs and storm-piled beaches indicate that wave action can be strong on this Bay of Fundy coast, but the large tides reduce effective wave energy. At Centreville a valley occupied by glacial drift contains incised Trout Creek, which flows out through a gravel beach into a tidal harbour, enclosed by tall breakwaters with wooden pilings, stone filling and boulder armouring. The sloping basalt coast continues WSW to the gap at Petit Passage, resumes in Long Island to Grand Passage, and ends on Brier Island. The inner coast, facing St. Marys Bay, has low cliffs and bluffs truncating at the base of the basalt dip-slope, small coves and some beach fringes. At The Sea Wall the coast is lined by granite boulders that have been dumped across the mouth of a salt marsh bay. The coast then curves out to Red Head, where cliffs have been cut in

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red Triassic sandstone, which underlies the Jurassic basalt. At the head of St. Marys Bay the coast of the Digby isthmus is low-lying on glacial drift, and at low tide the northern part of the bay has extensive brown gravelly mud and sand flats. The southern shores of St. Marys Bay are at first lowlying and marshy, with cliffs cut in glacial drift at Gilberts Head. To the WSW, marshes give place to beaches and longshore spits of sand and shingle derived from glacial

drift, as at Brooks Beach. The estuary of Sissiboo River enters below Weymouth. At Belliveau there is a little harbour, and a long curving shingle beach to west becomes a lobate foreland with multiple beach ridges at Pointe a Major (>Fig. 2.2.12). To the south, Rivière Grosses Coques flows out under an old road bridge to a broad salt marsh – Marais de Grosses Coques, behind a shingle barrier spit – the ­estuary opening near the SW end. At Ticken Cove a grassy dune ⊡⊡ Fig. 2.2.11 Digby Gut, a gap in the Long Mountain ridge. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.12 Gravel beach and parallel beach ridges at Pointe a Major. (Courtesy Geostudies.)

New Brunswick and Nova Scotia

barrier fronts a bay lagoon, and Church Point shelters a salt marsh, with gravel shoals backed by a beach with dunes fronting a lagoon. At Little Brook there are gentle grassy slopes down to the coast, and at Comeauville and Saulnierville low cliffs in glacial drift are fronted by shingle beaches. Southward along the coast, wave energy increases with widening fetch and diminishing tides as the Bay of Fundy opens into the Gulf of Maine and the Atlantic Ocean. Tide range diminishes from 8.5 m at Digby Gut to 5 m at Cape St. Marys. At Smugglers Cove (>Fig. 2.2.13) there is a bold cliffed coast cut in Lower Palaeozoic mudstones with steep to vertical cleavage. Some of the cleavage planes are exposed in the cliff face, disintegrating into angular gravel, which forms a cobble beach. At Cape St Marys the lighthouse stands on cliffs cut in the cleaved mudstone, which now trends at right angles to the coast so that there are narrow promontories between deep clefts and an arête. South of Cape St Marys is a fishing harbour, and a long curving sandy beach backed by dunes in Mavilette Beach Provincial Park. This ends in a cliff cut in brown glacial drift, followed by more cliffs in grey mudstones. There are several lagoons behind grassy barriers of sand and shingle. At Salmon River the stream has a narrow outlet to the sea, bordered by a wall of wooden piling. The coast runs southward with many inlets at the mouths of river valleys, past Yarmouth to Chebogue Point. Tide ranges diminish along the south-west coast Nova Scotia to 3.3 m

⊡⊡ Fig. 2.2.13 Bedding planes exposed in cliff face at Smugglers Cove. (Courtesy Geostudies.)

2.2

at Yarmouth, and wave action becomes more effective as the tides decline. There are cliffs cut in glacial drift and derived beaches and spits of sand and gravel. Long narrow peninsulas, as at Pinkneys Point, separate marine and estuarine inlets and the Tusket Islands are an outlying rocky archipelago. The Atlantic coast of Nova Scotia is indented, with long narrow peninsulas and rocky ridges separating elongated marine and estuarine inlets at valley mouths and many rocky islands. The rock formations are mainly Lower Palaeozoic (Cambrian and Ordovician) metasediment (sandstones and shales with some schist and gneiss), and several granite intrusions. The inlets were river valleys deepened by glaciation (an ice sheet that flowed N-S), then submerged by the Late Quaternary marine transgression. There is a mantle of glacial moraine with some drumlins, locally truncated by sea cliffs. Beaches, spits and barriers of sand and gravel have been derived largely from these glacial drift deposits. Coastal waters are generally ice-free in winter, but ice may form in estuaries and lagoons. Wave action, mainly from the SE, is strong in frequent stormy periods, and there are occasional storm surges of at least 1 m. Tides are semidiurnal, mean tide range diminishing from 3.7 m in the SW to 1.5 m at Halifax and less than a metre in the north of Cape Breton Island. Lahave River, downstream from Bridgewater, opens to the sea beside the Kingsburg peninsula. There are barrier beaches and spits enclosing lagoons, as in Hartling Bay,

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and at Rose Bay a spit shelters a salt marsh on Conrad Island. North of Ovens Point is gravelly Cunards Beach, and there are cliffs cut in slaty mudstones with clefts and Ovens Caves, cut out along cleavage planes. Some of the caves were adits from historical gold mining. St Margaret’s Bay is wide, with many wooded islands, and Peggy’s Cove to the east has sloping granite shores with glacial striations and a prominent lighthouse. At Sambro there are intricate peninsulas and inlets with seaweeddraped bouldery shores and occasional sandy beaches. Crystal Crescent Beach has grey sand and scattered boulders, with granite shore outcrops and patches of grassy marsh. Around Ketch Harbour the inlets are edged with pale grey granite boulders, and Portuguese Cove and Herr­ ing Cove are more inlets fringed by granite boulders and small sandy beaches. In all these locations the granite has glacially-smoothed slopes down into the sea, without shore platforms. Halifax and Dartmouth stand on either side of a broad marine inlet, Halifax Harbour. Upstream, through The Narrows, is the Bedford Basin, a former lake that became submerged during the later stages of the Late Quaternary marine transgression. As on McNab’s Island, erosion of bluffs in glacial drift has occurred, in the marine inlet south of Halifax (Manson 2002). On the outer coast east of Halifax there have been rapid changes. Cliffs are receding, particularly on drumlin headlands and islands, and there has been associated landward migration of beaches, spits and barriers, with overwashing and breaching which results in the movement of sediment into lagoons that have become more saline, with replacement of freshwater swamps by salt marshes. These rather rapid changes have been accelerated by the rising sea level indicated on the Halifax tide gauge (Taylor et al. 1996). Halifax Harbour is a marine inlet. To the east, the peninsula south of Dartmouth has an outer coast with headlands at Hartlen Point and Osborne Head. An intervening bay, Cow Bay, has a gravelly barrier beach (Silver Sands Beach) with multiple cobble beach ridges enclosing a lagoon – Cow Bay Lake – in a former embayment. On the eastern side is a cliffed drumlin, and there are records of former drumlins, now planed off as shoals, which were the main source of sediment for this barrier beach. In 1900 Cow Bay Lake was freshwater, with an outlet to the sea at the western end of the beach. However, after this was closed by storm waves, the lake overflowed and cut a new entrance to the east. A breakwater was then built to provide a boat harbour, and Cow Bay Lake has since been brackish. The barrier beach was reduced in width by sand and gravel extraction

in the 1950s and 1960s, and has migrated landward as the result of storm wave overwash. Cole Harbour is a branched embayment with bouldery shores with patches of grass marsh, backed by woodland that was damaged by Hurricane Juan on 29 September 2003. There is a sandy barrier spit at its mouth (Rainbow Haven) and extensive salt marshes, with some sandy beaches. To the east is Conrad Head, with a sandy cliff, and beyond it Conrad Beach – a sandy barrier beach backed by dunes and the marshes of West Marsh – curving round to cuspate Fox Point. Like Silver Sands Beach this was derived from sediment eroded from cliffed drumlins, and reduced by the quarrying of beach gravel in the mid-twentieth century. The barrier was breached during a storm in 1962 and remained open until 1989–1990, since when dunes have developed across the former inlet. The coastal features can be viewed from the top of Lawrencetown Head, a cliffed drumlin. The cliff has been receding at up to 0.8 m/year. Lawrencetown Beach has low grassy dunes, and the shingle beaches are fronted by sandy shores exposed at low tide. The sediment has been derived from cliffed drumlins and morainic drift deposits. A barrier beach of sand and gravel encloses elongated Law­ rencetown Lake. At Fishermans Reserve there are shacks and boats on a projecting spit, the north shore protected by massive boulders. To the east Rudeys Head is a cliffed drumlin. The barrier-enclosed lagoons are like New England ponds, marshy with Typha along their inner shores. Gaetz Head is another truncated drumlin headland (>Fig. 2.2.14). When these drumlin headlands were cut back, the barrier beaches derived from them have retreated landward. Cape Antrim, a drumlin headland, disappeared as the result of cliff recession at rates that attained more than 12 m/year in the 1970s, and the adjacent barrier beach receded at a similar rate. Chezzetcook Inlet has shoals and spits trailing NW behind islands. It is typical of the wave-dominated salt marsh fringed estuaries with only minor fluvial inflow on the coast east of Halifax (Carter et al. 1992). There are salt marshes with Spartina patens in the upper marsh, and Spartina alterniflora at lower levels, fringed by mudflats exposed at low tide. Some of the sediment has been supplied by rivers, but much has come from marginal shore erosion, soil creep on slopes (increased following forest clearance) and overwash of spits and barriers at the seaward end. Story Head Beach is a low gravel barrier beach 40–60  m wide that extended from Story Head, a cliffed drumlin, round to a shoal that is all that remains of the former Miseners Head drumlin (Carter et al. 1990). It has

New Brunswick and Nova Scotia

2.2

⊡⊡ Fig. 2.2.14 Gaetz Head, a cliffed drumlin. (Courtesy Geostudies.)

retreated landward as the result of frequent overwash by storm waves (forming washover lobes), and has been breached at its eastern and western ends, where trailing spits have formed. The development of these trailing spits on what is now the Story Head drumlin island and landward from the shoal marking the former Misiners Island drumlin, illustrates the transition from the original swashaligned barrier beach to a series of drift-aligned ridges transverse to the general outline of the coast. Miseners Long Beach is a relatively stable high stormpiled barrier beach of poorly sorted subangular grey pebbles and cobbles behind a sandy foreshore. The barrier beach is steep-fronted and reflective, and often has large beach cusps. Winter shore ice is sufficient to protect the upper beach face and displace the breaker zone seaward. The barrier beach stands in front of Misemers Lake. East of Petpeswick Inlet is Martinique Beach – another long barrier spit with slightly cliffed grassy dunes and a surfing beach. The indented rocky coastline continues north-east to Cape Canso, which shelters Chedabucto Bay. The narrow Strait of Canso separates Cape Breton Island and runs through to St. Georges Bay in the southern Gulf of St. Lawrence. It is crossed by the Canso Causeway, completed in 1955. The south-east coast of Cape Breton Island is again rocky and exposed to Atlantic swell and storm waves. St Peters Island, near Rockdale, has a cuspate spit on its northern (lee) coast, matched by a paired cus­ pate ­spit  at Potheir Point on the mainland. To the east a double tombolo links Michaux Point to the mainland near Gracieville.

The northern coast of Cape Breton Island is strongly influenced by the SW–NE trend of ridges and valleys following Appalachian folds and faults, with intricate bays and straits along submerged valleys. Crystalline rocks form uplands, and lowlands follow corridors of Carboniferous rock, as at Sydney Harbour. Scatari Island is hilly, with minor cliffing on its coasts. Glace Bay has a mean spring tide range of 1.0 m, and Table Head to the north has receding cliffs and isolated stacks bordering a low plateau. Johnson (1925: Fig. 152) provided a set of six photographs showing the dissection of a head­land near Table Head between 1900 and 1921. At Briton Cove there is a good example of a sloping cliff on steeplydipping Carboniferous sandstone, the cliff face coinciding with the bedding plane on some sectors (Johnson 1925: Fig. 83). West of Sydney Harbour Cranberry Head, Merrit Point and Point Aconi have receding cliffs in Carboniferous sediment, including coal seams. At Ingonish a barrier spit extends across South Bay, breached near its southern end, and a long curved barrier spit backs Aspy Bay at Dingwall, which has a mean spring tide range of 0.7 m. Beyond Aspy Bay a small peninsula runs out to Cape North beside Cabot Strait, at the entrance to the >  Gulf of St Lawrence.

4.  Sable Island Lying east of Nova Scotia is Sable Island (44°S, 60°W), an unusual arcuate deposit of unconsolidated sand, 200 km

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south-east of mainland Nova Scotia on the outer margin of the continental shelf. It is only 35 km long, with a maximum width of about 2 km. It has been shaped largely by the prevailing westerly winds and waves in the Gulf of St. Lawrence. Vegetated dunes are found in the central section. Cameron (1965) documented changes in the morphology and size of the end points of the island.

References Amos CL, Long BFN (1980) The sedimentary character of the Minas Basin, Bay of Fundy. In: McCann SB (ed) The Coastline of Canada, Geological Survey Paper Canada 80–10:153–180 Cameron HL (1965) The shifting sands of Sable Island. Geogr Rev 44: 363–376 Carter RWG, Orford JD, Forbes DL, Taylor RB (1990) Morphosedimen­ tary development of drumlin-flank barriers with rapidly rising sea level, Story Head, Nova Scotia. Sediment Geol 69:117–138 Carter RWG, Orford JD, Jennings SC, Shaw J, Smith JP (1992) Recent evolution of a paraglacial estuary under conditions of rapid sea level rise: Chezzetcook Inlet, Nova Scotia. Proc Geol Assoc 103:167–185

Dalrymple RW, Knight RJ, Middleton GV (1975) Intertidal sand bars in Cobequid Bay (Bay of Fundy). In: Cronin LE (ed) Estuarine Research, Vol 2, Academic Press, New York, pp 293–307 Dalrymple RW, Knight RJ, Lambiase JT (1978) Bedforms and their hydraulic stability in a tidal environment, Bay of Fundy, Canada. Nature 275:100–104 Johnson DW (1925) The New England-Acadian Shoreline, Wiley, New York Knight RJ (1980) Linear sand bar development and tidal current flow in Cobequid Bay Bay of Fundy Nova Scotia. In: McCann SB (ed) The Coastline of Canada, Geological Survey Paper Canada 80–10: 123–152 Manson GK (2002) Subannual erosion and retreat of cohesive till bluffs, McNab’s Island, Nova Scotia. J Coast Res 18:421–432 Owens EH, Bowen AJ (1977) Coastal environments of the Maritime Provinces. Maritime Sediments 13:1–13 Shaw J, Ceman J (1999) Salt marsh aggradation in response to late Holocene sea level rise at Amherst Point, Nova Scotia. The Holocene 9:439–451 Taylor RB, Shaw J, Forbes DL, Frobel D (1996) Field Trip Guidebook. Eastern shore of Nova Scotia. Coastal response to sea-level rise and human interference. Natural Resources, Canada

2.3  Gulf of St. Lawrence

Eric Bird

1.  Introduction

coastline is generally fringed by shore ice in winter, and exposed to wave action for up to 8 months in summer: there are variations in the length of winter and summer conditions between the southern and the northern parts. Atlantic ocean swell reaches the south coast of Newfoundland, and to a lesser extent the northeast-facing coasts of the island of Newfoundland and Labrador. Generally the wave regime on these coasts, and in the relatively protected Gulf of St. Lawrence, is dominated by local storm-generated waves.

The coast of the Gulf of St. Lawrence includes the St. Lawrence estuary and the south coast of Quebec (> Fig.  2.3.1). The southern coast of the Gulf is readily accessible and has been studied in detail, but much of the northern coast is rocky, barren and subarctic, difficult of access and little known geomorphologically. The Gulf of St. Lawrence is a microtidal sea environment with a maximum SW-NE extent of 800 km. The

⊡⊡ Fig. 2.3.1 Coastal regions of eastern Canada. 1-Bay of Fundy, 2-Gulf of St. Lawrence (south), 3-Gulf of St. Lawrence (north), 4-St. Lawrence Estuary, 5-Atlantic Nova Scotia and southern Newfoundland, 6-E. Newfoundland (including Labrador). Provinces: A-Quebec, B-New Brunswick, C-Nova Scotia, D-Prince Edward Island, E-Newfoundland. (Courtesy S.B. McCann.)

80 W

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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2.  Magdalen Islands

The north coast of Prince Edward Island has cliffs cut in glacial drift over soft red Permian sandstone (> Fig. 2.3.3) and beaches that pass west from Stanhope Beach into a chain of barriers, spits and barrier islands with grassy dunes (> Fig.  2.3.4) (McCann 1979). The barriers have developed across embayments and estuaries to enclose lagoons such as Tracadie Harbour and Malpeque Bay, and are basically transgressive, though downdrift, some have prograded dune ridge sequences. Most of the barrier coastline has spring tides smaller than 1.6 m, and tidal currents are only important near inlets. There are short period northeasterly waves generated by occasional storms within the Gulf, which drift sediment westward along the beaches; but the prevailing westerly winds blow from land to sea, and have little effect on the barriers except on sectors with westerly exposure. Tidal inlets that breach the barriers vary in size from small, temporary inlets, only tens of metres wide and 2–3 m deep, to large permanent inlets up to 1 km wide and 15 m deep. They separate barrier islands such as Hog Island which has low grassy dunes, a wide beach with small barchans that move northward to the sea, lumber on the backshore, and Spartina marshes on the lagoon shore, locally eroded to expose peat. Further west a high coast has maintained its dimensions during slow barrier retreat. Longshore transport rates are high at this location, and

Out to the northwest are the Magdalen Islands (> Fig. 2.3.2), where two long barrier beaches and dunes form tombolos joining a series of rock islands and enclosing shallow central lagoons (Owens and McCann 1980). Together with complex terminal spits they make a continuous narrow land area 70 km long. On the west coast of these islands are three sand bars, the outermost 700–1,000 m offshore in a water depth of 4–6 m.

3.  South Coast, Gulf of St. Lawrence West of Cape North, the coast of Cape Breton Island turns southward and faces the Gulf of St. Lawrence. Geologically Cape Breton Island and the north coast of Nova Scotia are part of the northern section of the Appalachian system, and consist predominantly of deformed Palaeozoic sedimentary rocks trending SW-NE. St. Georges Bay opens northward into the Gulf of St. Lawrence and west of Cape St. George, a generally low-lying north coast of Nova Scotia, passes along the southern side of Northumberland Strait, to the south of Prince Edward Island. There are several valley-mouth rias on this microtidal coast.

⊡⊡ Fig. 2.3.2 The Magdalen Islands. (Courtesy Geostudies.)

Bed rock Spits and Barriers Longshore transport

0

20 km

Gulf of St. Lawrence

2.3

⊡⊡ Fig. 2.3.3 Cliffs cut in Permian sandstone, Orby Head, Prince Edward Island. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.4 Low dunes behind sandy beach, with small tidal inlet at Dalvay Beach, north coast of Prince Edward Island. (Courtesy Geostudies.)

nearshore erosion of the sea floor appears to be contributing new sediment (Armon and McCann 1977). The south coast of Prince Edward Island, facing Northumberland Strait, is embayed, with broad promontories showing cliffs cut in soft Lower Permian clay and sandstone, as at Borden (> Fig. 2.3.5). Charlottetown has a mean spring tide range of 1.8 m, and muddy shores are exposed at low tide. At Gallas Point the receding cliffs are

cut in glacial drift, fronted by broad shore platforms on Lower Permian sandstone (> Fig.  2.3.6) which display weathering and abrasion features, and have a partial cover of marine algae. In the bays and estuaries are salt marshes dominated by Spartina and red muddy intertidal areas. The southern coast of Northumberland Strait in Nova Scotia and New Brunswick shows similar features bordering a broad low plateau of glacial drift. At Cape Jourimain

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⊡⊡ Fig. 2.3.5 Cliffs at Borden on the south coast of Prince Edward Island. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.6 Shore platform at Gallas Point, south coast of Prince Edward Island. (Courtesy Geostudies.)

there are slumping bluffs of till fronted by shore outcrops of almost horizontal Lower Permian pink sandstone exhumed from the drift cover. Sectors of low cliff are cut into the sandstone, as at Cape Bald (Cap Pelée). There is ­extensive nearshore bar topography, as in Kouchibouguac Bay, where there are two bars, the outermost 200–300 m offshore in about water about 2 m deep at low tide

(Davidson-Arnott and Greenwood 1976). North from Point Sapin, beyond the shelter of Prince Edward Island, the coast becomes exposed to stronger wave action across the Gulf of St. Lawrence, and there are spits and barrier beaches (Johnson 1925: 147). Buctouche Bar is a compound recurved spit that has: grown intermittently southward. The mouth of Miramichi Bay is encumbered with

Gulf of St. Lawrence

these depositional forms, with intervening tidal inlets maintained by strong currents, as at Portage Gully (Rein­ son 1977). Portage Island is a barrier island with recurved beach ridges indicating its southward growth. The inlets have large ebb and flood tidal deltas, and have maintained their positions throughout the period of historical record (about 200 years), though there have been changes in the pattern of channels and shoals. Landward transfer of sediment through these low, narrow transgressive barriers of New Brunswick has been by overwash, producing fans of sediment on the landward side, and through tidal inlets, past and present. The beaches and barriers come to an end at Miscou Point, on the southern side of the mouth of the broad Baie des Chaleurs which narrows westward to an estuary at Campbellton. On the northern side is a series of bays and headlands, extending to Bonaventure Island. Beyond this Le Rocher Percé (the Pierced Rock) is an elongated island of steeply dipping to vertical limestone bordered by cliffs up to 100 m high and penetrated by a tall natural arch. A stack at the eastern end is the remains of an earlier archway that collapsed in the nineteenth century. There is a broad gravelly intertidal zone. The Murailles to the west, on the southern side of Mal Bay, are vertical cliffs truncating a ridge of steeply dipping limestone that runs along the coast. They have cut back into the landward slope, and so are receding cliffs of diminishing altitude (Johnson 1925: 200–202).

⊡⊡ Fig. 2.3.7 Strandlines on salt marsh, St. Lawrence estuary. (Courtesy Geostudies.)

2.3

Mal Bay is backed by a barrier spit, and to the north is the Baie de Gaspé, the estuary of the Dartmouth River. Beyond this is the long narrow promontory that ends in Cape Gaspé. Here spectacular cliffs are fronted by shore platforms cut in shale, the outer edge of which has been trimm­ ­ed back by shore ice erosion, particularly when the shore ice is disrupted by spring storm waves (Trenhaile 1978).

4.  St. Lawrence Estuary Cape Gaspé marks the beginning of the St. Lawrence estuary, which has a long curving southern coast backed by the rising slopes of the Shickshock Mountains. Mean spring tide range increases from 1.0 m at Gaspé to 3.7 m at Rimouski and 4.9 m at Quebec as the estuary narrows to the outflow from the Great Lakes (> Great Lakes (Canada)). The intertidal zone widens correspondingly. Shore ice develops and persists for four months in winter, when there is a thick cover of mobile pack ice offshore. On the north coast of the Gaspé Peninsula are seaward sloping intertidal shore platforms cut across dipping strata that run roughly parallel to the coastline. The movement of shore ice to and fro has contributed to the abrasion of these platforms. There are salt marshes (schorre) threaded by high tide strandline litter (> Fig. 2.3.7) and strewn with pebbles, ­cobbles and large ice-­transported boulders with scoured furrows (> Fig. 2.3.8) (Dionne 1972, 1989).

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Boulder barricades are elongated rows of boulders emplaced just offshore on arctic or subarctic coasts by the grounding of boulder-laden ice rafts in the nearshore zone when the ice breaks up in spring. They have been described from the shores of the Gulf of St. Lawrence, notably at Cap la Baleine (Dionne 2003). There are also boulder pavements, assemblages of rocks in the intertidal zone on cold climate coasts flattened by impaction

beneath shore ice, usually smoothed and often striated (Hansom 2005). On the shores of the St. Lawrence estuary erratic boulders have been derived partly from the Appalachian Mountains and partly from the Canadian Shield to the north, and delivered by glaciers. Many boulders have been transported across the estuary from the north shore by drifting ice (Dionne 2002). Shore ice breakup in spring is

⊡⊡ Fig. 2.3.8 Boulder pushed by shore ice, Trois Pistoles, St. Lawrence estuary. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.9 Displaced raft of salt marsh turf, St. Lawrence estuary. (Courtesy Geostudies.)

Gulf of St. Lawrence

accompanied by erosion as storm waves push it across the salt marsh, displacing rafts of Spartina turf (> Fig. 2.3.9) Abrasion by shore ice has also contributed to the shaping of shore platforms, as at Trois Pistoles (Dionne and Brodeur 1988). At low tide a wide zone of muddy rocks and gravel is exposed in front of the salt marsh. The north coast of the St. Lawrence estuary is more embayed, with outcrops of hard Pre-Cambrian rock along the southern edge of the Canadian Shield and a coastal plain of glacial drift and glacifluvial deltas east from Cap Colombier. Beach sediment are often scarce, and largely derived from the erosion of glacial drift deposits. Emerged coastline deposits are found, particularly along the St. Lawrence estuary. Tide ranges diminish eastward to 2.6 m at Sept Iles, and as the estuary widens wave action becomes stronger. The eastern boundary of the St. Lawrence estuary is arbitrary, but can be taken at longitude 64° 35ʹ where the channels diverge to the Jacques Cartier Passage and the Honguedo Passage, north and south of Anticosti Island.

5.  Anticosti Island Anticosti Island is generally low and hilly, with broad plains of glacial drift and generally smooth coastlines with low cliffs, sand and gravel beaches and salt marshes. A series of 22 terraces planed across local geological structures at up to more than 400 feet above sea level was described by Twenhofel and Conine (1921), who considered them to be of marine origin, a view questioned by Johnson (1925: 144–146). There are separate discussions in the chapters on > Newfoundland and > St. Pierre et Miquelon.

6.  North Coast, Gulf of St. Lawrence The south-facing Quebec coast in the Jacques Cartier Passage and to the east is generally rocky and indented. The shore ice season increases from 3 to 4 months in the St. Lawrence estuary to 6 months in Labrador. An ice foot forms along these coasts in winter, and persists to the breakup season. The offshore ice cover in the south is usually dense, slowly moving pack ice, but in the north there is commonly a broad zone of fast ice extending seaward for many kilometres between the shore ice and the mobile pack. The presence of shore ice for many months each year reduces the period when wave processes are effective, but introduces other effects on the shore, ­particularly at breakup when erosion, transport and deposition by drift ice are key processes. Storm waves occur frequently.

2.3

The Natashquan River flows down to an estuary in a broad deltaic protrusion at Old Post Point, and to the east the low coast has numerous rias between rocky headlands. East of Harrington Harbour it is fringed by numerous small rocky islands. Boulder barricades are found along the mean low tide line and nearshore gravels have been derived from ice-rafted sediment. Wave energy diminishes as the coast passes alongside the Strait of Belle Isle. At Forteau the eastern border of Quebec crosses the coast, and the east coast of Labrador is within the province of > Newfoundland.

7.  Eastern Labrador At York Point on the northern side of the Strait of Belle Isle the Labrador coastline turns northward to Domino, then northwestward. It is a generally steep, rocky coastline, cut by numerous fjords and facing the northwest Atlantic Ocean. The tide range is small (1.9 m at Nain). Storm waves are frequently generated during the passage of cyclonic storms from west to east across the northern Atlantic and the Labrador Sea. There are pocket beaches of sand and gravel, derived from glacial drift, in some of the coves. Boulders on the shore are displaced by ice rafting (Hansom 2005). The effects of glaciation are evident throughout the Labrador region, which was covered by ice as recently as the late Wisconsin. Glacial erosion produced the fjords of eastern Labrador, glacial deposits cover lowlands and much of the shallow continental shelf, and changes in relative sea level due to deglaciation and isostatic rebound are important everywhere. The rocky coast with numerous fjords continues northward to Cape Chidley.

References Armon JW, McCann SB (1977) Longshore sediment transport and a sediment budget for the Malpeque barrier system, Southern Gulf of St. Lawrence. Can J Earth Sci 14:2429–2439 Davidson-Arnott RGD, Greenwood BB (1976) Facies relationships on a barred coast, Kouchibouguac Bay New Brunswick. In: Davies RA, Ethington RL (eds) Beach and nearshore sedimentation. Society of Economic Paleontologists and Mineralogists, Special Publication 24:149–168 Dionne JC (1972) Caracteristiques des schorres des regions froides, en particulier de 1¢estuaire du Saint Laurent. Z Geomorphol 15:137–180 Dionne JC (1989) The role of ice and frost in salt marsh development. A review with particular reference to Quebec, Canada. Earth Sci Rev 26 Dionne JC (2002) Dolomite erratics at Rivière-Blanche, south coast of the maritime estuary of the St Lawrence (in French). Can J Earth Sci 39:1239–1255

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Dionne JC (2003) The boulder barricade at Cap la Baleine, north shore of Gaspé; peninsula (Québec): nature of boulders, origin and significance. J Coast Res 18:652–661 Dionne JC, Brodeur D (1988) Erosion of rocky platforms by shore ice (in French). Z Geomorphol 32:101–115 Hansom JD (2005) Boulder pavements. In: Schwartz ML (ed) Encyclopedia of Coastal Science, Springer, Dordrecht, pp 208–210 Johnson DW (1925) The New England-Acadian shoreline. Wiley, New York McCann SB (1979) Barrier islands in the southern Gulf of St. Lawrence. In: Leatherman SP (ed) Barrier Islands from the Gulf of St Lawrence

to the Gulf of Mexico. New York, Academic Press, pp 29–63 Owens EH, McCann SB (1980) The coastal geomorphology of the Magdalen Islands, Quebec. In: McCann SBG (ed) The Coastline of Canada, Geol Surv Can Pap 80–10:123–152 Reinson GE (1977) Tidal current control of submarine morphology at the mouth of the Maramichi estuary, New Brunswick. Can J Earth Sci 14:2524–2532 Trenhaile AS (1978) The shore platforms of Gaspé. Quebec. A Assoc Am Geogr 68:95–114 Twenhofel WH, Conine WH (1921) The post-glacial terraces of Anticosti Island. Am J Sci 1:268–278

2.4 Saint-Pierre-et-Miquelon Islands Jean-Marie M. Dubois 

1.  Introduction The Saint-Pierre-et-Miquelon archipelago is a small French territory, only 242 sq. km, located 22 km from the southern coast of Newfoundland, between 46° 45' and 47° 10' N and 56° 05' and 56° 25' W. It constitutes an enclave on the Saint-Pierre Bank within Canadian territorial waters. The archipelago is formed by a main island (Saint-Pierre) to the east, where the prefecture and administrative centre are located, by two more important islands (Miquelon and Langlade, joined by a N-S tombolo) as well as six other islands (Le Cap, Grand Colombier, Verte, aux Marins, aux Vainqueurs and aux Pigeons islands) and various islets and reefs (Aubert de la Rue 1970) (>Fig. 2.4.1). The Labrador Current, which flows through the archipelago from east to west, is responsible for cool and foggy summers (85–120 days) and the average annual tem­ perature in this subarctic marine environment is barely 5.5°C. Average annual precipitation is about 1,424 mm – about a third as snow. Winds exceed 60 km/h for almost half the year, and storms are frequent. The storm of January 1987 caused up to 5 m of recession on cliffs cut in glacial drift deposits, the February 1988 storm, with winds of up to 160 km/h cut back cliffs by up to 14 m and the storm of December 1989 caused 7 m of erosion (Rabottin 1990). Much of the archipelago has scrubby vegetation following long-term deforestation, but woodland occupies 20% of Langlade Island. Peatland is extensive, and heathland and grassland are found over coastal sand and pebble areas. The main algae found along the coasts are laminaria, brown algae, red algae and lime-secreting algae. The geology is varied (Dubois 2006). Le Cap Island is  mainly composed of metamorphic crystalline PreCambrian rocks (schists, quartzites, paragneiss and migmatites), with a small intrusive granitic stock. Miquelon Island consists of slightly metamorphosed post-­Ordovician volcanic rocks, mainly rhyolites with some basalts, andesites and breccias. Langlade Island has slightly meta­morphosed Appalachian Palaeozoic sedimentary rocks (sandstones, schists, phyllites, limestones, quartzites and conglomerates) trending SW-NE with faults. Saint-Pierre and adjacent islands are composed of post-Ordovician intrusive rocks, mainly rhyolites, brecchia and tuffs with a few andesites.

The islands are mostly rocky with superficial deposits of glacial till, wind-blown sand, alluvium, weathered rock, landslide debris, colluvium and peat. The islands are the remains of an ancient erosion surface except for Langlade Island, which has a tabular relief. On the main islands there are a number of summits: Morne de la Grande Montagne on Miquelon island (240 m), Le Trépied on Saint-Pierre island (207 m), Tête de Cuquemel on Langlade Island (190 m) and a hill with an altitude of 153 m at the NE end of Le Cap island. The thick superficial deposits are located mainly on Miquelon and Langlade Islands, while silty clay till, superficially reworked by the sea, is present over most of the periphery of Miquelon island and in NW Langlade Island, often affected by landslides. Tombolos of sand and gravel link Langlade and Miquelon Islands, Miquelon and Le Cap Island, and will eventually link aux Vainqueurs and aux Pigeons islands. The archipelago was covered by ice from Newfoundland during the Middle Wisconsin, as shown by the nature of the erratics, but not during the Late Wisconsin. Glaciation left moraines to the south of Saint-Pierre Island and till over most of the islands. A subsequent marine submergence, around 38,000 years bp left shore deposits and indications of marine erosion up to at least 25 m above sea level. The southern limit of the last glaciation was along the south coast of Newfoundland (Piper and Macdonald 2001), and evidence of periglaciation – including frost wedges – is found on all the islands (Lauriol and Dubois 1981). However, recent shore formations at 10 m on the north of Miquelon Island, 6.4 m in the north of Langlade Island and present sea level on Saint-Pierre Island, could indicate differential Post­ glacial emergence. At Quine Point, north of Grand Barachois, the base of a peat bog, 1 m below present sea level, was dated at 8,460  ±  150 years bp, indicating recent submergence (Lauriol and Dubois 1981). The base of another peat bog close to present sea level in the De Savoyard cove, on the south coast of Saint-Pierre Island, had an age of 2,000 ± 60 years bp. Global warming is expected to cause submergence of about 2 mm/year, taking account of continuing glacioisostatic subsidence (shaw and Forbes 1990). In addition to wave action there are longshore currents which have contributed to the shaping of the tombolos on Langlade, Miquelon and Le Cap Islands (Dubois 1980).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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       Saint-Pierre-et-Miquelon Islands

⊡⊡ Fig. 2.4.1 The Saint-Pierre-etMiquelon archipelago.

Generally, the longshore currents flow in a clockwise direction, but there are local reversals in the vicinity of large capes on the eastern shore of the Langlade isthmus and the tombolo between Miquelon and Le Cap Islands. Sediment are certainly carried offshore by convergent longshore currents to the west of Miquelon Island. There is historical evidence of coastal erosion in the archipelago since the eighteenth century, but there have

also been local phases of accretion, as on the Langlade isthmus. Using landmarks established between 1984 and 1989 and comparisons of the 1952 and 1985 aerial photographs, Rabottin (1990) concluded that 88% of the coasts in superficial deposits on Langlade and Miquelon Islands had receded at an average annual rate of slightly more than 1 m, and that this rate was increasing. On lagoon barriers, erosion sometimes reached 2–3 m/year. On the morainic

Saint-Pierre-et-Miquelon Islands     

shores of Miquelon Island the annual recession rate is smaller, between 0.1 and 0.4 m. Further­more, according to fishermen, it seems that during the last 80 years winters have been less rigorous and the winter ice-foot offers less littoral protection than before (Lauriol et Dubois 1981). There are only nine areas of accretion, all on the Langlade isthmus and the tombolo between Miquelon and Le Cap Islands. Rabottin (1990) found that the average rate of accretion was slightly more than 1 m/year. Accretion occurs near sediment sources such as eroding cliffs and around lagoon outlets. The rocks of the cliffs are relatively friable and are disintegrated by cycles of freeze-and-thaw, after which waves remove the fallen material.

2.  Island Coastlines Le Cap Island is linked to Miquelon Island by a sand and pebble tombolo more than 3 km long, and in places, less than 100 m wide. This double tombolo, which was built before the sixteenth century, contains a lagoon with an ­outlet on the eastern side. Since its altitude is only 1–4 m, the tombolo is often inundated during storms coinciding with high tides. During the winter of 1947–1948 a 2.5 m rampart of ice and pebbles was pushed up by storm waves, and this was still visible at the end of May. Le Cap Island has an indented coast with high rocky cliffs which are eroding. Miquelon Island has low morainic and slightly in­­ dented coasts with numerous eroding peat bogs at times. On the west coast cliffs cut into till have a height of 5–10 m, and are receding (Lauriol et Dubois 1981), but a few rocky points and reefs protrude. On the lower east coast some cliffs are rocky, but numerous beach barriers of sand and pebbles enclose lagoons of varying size. The Langlade isthmus is a 12 km long and 100 m to 6 km wide tombolo composed of sand and pebbles with an altitude varying from 1–4 m, and dunes west of Grand Barachois reaching an altitude of 26 m. Captain James Cook mapped an opening about 1.8 km wide in the isthmus in 1763, but this has been closed since 1784. There are parabolic dunes 10–20 m high (locally named buttereaux) at both ends, where spits shelter two large lagoons with outlets towards the east. The west coast is strewn with shipwrecks. It is estimated that there have been more than 650 shipwrecks around the archipelago since 1816 and this coast has become known as the ship necropolis (Sanguin 1983). About 20% of these wrecks occurred between the end of August and the beginning of October, close to the Fall Equinox (Dubois 1980). Langlade Island has a high plateau with an altitude of 130–140 m, incised by many valleys with permanent water

2.4

courses, notably the Belle Rivière – the only real river in the archipelago. The island is surrounded by high receding rocky cliffs up to 25 m high, except in the north where the coastline is lower. This low coastline, cut mainly in till and glacio-fluvial deposits, with recent shore deposits to the NW, is mostly receding (Dubois 1980). In the north and west of Saint-Pierre Island, the coasts are rocky and rectilinear, with cliffs up to 75 m high, sometimes receding. In the south the coast is low, rocky and indented, with reefs and numerous coves containing pebble and boulder beaches (Mailhot et Dubois 1983). There are only a few sandy beaches, notably at De Savoyard and Alumette, where unfortunately sand extraction occurred in the past. Fortunately this practice has been prohibited since 1984 (Rabottin 1990). A fault zone 50–100 m wide may separate the two types of coast. Saint-Pierre Island is separated from Langlade Island by a strait often referred to as La Gueule de l’Enfer (the Mouth of Hell) because of its violent currents. Breakwaters have been built to shelter the harbour at Saint Pierre. Since 1974 breakwaters have also been built to counter erosion in the inner parts of the coves, but erosion is widespread on the coasts of the smaller islands (Dubois 1980).

References Aubert de la Rüe, E (1970) Saint-Pierre et Miquelon. Horizons de France, Paris, 174 Dubois JM (ed) (1980) Géomorphologie des îles Saint-Pierre et Miquelon: premier rapport d’étape. Département de Géographie, Université de Sherbrooke, 47 Dubois J-MM (2006) Saint-Pierre-et-Miquelon: un laboratoire de terrain en miniature pour les géographes physiques. Enjeux géographiques 2:36–44 Lauriol B, Dubois JM (1981) Géomorphologie des îles Saint-Pierre et Miquelon, deuxième rapport d’étape: le Grand Barachois et la côte ouest de l’île de Miquelon. Département de géographie, Université de Sherbrooke, 61 Mailhot R, Dubois JM (1983) Géomorphologie des îles Saint-Pierre et Miquelon, troisième rapport d’étape: classification de la zone côtière des îles Saint-Pierre et Langlade. Département de géographie, Université de Sherbrooke, 40 Piper DJW, Macdonald A (2001) Timing and position of Late Wisconsinan ice-margins on the upper slope seaward of Laurentian Channel. Géographie physique et Quaternaire 55:(2):131–140 Rabottin JL (1990) Saint-Pierre et Miquelon, un archipel d’outre-mer assiégé par les flots. Cahiers Nantais 35–36:37–47 Sanguin A-L (1983) Saint-Pierre et Miquelon. Département français d’Amérique du Nord, Norois, Poitiers, p 110 Shaw J, Forbes DL (1990) Short and long-term relative trends in Atlantic Canada, National Research Council of Canada. Proceedings of the Canadian Coastal Conference 1990, Kingston, Ottawa, Canada, pp 291–305

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2.5  Newfoundland

Eric Bird

1.  Introduction The province of Newfoundland and Labrador includes the large island of Newfoundland (405,720 sq. km). This chapter deals with the island of Newfoundland and the Labrador coast up to Cape Chidley. Much of the coastline is rocky, barren and subarctic in character.

2.  Newfoundland The island of Newfoundland partly blocks the eastern end of the Gulf of St. Lawrence and is separated from Labrador to the north by the Strait of Belle Isle and from Cape Breton Island to the southwest by the wider Cabot Strait. It consists of Appalachian structures, with Palaeozoic rock formations, folded and faulted mainly along SW–NE alignments, in contrast with the Laurentian Shield structure of Labrador. The interior consists of a broad plateau with low hills and a higher ridge (Long Range Mountains) runs NNE along the western peninsula. The rugged island coastline is about 16,000 km long, with peninsulas and bays following the SW–NE trend, and much of the coast is backed by hills or low mountains, except for the lowlands in the northeast, where Pleistocene glaciation has left numerous fiords between narrow peninsulas. There are fiords on the steeper western and southern coasts. The Newfoundland coast is rising as the result of postglacial isostatic recovery. Coastline changes have, however, been generally very slow and the outline charted by Captain Cook in 1775 has been little modified. Newfoundland has cold winters and cool to mild summers. Shore ice is present along much of the seaboard in winter, only the outer coasts of southern Newfoundland being ice free, although even there ice may form in sheltered bays. The Strait of Belle Isle has sea ice for up to 10 months each year and is subject to strong tidal currents. Storms and strong winds are frequent, and the coast is often foggy. Westerly winds over the Gulf of St. Lawrence produce waves that generate northward drift along the west coast, where beaches, spits and barriers of sand and gravel have been derived from cliffs cut in glacial drift. Cuspate

spits have formed on the lee of islands, one of which (Cow  Head) is attached by a tombolo. Coastal dunes ­drifting inland have overrun a forest near Western Brook in the Gros Morne National Park. Long Point and St. George’s are major spit formations. Atlantic Ocean waves reach the southern and eastern coasts, which are generally steep and cliffy. There are many small islands off the coast, including the French archipelago, > St. Pierre et Miquelon, off the south coast. To the southeast is the submerged Grand Banks, sustaining major fisheries.

3.  Labrador At York Point on the northern side of the Strait of Belle Isle the Labrador coastline turns northward to Domino, then northwestward. It is a generally steep, rocky coastline, cut by numerous fiords and facing the northwest Atlantic Ocean. Many hundreds of kilometres of coastline are difficult to reach and remain little explored in terms of coastal geomorphology. The tide range is small (1.9 m at Nain). Storm waves are frequently generated during the passage of cyclonic storms from west to east across the northern Atlantic and the Labrador Sea. There are pocket beaches of sand and gravel, derived from glacial drift, in some of the coves. Boulders on the shore are displaced by ice rafting (Rosen 1979, 2005). The effects of glaciation are evident throughout the Labrador region, which was covered by ice as recently as the late Wisconsin. Glacial erosion produced the fiords of eastern Labrador, glacial deposits cover lowlands and much of the shallow continental shelf, and changes in relative sea level due to deglaciation and isostatic rebound are important everywhere. The rocky coast with numerous fiords continues northward to Cape Chidley.

References Rosen P (1979) Boulder barricades in central Labrador. J Sediment Petrol 49:1113–1123 Rosen P (2005) Boulder barricades. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 204–206

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

2.6  Northern Canada

Wayne Pollard

1.  Introduction The northern coast of Canada is roughly 172,950 km long, within the Arctic (Shaw et al. 1998). The Arctic is marked by extremely cold winters and year-round freezing air temperatures that are together responsible for widespread permafrost and glaciers as well as the seasonal and perennial occurrence of sea ice and snow. The Arctic is also the traditional home for roughly 45,000 Inuit, the only indigenous population residing in 35 coastal communities. The aim of this discussion is to describe the nature and pattern of the northern coastline of Canada from Battle Harbour in Labrador (Labrador Sea) to Demarcation Point at the Yukon–Alaska border (Beaufort Sea), including the Arctic Archipelago, Ungava Bay, Hudson Bay and the northern mainland. A modified version of Owens’ (1994) regional coastal framework is used to subdivide this region into 11 coastal zones (>Fig. 2.6.1). These zones reflect physiographic patterns linked mainly to bedrock geology and location, from crystalline Pre-Cambrian (Archean and Proterozoic) rocks in Labrador, Ungava Bay, eastern Baffin and Ellesmere Islands, Palaeozoic sedimentary rocks in the central Arctic Archipelago and southern Hudson Bay to poorly consolidated Mesozoic and Cenozoic sediment along the north-westerly edge of the Arctic islands and the Beaufort Sea (Bird 1967). The geomorphic diversity of these northern coastal environments includes deltas, estuaries, lagoons, barrier islands, spits, fringing sand and gravel beaches, steep bedrock cliffs, and several systems unique to northern environments such as fiords, ice coasts (tidewater glaciers and ice shelves), ice-rich permafrost coasts, breached thawlake coasts, and thermokarst coasts. The complexity of these systems is increased by regional patterns of isostatic adjustment and wave climate. Sea level rise and littoral processes related to changing storm patterns and changing sea-ice conditions add to the diversity. The dynamic nature of any coastal system is a function of the physical and morphologic characteristics of the coastline materials as well as the environmental forces operating on it. There are dramatic contrasts in tide range. According to Shaw et al. (1998), the tide range decreases northward

and westward across the Arctic Archipelago. The largest range occurs in Frobisher Bay and Ungava Bay where it is about 13 m. A tide range of more than 4 m occurs in both southern (Hudson Strait) and northern Baffin Island (Nares Strait), while tides are less than 1 m in east central Baffin Island. Tides are more than 4 m in the Foxe Basin, 1–2 m in Barrow Strait and typically less than 0.5 m along the western and northern edge of the Arctic Archipelago and the Beaufort Sea. In Hudson Bay, the tide range is less than 4 m. In some areas, for example, the Beaufort Sea, storm surges can add 1–2 m of water-level rise to the tides. Relative sea level change is determined from either tide gauge records or calculated from coastal emergence curves. In the Canadian Arctic, there are only two locations where tidal records are long enough to detect current trends. At Tuktoyaktuk, tide records indicate a relative rate of sea level rise (submergence) of about 4 mm/year for the Beaufort Sea area. By comparison, tide gauge records from Churchill indicate an emergence rate of the order of 8–9 mm/year for western Hudson Bay. Emergence curves for the rest of the Canadian Arctic suggest emergence rates of up to 7 mm/year across most of the central Arctic including Hudson Bay and submergence of the order of 5 mm/year for eastern Baffin Island. The geomorphic evidence for emergence includes raised beach and delta deposits, marine sediment, marine shells, fossils and driftwood. Under current scenarios of climate change, many Arctic coastlines and several communities will be susceptible to flooding.

2.  The Coastlines Canada’s northern coasts are divided into regions based on common physiographic character and geology. The Labrador Coast (Zone 1) is divided into southern and northern segments. The southern segment from Battle Harbour to Nain consists of roughly 600 km of rocky coastline dominated by numerous small bays and inlets, rugged headlands and islands. The hard crystalline nature of the Pre-Cambrian rocks south of Lake Melville maintain low bluffs and cliffs and a steep nearshore

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

2.6

Northern Canada

⊡⊡ Fig. 2.6.1 Map of northern Canada showing coastal regions based on common geology and physiography.

7 8 11 8 YUK ON

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NWT

9

10

11

Coastal Regions of Northern Canada

4

5 10 NUNAVUT

1 2 3 4 5 6 7 8 9 10

4

6

COASTAL SUBDIVISIONS Labrador Hudson Strait-Eastern Baffin Is. Hudson Bay and James Bay East Baffin Bay to Lincoln Sea Jones, Lancaster and Prince Regent Sounds West Ellesmere and Axel Heiberg Islands Northern Ellesmere Island Northwest Arctic Archipelago Central Arctic Archipelago Southern Arctic Archipelago and Mainland coast Beaufort Sea

bathymetry that limit the development of littoral deposits and landforms to a few sheltered bays. The Groswater– Lake Melville area contains unconsolidated materials that form extensive beaches, beach ridges and coastal dunes. Small estuaries occur in some of the sheltered bays. Cape Harrison to Nain displays a mix of rugged and steeply sloping rocky coasts with small bays and headlands. In the Nain area, low relief and gentle nearshore slopes develop wide boulder-strewn tidal flats (>Fig. 2.6.2), bolder barricades and shingle beaches.

4 10

2 1

3

NUNAVIK

LABRADOR QUEBEC

0

Kilometres

1000

(Source: E. Owens 1994)

The north segment from Nain to Cape Chidley is roughly 450 km of rugged coastline dominated by the Torngat Mountains that rise abruptly to elevations over several hundred metres. The hard igneous rocks of the Canadian Shield (Nain Province) have been glacially eroded producing a coastline dominated by numerous large bays, islands and several deep fiords. For example, Nackvak and Saglek Fiords extend 20–40 km into the Torngat Mountains. In general, this section of the coast is marked by a paucity of beaches and tidal flats except for

Northern Canada

2.6

⊡⊡ Fig. 2.6.2 Boulder-strewn tidal flat along the Labrador coast. (Courtesy P. McLaren, Geological Survey of Canada.)

several small deltas that form at the head of fiords and in larger bays. Hudson Strait–Southeastern Baffin Island (Zone 2) is characterised by moderate-to-high relief coastlines underlain by hard Proterozoic crystalline rocks. Ungava and Frobisher Bays are known for their large tidal range and wide intertidal mudflats with boulders. The south side of Hudson Strait has a steep cliff coastline with a few bays and fiords. The north side is irregular, rugged and rocky with several large bays (e.g. Frobisher Bay) and fiords. The coast of Ungava Bay is low and rocky on both the east and west sides, while the south coast is low with wide intertidal mudflats and estuaries (e.g. George River). Small bays on the west side of Ungava Bay also display wide tidal flats. The Hudson Bay–James Bay area (Zone 3) fronts a large (1.23 million square kilometres), shallow inland sea that is covered with ice for 8–10 months of the year. The east coast of Hudson Bay is characterised by gently dipping (seaward) sedimentary strata forming a series of low cuesta islands parallel to the shore and farther offshore (e.g. Belcher Islands). The coast is gently sloping with low beaches and narrow tidal flats. The east coast of James Bay is characterised by Archean crystalline rocks mantled by till, glacial fluvial and marine (Tyrell Sea) deposits. The irregular coastline consists of numerous low rock outcrops, sand and gravel beaches, tidal flats and salt marshes. The coast on the west side of James and Hudson Bays is dominated by several estuaries with wide salt marshes, separated by coastal segments with tidal flats with sand

and gravel ridges oriented normal to the shore and ­shore-parallel sand-gravel beach ridges fronted by tidal flats and salt marshes. The extensive estuarine coastal systems reflect the importance of the Hudson Bay as the destination for many drainage systems in central mainland Canada. The northern part of the west side of Hudson Bay, including the south side of Southampton and Coates Islands, is characterised by low sandy beaches and wide tidal flats. The East Baffin Bay to Lincoln Sea (Zone 4) includes the western coast of Baffin, Devon and Ellesmere Islands. The coasts in this zone are underlain mainly by hard Pre-Cambrian crystalline rocks except for the Eclipse Sound area of northeastern Baffin Island where there are Cretaceous and Tertiary shales and sandstones, and parts of eastern Devon and northern Ellesmere Islands where there are pockets of Cambrian–Ordovician sedimentary rocks (Shaw et al. 1998). Unconsolidated Quaternary sediment are also locally important. All three islands are mountainous and heavily glacierized with active glacial systems that impinge on this coast, including ice caps, outlet glaciers and tidewater glaciers. The coastal relief is generally high and incised by long, winding steep-sided and talus-banked fiords (>Fig. 2.6.3) up to 98 km long. This zone has Canada’s greatest concentration of fiords (Syvitski et al. 1987). Unconsolidated shores are restricted to low coastal forelands and fiord heads where there are extensive sandur and deltaic deposits. In these areas, coastal landforms include low coastal bluffs, broad tidal flats with boulder barricades, fringing and barrier beaches, and

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Northern Canada

⊡⊡ Fig. 2.6.3 Long and twisting fiord from northeastern Baffin Island. (Courtesy R. Taylor, Geological Survey of Canada.)

raised beaches and deltas. The Cape Ashton area has an extensive beach–barrier island complex with aeolian dunes (Taylor et al. 1987). Bylot Island displays a range of barrier beach and nearshore bar complexes. Ice coasts comprise less than 1% of the coastlines and are most common along parts of Devon and southeastern Ellesmere Islands. The Jones, Lancaster and Prince Regent Sound (Zone 5) coastal zone includes northwestern Baffin Island, Devon Island, southern Ellesmere Island, southeastern Cornwallis Island and eastern Somerset Island. This region is underlain by relatively weak, flat-lying Palaeozoic limestone and dolomite. It is described by Shaw et al. (1998) as ‘a plateau region, deeply dissected by fiords. Scree-banked coastal cliffs 100–600 m high dominate the coastline except where the coastal plain is sufficiently wide to allow the formation of depositional beach features.’ Sea-ice conditions are considered severe and result in ice scour and ice push in the littoral zone. During the relatively short open water period (<2 months), beaches and barrier complexes in this area can be dramatically reworked by storm waves. Two coastal patterns repeat throughout the area, including: (a) steep, scree-banked cliffs that in some locations are fronted by gravel beaches, and (b) low gravel/cobble beach–coastal plain complexes sometimes backed by tundra ponds or shallow lagoons. In places, gravel beach ridges are common and where coupled with a history of emergence has led to well-developed strandlines and raised beaches typical of emergent coastlines in many parts of the

Arctic (>Fig. 2.6.4). Locally, bay and fiord head deposits include typical delta, low beach and intertidal flats. The West Ellesmere and Axel Heiberg Islands (Zone 6) zone lies on the eastern edge of the Sverdrup Basin, and is characterised by poorly consolidated folded Mesozoic and Cenozoic sedimentary rocks intruded by more resistant volcanics. Eureka Sound, the narrow passage that separates the two islands, is bordered by extensive lowlands composed of poorly consolidated Cenozoic sediment mantled by a veneer of marine sediment. Several long fiords (Canon, Tanquary, Gibbs, Mokka, Slidre and Greely Fiords) and bays feed into Eureka Sound. The coastline is made up of several straight segments characterised by: (a) gentle slopes with narrow sand and gravel beaches, (b) low-to-moderate bedrock bluffs fronted by narrow beaches, and (c) bays and fiords characterised by welldeveloped sand and gravel beaches, deltas, glacial flood plains, and in places, raised beaches and deltas. A few tidewater glaciers provide a continuous supply of small icebergs that become stranded in shallow bays. The south and west sides of Axel Heiberg Island are dominated by several large steep-sided fiords (>Fig. 2.6.5) (Li, Middle, Strand, Expedition, Glacier, Wolf and Skaare Fiords) and bays (Good Friday, Sand and Whitsunday Bays). Run-off from two large ice caps (Müller and Stacie Ice Caps), several smaller highland ice caps and valley glaciers (ice covers roughly 31.5% of the island) contribute to extremely high sediment loads that form large sand and gravel flood plains at the head of most fiords. Fiord sides are generally

Northern Canada

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⊡⊡ Fig. 2.6.4 Gravel beach backed by raised beach ridges, Somerset Island.

⊡⊡ Fig. 2.6.5 Fiord head with a fluvial delta and glacial flood plain, west Axel Heiberg Island.

steep with narrow beaches, or vertical with no beach. Low coastal plains with sand and pebble beaches, raised beaches and muddy lagoons occur in areas separating fiords on the west side of the island. The presence of sea ice year round limits littoral processes but contributes to localised ice push.

Northern Ellesmere Island (Zone 7) includes some of the most northerly coasts in the northern hemisphere and is considered an ice coast. Northern Ellesmere Island is underlain by Palaeozoic and Proterozoic rock covered by a veneer of Quaternary sediment and ice. The Grant Land Mountains is a high, ice-capped fold sequence responsible

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Northern Canada

for the high relief (up to 1,200 m) coastal landscape dissected by ice-choked fiords, epishelf lakes and ice shelves. The absence of open water results in a paucity of littoral processes and landforms. Over the past century, the once near-continuous ice shelf that rimmed most of northern Ellesmere Island has disintegrated into a series of smaller ice shelves (Jeffries 1987), which in recent years have almost disappeared (e.g. Ward Hunt and Ayles ice shelves). Northwest Canadian Arctic Archipelago Coast (Zone 8) consists of a series of low-relief (<60 m) islands situated on the northwestern perimeter of the Arctic Archipelago. It is underlain by poorly consolidated Mesozoic and Cenozoic sedimentary rock and includes all or part of Ellef Ringnes, Amund Ringnes, Cornwall, King Christian, Lougheed, Prince Patrick, Eglington, Borden, Mackenzie King and western Banks Islands. These low coastlines are ice-locked for more than 11 months of the year, and are characterised by a combination of low-energy littoral and sea-ice processes. The western perimeter is undergoing submergence as indicated by flooded tundra polygons, drowned river mouths and barrier beaches, and a lack of raised coastlines. However, raised beach deposits found along coasts of the inner islands display more typical emergence histories. In general, coastal landforms are dominated by low-energy littoral, fluvial and sea-ice processes and include coalescing deltas, lagoons, sand and mudflats, low-barrier beaches, spits and ice-push structures. Cornwall Island is distinguished by low bedrock bluffs and gravel barrier beach structures. Ellef Ringnes

and Prince Patrick Island coasts also have coastal landforms caused by mass wasting. Central Canadian Arctic Archipelago Coast (Zone 9) includes Bathurst, Melville and Byam Martin Islands, and northwestern Cornwallis and Devon Islands (as well as smaller islands contained within the Innuitian fold belt). These islands comprise a group of low-relief inner islands in the central Arctic Archipelago. The underlying geology is characterised by easily eroded Palaeozoic to Mesozoic age sandstones, siltstones and shales. Coastal morphology is influenced by a series of east–west trending folds where resistant lithologies form headlands and peninsulas and the weaker materials bays; this pattern is strongest on western Bathurst and eastern Melville Islands. Gravel beaches commonly develop on shale, limestone and dolomite lithologies, while sand beaches typically develop on sandstone (Shaw et al. 1998). Coastal morphologies combine patterns and landforms seen in adjacent zones including fiords and bays on western Melville Island and low-barrier beaches and sand flats on eastern Melville and Bathurst Islands. Areas with coarse sediment are characterised by barrier beaches, spits and beach ridge complexes separated by flooded tundra ponds. The Sabine Peninsula coastline is marked by numerous deltas and mudflats. Southern Arctic Archipelago and Mainland Coast (Zone 10) includes the eastern half of Banks Island, Victoria Island, Prince of Wales Island, Western Somerset Island, northern Southampton Island, southwestern Baffin Island and the mainland coast between the Beaufort Sea and

⊡⊡ Fig. 2.6.6 Low permafrost bluff formed in icy unconsolidated sediment, southern Beaufort Sea coast.

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⊡⊡ Fig. 2.6.7 Barrier island–tidal lagoon complex, southern Beaufort Sea coast.

⊡⊡ Fig. 2.6.8 Flooded tundra indicative of coastal submergence, Beaufort Sea coast.

Hudson Bay. Underlain by both resistant Pre-Cambrian crystalline rocks and less-resistant Palaeozoic to Cenozoic rocks, it is the largest of the coastal zones identified by Owens (1994). As expected for such a vast area, the coastline in this zone is extremely variable and according to Shaw et al. (1998) includes nearly every coastal type. The Beaufort Sea Coast (Zone 11) extends from Cape Bathurst to Demarcation Point, and includes the coast of the west side of Banks Island as well. It is composed of icerich glacial, fluvial glacial and marine sediment that form the Yukon Coastal Plain, the Mackenzie River Delta, a series of small Tundra Islands (e.g. Richards, Gary, Pelly and

Hooper Islands), the Tuktoyaktuk Peninsula and western Banks Island. Coastal cliffs, which represent 52% of the coastline, are fronted by narrow sand and gravel beaches (>Fig. 2.6.6). They are mostly low-lying, generally <10 m in height, although in places are up to 90 m high. Barrier island and spit formations (>Fig. 2.6.7) are also common, especially along the Yukon Coastal Plain and the Tuktoyak­tuk Peninsula. Wetlands, tidal flats and delta deposits characterise the mouth of the Mackenzie River and a number of other smaller rivers. Other landforms include submerged tundra (>Fig. 2.6.8) and breached thermokarst lakes, which are widespread along the Tuktoyaktuk

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⊡⊡ Fig. 2.6.9 Erosion of ice-rich unconsolidated sediment, Beaufort Sea coast.

Peninsula. This coast is erosional in nature with an average annual retreat of about 1 m/year and a maximum annual retreat of 22.5 m/year (Lantuit and Pollard 2008). Areas of exposed massive ground ice develop retrogressive thaw slumps and thermokarst exposures (>Fig. 2.6.9). Dallimore et al. (1996) have shown that degradation of sub-sea permafrost in the nearshore also contribute to shore face as well as cliff retreat. This coast has been identified as highly vulnerable to the effects of sea level rise and global warming (Shaw et al. 1998). Storm waves and storm surges play an important role in coastal dynamics.

References Bird B (1967) The physiography of Arctic Canada. John Hopkins Press, Baltimore, MD, 336 pp

Dallimore SR, Wolfe SA, Solomon SM (1996) Influence of ground ice and permafrost on coastal evolution, Richards Island, Beaufort Sea Coast, N.W.T. Can J Earth Sci 33:664–675 Jeffries M (1987) The growth, structure and disintegration of Arctic ice shelves. Polar Record 23:631–649 Lantuit H, Pollard WH (2008) Fifty years of coastal erosion and retrogressive thaw slump activity on Herschel Island, southern Beaufort Sea, Yukon Territory, Canada. Geomorphology 95:84–102 Owens E (1994) Canadian Coastal Environments, Coastline Processes, and Oil Spill Cleanup Environmental Protection Series Report 3/SP/5 Environment Canada, Ottawa, 328 pp Shaw J, Taylor R, Forbes M, Solomon D (1998) Sensitivity of the coasts of Canada to sea level change. Geol Survey Canada Bull 505:79 Syvitski J, Burrel D, Skei J (1987) Fjords: processes and products. Springer Verlag, New York, 379 pp Taylor R, Praeg D, Syvitski J (1987) Coastal morphology and sedimentation, Eastern Baffin and Bylot Islands, N.W.T. In: Sedimentology. Experiment Data Report, Vol 3. Geological Survey of Canada Open File 1589

2.7 Great Lakes (Canada)

Mary-Louise Byrne

1.  Introduction The Great Lakes are five interconnected freshwater lakes in central eastern North America. Closest to the continental interior is Lake Superior, which flows through the St. Mary’s River into Lakes Michigan and Huron. These lakes empty into the connecting channels of the St. Clair River, Lake St. Clair and the Detroit River that in turn flow into the southernmost of the Great Lakes, Lake Erie, which empties into the Niagara River and Welland Canal, and then Lake Ontario. Finally, Lake Ontario empties into the St. Lawrence River. Overall, the Great Lakes span more than 1,200 km from the head of Lake Superior to the St. Lawrence River. According to the Great Lakes Atlas (2000), the lakes contain about 23,000 km3 of water covering 244,000 km2 of area. They have a total coastline length of more than 17,000 km in both Canada and the United States. The drainage area that empties into the lakes covers more than 521,000 km2. The climate of the Great Lakes basin is greatly influenced by its location in the continental interior (Carter and Haras 1985). Weather patterns are complex and change rapidly because of the location of the basin at the boundary of the polar and tropical air masses. The lakes modify temperature and humidity by acting as heat sources in winter and heat sinks in summer. Precip­­ itation  ranges considerably in this area. For example, mean annual precipitation varies from a total of 74.6 cm (55.9 cm rainfall and 18.7 cm snowfall) in Thunder Bay in  the northeast, to 89.3 cm (80.4 cm in rainfall and 8.9 cm in snowfall) in Kingsville in the south, to 100.1 cm (79.4 cm in rainfall and 20.7 cm in snowfall) in Cornwall in the southeast. The prevailing westerly winds create surface waves and currents in the Great Lakes, and can cause short-term fluctuations in lake levels during intense storms. The influence of these variations is greatest in Lake Erie, the shallowest of the lakes. There are also seasonal and longer term fluctuations in lake levels. Long-term fluctuations are influenced by variations in the balance of precipitation and evaporation. Variations between record high levels

and record low levels range from 1 to 2 m. The greatest variations are on the middle lakes (Erie, Huron and Michigan) and lowest on Lake Superior (the largest) and Lake Ontario (the most controlled). The coastlines of the Canadian side of the lakes ex­­­ hibit great variability from Lake Superior through to the St. Lawrence. The Upper Great Lakes consist of Lake Superior, Lake Michigan and Lake Huron. Lakes Superior and Huron have coastlines in Canada, but Lake Michigan lies wholly within the United States and will therefore not be included here. Geologically, the basins of these lakes appear to have their origins in the Pleistocene glaciations of the last 2.5 million years (Eyles 2002). Our knowledge of their geology prior to that time is very limited. The ice sheets that advanced over Canada and the northern United States deeply eroded river valleys producing the basins that the present lakes fill. During glacial retreat, the lakes had several configurations, approximating the current five lakes and connecting channels by about 8,000–10,000 years bp. The exposed shield rock in the northern portions of the Upper Lakes shows evidence of erosion during glaciation, while the deep Pleistocene deposits of the Lower Great Lakes dominate the present coastlines.

2.  Lake Superior Lake Superior, the most northerly and least heavily settled of the Great Lakes, lies at 183 m above sea level and is shaped by the mid-continent rift system. Pre-Cambrian rock of the Canadian Shield is exposed on its northern coastline (Eyles 2002). Total coastline length in Canada is 1,394 km with an additional 990 km on islands. The coast is dominated by rock outcrops with a thin cover of glacial deposits, with glacial lake sediment close to the shore, and unstratified ground and end moraine further inland. The present coastline consists of high cliffs with coarsegrained pocket beaches, punctuated by deep bays like that of Thunder Bay and Nipigon Bay (Carter and Haras 1985).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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3.  Lake Huron Lake Huron (176 m above sea level) has the longest coastline on the Canadian side, more than 2,044 km on the mainland and 2,768 km on the many islands that dot the lake and its large bays. Lake Huron is separated from Lake Michigan by the wide Straits of Mackinaw, and the two lakes are considered hydrologically one. Water flows from the Straits to the main basin of Lake Huron. To the north is Manitoulin Island, separated from the mainland by the North Channel. Manitoulin Island is the world’s largest freshwater island, and is composed of low-sloping limestone coast on the main lake side and steeper cliffs on the North channel side. The north coast of North Channel and the east coast of Georgian Bay are composed of Pre-Cambrian rock, with rare pocket beaches of sand as in Killbear Provincial Park. The eastern side of the Bay is known as the 30,000 Islands. Further south in Georgian Bay, the eastern coast becomes more gently sloping, sand is more abundant, and features like Wasaga  Beach in Nottawasaga Bay have accumulated. This area marks the transition from the Pre-Cambrian basement to the overlying Ordovician, and sub­ sequently  the Devonian and Silurian Rocks of the Niagara Escarpment that outcrop on the western side of

Georgian Bay. Spectacular limestone and dolostone cliffs plunge into the waters of the Bay, the water being more than 160-m deep immediately off the northern tip of the Bruce Peninsula (>Fig. 2.7.1). The cliffs have been dissected along faults and joints (>Fig. 2.7.2). The Lake Huron side of the Bruce Peninsula is characterised by the same gently sloping limestone seen in Manitoulin Island. In the north, the bedrock is often exposed (>Fig. 2.7.3). Moving south along the Lake Huron coast, steep bluffs of Pleistocene till are common. The easily eroded till supplies a wealth of sand that has accumulated in the southern end of a littoral cell at Grand Bend and Pinery Provincial Park, dominated by wide low-­sloping sand beaches and dunes (Dech et al. 2005) (>Fig. 2.7.4). Ice plays a role in the transport of sediment (>Fig. 2.7.5), and waves and currents generate sediment movement in the ice-free season (Houser and Greenwood 2005).

4.  Lake Erie Moving through the connecting channels of the St. Clair River, Lake St. Clair and the Detroit River, the transition is made from the Upper to the Lower Great Lakes. Lake Erie (173 m above sea level) has a coastline length in Canada

⊡⊡ Fig. 2.7.1 Cliffs and shore platforms with cobble and gravel beaches on the east side of the Bruce Peninsula, Lake Huron. Cobbles are derived from the cliffs.

Great Lakes (Canada)

2.7

⊡⊡ Fig. 2.7.2 The Grotto, a cave eroded along a plane of weakness in the limestone of the Niagara escarpment in Bruce Peninsula National Park, Lake Huron.

⊡⊡ Fig. 2.7.3 Singing Sands beach on the Bruce Peninsula, Lake Huron. A thin veneer of sand overlies limestone.

of 592 km on the mainland and 47 km on islands. Lake Erie, the smallest in volume and shallowest of the Great Lakes, has a coastline dominated by outcrops of glacial till. The lake can be divided into three distinct basins. The western basin is the shallowest and the north coast is composed of 3–20-m high bluffs of glacial drift that decline to sediment-starved beaches

(>Fig. 2.7.6). This basin extends from the Detroit River to Point Pelee, the southernmost part of mainland Canada. Recently, under relatively low lake levels, erosion has been a problem in this part of the lake where the lakebed is composed of cohesive clay, and new models are needed to predict the outcomes (Trenhaile 2008). The central

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⊡⊡ Fig. 2.7.4 Vegetated sand dunes at Pinery Provincial Park, Lake Huron, formed as lake level fell during the past 10,000 years. The sand deposits extend at least a kilometre inland.

⊡⊡ Fig. 2.7.5 The beach at Pinery Provincial Park, Canadian coast Lake Huron, in winter. Often the beach is frozen, with several ice ridges developing offshore. These contain an abundance of sand that melts out in places in spring, and is added to the beach.

basin, extending from Point Pelee to Longpoint has a north coast composed of bluffs of glacial till up to 40-m high (>Fig. 2.7.7). The bluffs are easily eroded and the nearshore waters are often heavily loaded with sediment (dominantly silt and clay). On the eastern end of this

stretch, the bluffs are capped by the sand of the Norfolk sand plain, and cliff-top sand dunes are common. The eastern basin extends from Longpoint eastward to the Niagara River. The north coast of this basin is composed of limestone outcrops and up to 15-m high glacial bluffs.

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⊡⊡ Fig. 2.7.6 Beach sand deficit on the NW coast of Lake Erie has prompted coast protection works. Block armouring has been introduced just east of Point Pelee, with a groyne that has failed to intercept longshore drifting sediment.

⊡⊡ Fig. 2.7.7 Bluffs on the coast of Lake Erie, cut in soft glacial till and capped by thin deposits of sand, forming cliff-top dunes.

Pocket beaches are common between the limestone headlands. Lake Erie empties into the Niagara River and Welland Canal, which open into Lake Ontario.

5.  Lake Ontario Lake Ontario (74 m above sea level) is the smallest lake in surface area. However, its volume is large due to its great

depth, the average being over 86-m deep and the deepest part more than 244 m below low water datum. Depths increase rapidly offshore in the west, while the east sees a more gradual deepening. The Canadian coastline of Lake Ontario is more than 537 km long, with an additional 80 km on islands. The coastline of the western end of the lake is comprised of bluffs of glacial till and outwash. These are up to 90-m high in the Scarborough bluffs in the east of Metro Toronto (Eyles 2002). The bluffs decrease in size to

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⊡⊡ Fig. 2.7.8 Gently sloping beach and small dunes at Presquile Provincial Park, Lake Ontario. Vegetation has been planted to trap wind-blown sand and prevent it blowing inland.

the east where marshes and barrier dune complexes have developed. Presquile and Sandbanks Provincial Parks are typical of barrier dune and marsh complexes (>Fig. 2.7.8). Prince Edward County is a limestone plateau at the eastern end of the lake with an irregular coast containing bays that are partly blocked by spits or barrier islands.

References Carter CH, Haras WS (1985) Great lakes. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 253–260

Dech JP, Maun MA, Pazner MI (2005) Blowout dynamics on Lake Huron sand dunes: analysis of digital multispectral data from colour air photos. Catena 60:165–180 Eyles N (2002) Ontario rocks. Fitzhenry and Whiteside, Markham, Ontario, p 399 Great Lakes Atlas (2000) The Great Lakes: An environmental Atlas and Resource. Book http://www.on.ec.gc.ca/great-lakes-atlas/intro.html Hough JL (1958) Geology of the Great Lakes. University of Illinois Press, Urbana, IL Houser C, Greenwood B (2005) Profile response of a lacustrine multiple barred nearshore to a sequence of storm events. Geomorphology 69:118–137 Trenhaile AS (2008) Modeling the erosion of cohesive clay coasts. Coastal Eng 56:59–72

3.0 Pacific Central America – Editorial Introduction

Central America is here taken to extend from Mexico south to Colombia. It is dominated by mountain ranges that continue the trend of the Rocky Mountains of North America southward through an isthmus to the Andes Mountains of South America, but there are intricate structures resulting from a complex tectonic history, and some associated vulcanicity. Rock formations range from Palae­ ozoic to Holocene (Weyl 1980; Yanez-Arancibia 2005).

1.  Pacific Coast On the Pacific coast narrow coastal plains alternate with steep sectors. Earthquakes occur, and there are many uplifted beaches and terraces on the coastal slopes. Ocean swell from the SW breaks on beaches supplied with sand mainly from rivers draining the steep slopes of coastal ranges. Beaches pass laterally into barriers in front of coastal lagoons and swamps. In Costa Rica and Panama the coastline becomes irregular in response to geological structures, with long sectors of steep and rocky coast, and shallow embayments, notably the Gulf of Panama, in which wave action is attenuated. Mean spring tide range is generally small (<2 m), but increases in these embayments: Balboa at the head of the Gulf of Panama has a 5 m range. There is a transition from an arid climate in northern Mexico to humid tropical conditions in Nicaragua, Costa Rica, Panama and Colombia. Mangroves occupy the shores of estuaries and deltas, notably in the Gulf of Fonseca, while coral reefs are found away from these turbid sea areas, particularly on islands offshore. Cliffs border the coastal ranges extending from southern Panama into

northern Colombia, but in southern Colombia a wide coastal plain is fringed by mangroves in numerous inlets and estuaries.

2.  Caribbean Coast The Caribbean coast of Central America generally has wider coastal plains – notably in northern Colombia – and a wide continental shelf. Wave energy is generally low (except during occasional hurricanes) and tide ranges are very small (mostly <0.5 m). Much of the coastline has narrow sandy beaches and barriers enclosing lagoons, and mangroves are extensive, particularly in inlets and estuaries. Coral reefs are found away from river mouths and around offshore islands. On the Caribbean coast of Mexico there are limestone cliffs bordering the plateau on the Yucatan Peninsula, and beaches and barriers with lagoons extend around the Bahia de Campeche, backed by a wide coastal plain. As the climate becomes drier northward the lagoons become more brackish, and saline flats are associated with the Mexican Laguna Madre towards the Rio Grande delta.

References Weyl R (1980) Geology of Central America. Gebruder Borntraeger, Berlin Yanez-Arancibia A (2005) Middle America, coastal ecology and geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 639–645

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_3.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

3.1 Mexico, Pacific Coast

Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez

1.  Introduction The Pacific coastline of Mexico is about 6,500 km long, including the coasts of Baja California. The northern part of Mexico has a warm temperate climate, becoming arid in Baja California and Sonora under the influence of NE trade winds. To the south the climate becomes tropical, with high rainfall (about 2,000 mm/year) in the coastal fringe. Ocean swell comes in to the Pacific coast, but the NE trade winds blow offshore. There are occasional hurricanes. This coast is subject to seasonal changes in the direction of wave approach and resulting longshore drifting. In winter (October–March) NW waves generate southeastward drifting, while in summer (April–September) SW waves generate northward drifting. Tide ranges are larger (0.8–1.6 m) during the solstices, and reduced (0.2–0.4 m) in the equinoxes. Tide ranges increase into the Gulf of California – Bahia San Quentin has 1.4 m, Puerto San Bartolomé 2.5 m, the Gulf of Cal­ifornia La Paz has 1.3 m and Santa Rosalia 1.4 m. To the south Mazatlan has a tide range of the order of 1.4 m and Acalpuco has 0.9 m.

2.  The Coastline The Pacific coast of Mexico is backed by mountain ranges formed by tectonic uplift, folding and faulting, with some volcanic activity. Baja California continues the Coast Ranges of southern California. The coast from the Colorado River delta to Puerto Vallarta has been classified a neo-trailing-edge coast and from Puerto Vallarta to Guatemala, it has been classified a continental collision coast (Inman and Nordstrom 1971; Carranza-Edwards et al. 1975). On Baja California the morphological and tectonic units are the Sierra de Juarez (Mesozoic) and Sierra La Giganta (Tertiary), the basins of Victoria and Vizcaino (Cretaceous), and Purisima Iray (Tertiary). The west cost is rugged, but there is lowland on the eastern side of Magdalen Bay and behind Vizcaino Bay. Vegetation is too sparse to retain sand blown inland from the beach and

there are bare salt flats with barchanoid dunes. Guerro Negra is a coastal lagoon. East of the Gulf of California, the coast south to Puerto Vallarta is the Sierra Madre Occidental (MesozoicTertiary), a large cordillera that extends from Sonora to the Eje Neovolcanico. The Rio Grande de Santiago is located within this tectonic unit and drains the tectonic lake of Chapala. South central Mexico includes the Sierra Madre del Sur (Late Mesozoic to Tertiary), the Sierra de Oaxaca (Mesozoic) and the Eje Neovolcdnico (Late Tertiary). The Balsas River has cut a valley across the Sierra Madre del Sur, which extends parallel to the coast from Puerto Vallarta to Salina Cruz. To the south, the main tectonic units are the Isthmian metamorphic zone (Palaeozoic-Mesozoic) and the Macizo de Chiapas (Palaeozoic-Mesozoic). The continental shelf bordering the Sierra Madre del Sur is very narrow, The width varies from about 50 km in northern Sonora to 5 km just south of the MichoacanColima boundary and in western Guerrero (Shepard 1973). At least one major submarine canyon cuts the shelf (Reimnitz and Gutierrez-Estrada 1970). The coastal mountain flank is wide in Michoacan and Oaxaca, where it averages 70 and 90 km respectively. It can be divided physiographically into basins of coastal rivers. In addition, the mean elevation of the mountains is lowest in the northwest, increasing to 2,200–2,700 m in the southeast. Some of the larger coastal basins are those of the Tehuantepec, Ometepec, Coahuayana, Ameca, Baluarte, Huamaya and Concepcion rivers. The Pacific coast can be divided into rocky eroded sectors and sectors prograded by deposition (>Fig.  3.1.1). The former are characteristic of the Peninsula de California coast, where the relatively high relief has caused the development of steep cliffs and coastal canyons; discontinuous beaches of sand with a bouldery substrate are present. Similar features are seen between Puerto Vallarta and Chiapas. Where the coastal landforms consist of ridges and valleys the coast is irregular, with capes and bays, pocket beaches and esteros (Zihuatanejo). Valleys and canyons descend to the coast of Sinaloa, where large deltas have developed (Gutierrez-Estrada 1972). Where the coast

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_3.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Mexico, Pacific Coast

⊡⊡ Fig. 3.1.1 Predominant coastal landforms of the Mexican Pacific coast. (Courtesy Geostudies.) N A Y A R I T

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is relatively low the shore is variable in width, and beaches fringe deltaic coastal plains (Tehuantepec, Nayarit, Sinaloa, and Sonora). Coastal lagoons are common, separated from the sea by sandy barriers with beach ridges and dunes (Yavaros lagoon) (Ayala-Castanares et al. 1980). Coasts with active river sedimentation are arcuate in shape and the shore is wide with many transverse channels and extensive mangrove swamps, as at the mouth of Balsas River (Gutierrez-Estrada 1971).

References Ayala-Castanares A, Maplica-Cruz VM, Gutierrez-Estrada M (1980) Marine geology of the Yavaros Region, Sonora, Mexico. An Centro

CLIFFS & BLUFFS

SANDY

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SALT MARSHES & SWAMPS

Cienc del Mar y Limnologia Univ Nal Auton Mexico (in Spanish) 7(2):275–289 Carranza-Edwards A, Gutierrez-Estrada AM, Rodriguez-Torres R (1975) Morphotectonic units of the Mexican coast. An Centro Cienc del Mar y Limnologia Univ Nal Auton Mexico (in Spanish) 2(1):81–88 Gutierrez-Estrada M (1971) Physiography and sedimentology of the delta of the Rio Balsas, Michoacan, Mexico. Mexico Univ Nal Auton Inst Geol Bol (in Spanish) 93:1–58 Gutierrez-Estrada M (1972) Preliminary classification of the Mexican coast (in Spanish). Union Geofisica Mexicana, Reunion Anual, 1972, Programas y Resumenes Inman DL, Nordstrom CE (1971) On the tectonic and morphological classification of coasts. J Geol 79:1–21 Murray GE (1961) Geology of the Atlantic and Gulf coastal province of North America. Harper, New York Reimnitz E, Gutierrez-Estrada M (1970) Rapid changes in the head of the Rio Balsas submarine canyon. Mar Geol 8:245–258 Shepard FP (1973) Submarine geology. Harper & Row, New York

3.2 Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama Anja Scheffers · Tony Browne 

1.  Introduction The Pacific Ocean coastline of Central America extends along the western to southwestern shores of six countries: Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama (>Fig.  3.2.1). This is about 2,200 km overland. There is considerable geological and geomorphological diversity from one end to the other (>Fig.  3.2.2) (Yanez-Arancibia 2005). There are no known Pre-Cambrian rocks in Central America, the oldest formations being the Maya Series of  Palaeozoic metasediment in Belize (Moody 1975). The Guatemala Massif metamorphics are also of Palaeo­ zoic age. The Mesozoic is represented by Triassic and Jurassic metamorphic rocks in the Honduras Massif, and on the Santa Helena peninsulaby Triassic and Jurassic dark shales near Tegucigalpa and Todos, the Santos red  beds in Guatemala, and by a variety of Cretaceous sediment and volcanic formations. Cenozoic volcanic and sedimentary rocks are widespread, and include Palaeocene sediment in the Peten lowland, the thick Eocene clastic beds of the Toledo Formation in Guatemala and Eocene volcanics in Costa Rica; Oligocene and Miocene clastic sediment, limestones, and volcanic rocks in Costa Rica, Nicaragua and Panama, and marine Pliocene formations in Panama. The most recent volcanic rocks in the region are of Pleistocene and Holocene age (Weyl 1980). There are three geotectonic provinces in Central America: (1) the Peten lowland, a continuation of the Campeche-Yucatan basin, in Guatemala and Belize; (2) the fault block mountainous massifs in Belize, Guate­­ mala and Honduras; and (3) the volcanic terrain made up  of the Isthmian link – the Pacific coastal plain, the Nicaraguan volcanic upland and the volcanic ranges of the Pacific ­cordillera from Mexico to Panama. The generally accepted plate tectonic theory concerning Central America holds that the region has been thrust northeastward, as

part of the Caribbean-East Pacific plate, between the essentially westward moving North and South American plates. Tides on the Pacific coast of Central America are semi-diurnal. The tide range is 1.5–2 m at Guatemala, 2–3 m at El Salvador and Nicaragua, 4 m in the Gulf of Fonseca and 4.5–6 m in the Gulf of Panama (GierloffEmden 1982). The coast is exposed to Pacific ocean swell, with waves in the surf zone up to 2 m high (GierloffEmden 1982). The predominant longshore drifting is southeastward.

2.  Guatemala The Pacific coast of Guatemala is about 250 km long and consists essentially of sandy beaches, beach ridges and dunes, with no cliffs. The coastline is gently curving because sand supplied by longshore drift has sealed former inlets. Some are now lagoons, bordered by mangroves or bare hypersaline tidal flats. River mouths have been deflected behind beach ridges, and deltas are blunted by high wave energy over a narrow continental shelf. Beach and dune biota include the creeping vines Ipomea and Canavalia, erect plants like Uniola, Scaevola and Distichlis, and burrowing invertebrates such as lamellibranchs, polychaetes and crustaceans.

3.  El Savador The northwest coast of El Salvador is rocky and cliffy, with red and green algae in the sublittoral zone. Barnacles, chitons, crabs, mussels, periwinkles and limpets are found in the tide pools and on the rocks. Barrier islands dominate the central sector of the coast (>Fig. 3.2.3). The barriers are 1–2 m high and the lagoons behind them are 1–3 m deep, with tidal channels to a depth of about 10 m, and a

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Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

⊡⊡ Fig. 3.2.1 Location map of the Pacific coast of Central America, showing drainage patterns. (Courtesy Geostudies.)

dense mangrove thicket with meandering tidal channels. The southeast coast of El Salvador, outside the Gulf of Fonseca, has sandy beaches backed by dunes. The Gulf of Fonseca is flanked on the west­­ern  side by the high Conchagua Volcano (1,258 m), and has  a mangrovefringed northern coast (Gierloff-Emden 1959).

4.  Honduras The Pacific coast of Honduras is a 70 km long deltaic sector behind the Gulf of Fonseca. This Gulf, bordered by El Salvador, Honduras and Nicaragua, extends about 50 km inland and is 70 km wide, backed by a largely mangrove fringed deltaic coastline. The largest river entering the gulf is  the Goascoran, on the border between El Salvador and

Honduras. There are three volcanic islands, each 5–10 km in diameter, in the gulf: El Tigre, Conchaquita and Meanguera.

5.  Nicaragua Southeast of the Gulf of Fonseca, the coast of Nicaragua, more than 300 km long, includes the mountainous ­volcanic Cosiguina Peninsula. To the southeast extensive beaches and beach ridges are cut by small mangrove-fringed rivers. Longshore drifting is from south to north. At the southeastern end the Nicaraguan coast becomes steep with rocky shores and coves with pocket beaches separated by cliffs. On 2 September 1992 an earthquake SW of Managua caused a tsunami of up to 10.7 m along about 200 km of the coastline.

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

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⊡⊡ Fig. 3.2.2 Geomorphological map of the Pacific coast of Central America. (Courtesy Geostudies.)

6.  Costa Rica

7.  Panama

Steep and rocky coasts extend along most of Costa Rica. Plant and animal communities are like those of northern EI Salvador. Emerged terraces are prominent along some stretches of this steep coast. The peninsula of Capo Sta. Elena follows an east–west anticlinal trend, and has drowned valleys along its northern coast. The Nicoya peninsula has marine terraces deformed by tectonic movements, and has been subject to many tsunamis (Fernandez et al. 2000). The Gulf of Nicoya is a wide indentation bordered by sandy beaches, with a northwest extension to the mouth of a mangrove fringed estuary. Longshore drifting from the southeast and east has fed the long spit of Puntarenas, backed by a dense mangrove forest. The large open Bahia de Coronado shows blunt deltas and barred river mouths with beach ridges and mangroves in swales on coastal barriers. To the south east cliffs and bays with small beaches border the Peninsula de Osa and the Golfo Dulce. The gulfs occupy grabens between uplifted mountain chains.

Panama has the longest Pacific coastline (more than 900 km) of Central America. Indented rocky shores with steep bluffs border the coast of the Gulf of Chiriqui, with the high outlying volcanic island of Coiba rising to 400 m and the steeply-sided Peninsula of Azuero to the east. To the east the steep coasts are interrupted by bays with barrier islands, beach ridges, blunt deltas and mangroves. High rainfall feeds rivers draining steep hinterlands, which have delivered large quantities of sediment to the coast. The Gulf of Panama is about 180 km long from east to west and 140 km from north to south, and is rather shallow with depths of only 50–100 m in the central region. The western and eastern coasts are steep and rocky but intertidal flats are extensive along the coast west of the Panama Canal, and to the east are large areas of marsh and mangrove swamps. The indented coastline to the east includes submerged valley mouths bordering parallel mountain chains. The main coral reefs on the Pacific coast of Central America are located in Panama off the Azuera

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⊡⊡ Fig. 3.2.3 The coast of El Salvador between the Rio Lempa and the Rio Grande de San Miguel, showing beach ridge patterns on the outer barrier and a segment of an earlier inner barrier. Mangrove swamps are shown in green. (Courtesy H.G. Gierloff-Emden.)

190 Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

Peninsula and Coiba Island, in clear water away from river mouths and suspended sediment.

References Fernandez M, Molina E, Havskov J, Aatakan K (2000) Tsunamis and tsunami hazards in Central America. Nat Hazards 22:91–116 Gierloff-Emden HG (1959) Die Küste von El Salvador. Dtsch Hydrogr Z 12:14–24

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Gierloff-Emden HG (1982) Central America, coastal morphology. In: Schwartz ML (ed) The Encyclopedia of Beaches and Coastal Environments. Hutchinson and Ross, Stroudsburg, PA, pp 188–191 Moody JD (1975) Central America regional review. In: Fairbridge RW (ed) The Encyclopedia of World Regional Geology, part 2. Dowden, Hutchinson and Ross, Stroudsburg, PA, pp 228–237 Weyl R (1980) Geology of Central America. Bornträger, Berlin Yanez-Arancibia A (2005) Middle America, coastal ecology and geomorphology. In: Schwartz M (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 639–645

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3.3 Pacific Coast of Colombia

Iván Correa · Robert Morton

1.  Introduction The Pacific coast of Colombia (for its Caribbean Coast, > Caribbean Coast of Colombia) extends to the west of the Cordillera Occidental, the westernmost branch of the South American Andes. The region, known for its heavy rainfall, numerous rivers, and luxuriant vegetation, remains mostly uninhabited and undeveloped. No roads reach this coast, so access is only by boat. Overland access from the interior of the country is only possible to the cities of Bahía Solano, Buenaventura, Guapi and Tumaco, the main commerce centres of the region. Because the Nazca plate is being subducted under the South American plate, the Pacific Coast of Colombia is an active seismic zone with a well-documented record of high-magnitude earthquakes. Seismic observations for the past century record at least three major seismic events (1906, 1979 and 1992), accompanied by subsidence, soil liquefaction, and extensive landsliding that caused abundant river damming. The 1906 and 1979 events caused littoral co-seismic subsidence up to 1.5 m along the southern Pacific Coast between Buenaventura and the Mira delta. Both events triggered tsunami waves 2.5 m high that caused more than 3,500 deaths of inhabitants of the low coastal areas (Herd et al. 1981). Between Punta Ardita (at the Panamá border) and Cabo Corrientes the coastal zone corresponds to the western flanks of the Baudó Range (Serranía del Baudó), a 600  m high region composed predominantly of oceanic basalts, diabases, and associated cherts and radiolarites. The Baudó Range is a structurally controlled relief, densely fractured and faulted in the preferential directions NNESSE and N 60º E – N 30º W that control the general orientation of valleys and ridges. South of the Baudó Range, the relief of the Pacific Coast is dominated by 20–100 m high hills cut into Tertiary sedimentary sequences and Plio-Quaternary deposits at the piedmont of the Cordillera Occidental (Martínez et al. 1995). Conspicuous terraces between 20 and 30 m above present sea level are found along the hills between Cabo Corrientes and Buenaventura Bay.

Coastal relief south of the Baudó Range abruptly limits the Holocene prisms of accretionary deposits that characterise the Pacific Coast of Colombia. They include the three major Plio-Quaternary deltaic prisms of the San Juan, Patía, and Mira Rivers, and two narrow (<10 km wide) fringes of coalesced deltas between the Cabo Corrientes-Togoromá and Buenaventura-Guapi stretches of coast (Martinez et al. 1995). Holocene deposits along the narrow littoral fringe of the Pacific Coast form distinctive sets of contemporaneous barrier islands, extensive mangrove swamps and from the San Juan delta to the Ecuador border, ancient barrier systems cut by funnel-shaped estuaries interconnected by a dense network of tidal channels. Extensive, highly unstable, barren sandy and muddy tidal flats are common at all the main tidal inlets of the Pacific Coast. The climate of the Pacific Coast is humid tropical, dominated by the annual migration of the Intertropical Con­­vergence Zone (ITCZ) and by high relief associated with the Cordillera Occidental. Mean daily temperatures are about 26 ºC with minimum values of about 14 ºC. Annual rainfall reaches a maximum of about 10 m at the Buenaventura-San Juan River delta area, and decreases to the north and south to about 3 m/year at the Panamá border and at the Tumaco-Mira River delta area. El Niño events regularly cause noticeable anomalies in rainfall, ocean-surface temperature, and sea level along the Pacific coast of Colombia. The 1997–1998 event was the most intense of the twentieth century. Tides along the Pacific Coast are mixed semidiurnal, with mean amplitudes between 2 m at Tumaco and 4 m at Buenaventura. Maximum spring tide amplitudes increase slightly from south to north, and vary between 3.5 m at Tumaco to 4.5 m at Buenaventura and Juradó. ­Predomi­nant winds along the Pacific Coast come from the south at the latitudes of the Baudó Range, and from the west to southwest along the central and southern sectors of the Pacific coast. Swell waves along the Pacific Coast are about 0.5–1.5 m high during calm periods, but can be as high as 2.5–3.5 m during strong wind periods. High waves are focused on the extensive subtidal and intertidal bars that front the major tidal inlets.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_3.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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2.  Northern Pacific Coast The Northern Pacific coast extends for about 650 km between Punta Ardita and the bays of Málaga and Buenaventura. The 250 km long coastline of the Serranía del Baudó, between Punta Ardita and Cabo Corrientes is dominated by steep cliffs, up to 70 m high, cut into diabases and chert. Cliffs alternate with small sandy to shingle and cobble pocket beaches that are located within the minor coastal indentations. Wide sandy beaches and sandy-muddy tidal flats front the major coastal embayments and river mouths. Small mangrove swamps are located landward of the beach and along the low courses of the main rivers, where tidal waters penetrate 0.5–1 km upstream from the coastline. Recession of the Baudó cliff coast is negligible in most areas, but saprolite landslides and rock falls are common during heavy rain, and/or strong onshore winds and seismic events. The strong structural control on the coastal morphology of the Serranía del Baudó is exhibited well in the Bahía Solano-Bahía Utría zone, where the 8 km long Utría Bay depression closely coincides with the Bahía Solano fault zone. The fault is interpreted as a reverse fault with a slip rate of 0.2–1 mm/year. The Bahía Solano village was partly destroyed and suffered local subsidence of 20–30 cm as a result of the September 26 1970 earthquake, associated with the Bahía Solano fault. There have been at least three coastal subsidence events in the area during the last 800 years, probably associated with earthquakes of similar magnitude. The Cabo Corrientes-Togoromá coastline is about 100  km long and has a general north-south orientation. This depositional fringe varies in width from 1 km at the mouth of the Baudó River to 9 km at the Docampadó River. Coastal morphology is characterised by a chain of barrier islands, backed by estuarine lagoons, mangrove swamps, and a narrow belt of freshwater swamps. The barrier islands of this coast, some of them regressive, are no more than 2 m high, 3.2–15.8 km long, and have areas between 0.6 and 25.2 sq km (Martínez et al. 1995). Holocene deposits between Cabo Corrientes and Togoromá are limited to the east by paleo-cliffs cut into Tertiary sedimentary rocks. The paleo-cliffs mark the maximum advance of the Flandrian transgression in the area. The coastal relief, which suggests tectonic uplift, shows well-defined structural controls according to vertical fracturing N 60º E and N 30º W, and a highly dissected terrace, which is 80 m above present sea level. The San Juan River delta coastline is about 50  km long  and is characterised by the alternation of

­ arrier  islands, 8–12 km long, interrupted by funnelb shaped estuarine inlets (Martínez et al. 1995). The San Juan delta has an area of 800 sq km, a drainage basin of 16,500 sq km, and the highest discharge (2550 m/s), sediment load (16,000,000 t/year), and basin-wide sediment yield (1150 t/sq km/year) on the west coast of South America (Restrepo et al. 2002). The San Juan delta coastline is changing, especially along its northern and southern extremes (Boca Togoromá, Charambirá and Boca Chavica) where beach retreat has been about 100 m in the last 50 years (Correa and González 2000; Morton et al. 2000). The historical evolution of El Choncho barrier island (>Fig.  3.3.1), at the southern part of the delta, shows a general pattern of barrier narrowing and breaching that is becoming common along other barrier islands of the Pacific coast, including Charambirá (northern part of the delta), and along the Patía delta and Mira deltas to the south. According to radiocarbon dates, the Isla El Choncho barrier island began forming about 500 year bp., and grew to the southeast by successive beach-ridge accretion (González and Correa 2001). Its evolution in the last 50 years records a succession of events that segmented the island. The general retreat of the central part of the barrier coincided with the formation of an extensive sandy tidal flat at its northern end (Boca Chavica), which diminished sand supply from the north. Shore erosion was accelerated after a 20–30 cm co-seismic subsidence event generated by the November 19, 1991 earthquake. After the 1991 subsidence, the barrier flooded from 2 to 17 times each year ­during highest spring tides. Final breaching of the barrier occurred in 1997 and coincided with several overwash events associated with positive sea level anomalies of 30 cm caused by El Niño (Morton et al. 2000). Following a prudent strategy, El Choncho village was relocated to the Santa Bárbara beaches, a beach ridge remnant located 200  m inland from the coastline (Correa and Gonzalez 2000). To the south of the San Juan delta, the coastline of Málaga Bay and the northwest side of Buenaventura Bay are characterised by vertical to sub-vertical cliffs, 10–20 m high, cut into horizontal to sub-horizontal Tertiary sandstones and mudstones. The coastal relief is characterised by low hills and a terrace level of 20 m. Erosional segments of cliff shore alternate with small pocket beaches supplied with sand produced by rapid shore retreat. The numerous stacks and arches are evidence of the retreat. Depositional zones are represented by mangrove-vegetated tidal flats at the eastern sides of Málaga and Buenaventura Bays, and by the Juan Chaco beaches (>Fig.  3.3.2), an unstable supratidal feature formed with sand from the San Juan delta. South of Málaga Bay, the coast is dominated by steep,

Pacific Coast of Colombia

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⊡⊡ Fig. 3.3.1 Central part of El Choncho barrier island after breaching of the new inlet, near El Choncho village.

plunging cliffs, 10–30 m high, cut into horizontal to ­sub-horizontal Tertiary mudstones and sandstones, with notches and evidence of bioerosion at the intertidal levels.

3.  Southern Pacific Coast The Southern Pacific coast has an overall length of 650 km between Buenaventura Bay and the Mira River delta. The coast from Buenaventura Bay to Guapi Bay extends for about 150 km, with a general S 40º W trend. The southern Buenaventura Bay shore is supplied by the Anchicayá and Dagua Rivers and is a low depositional area characterised by segmented sand bars and extensive, intertidal mudflats and mangrove swamps that penetrate more than 10 km inland from the coast. Farther south, the littoral fringe represents a system of barrier islands mangrove and transitional swamps (>Fig.  3.3.3), and is supplied by numerous rivers. The largest are the Cajambre, Naya, and Micay Rivers that drain the western slopes of the Cordillera Occidental. A dense network of tidal channels interconnects the estuarine lagoons and mangrove interdistributary areas, and the freshwater swamps are backed by cliffs or gently sloping platforms cut into Tertiary sedimentary rocks. Modern barrier islands of this stretch are 2.8–5.6 km long and 0.2 –1.1 km wide and are mostly transgressive features (Martínez et al. 1995). The Patía River drains about 23,000 sq km of the Central and Western Cordilleras of the Colombian Andes. Its delta consists of a northern ancient delta plain and the

western, more recent delta plain. This zonation results from the progressive NE to SW delta migration. Both delta zones have been influenced by Quaternary tectonic movements. The Patía delta coastline is about 120 km long and extends from the Guapi Bay to Punta Cascajal, at the northwestern extreme of Tumaco Bay. The northern Patía delta fringe is a subsiding, tide-dominated area characterised by frontal, erosional lobes of sandy and muddy tidal flats backed by mangrove swamps and an ancient complex of beach ridges. Numerous funnel-shaped estuaries and tidal channels penetrate more than 15 km to the interior of the northwestern Patía delta plain, where mangrove and transitional swamps are rapidly eroding. The erosion produces a pattern of numerous isolated small islands. The western and southern shores of the Patía delta consist of external sets of barrier islands and mangrove and transitional swamps backed by intricate sets of point bars and sandy fluvial deposits of numerous abandoned courses of the Patía River. Barrier islands of this sector have the best dune development (2 m high) of the barrier islands of the Pacific Coast of Colombia. Nevertheless, the barriers exhibit transgressive characteristics with abundant breaches and island segments (Martínez et al. 1995). The entire Patía delta has been affected in historical times by co-seismic subsidence, best recorded in the December 12, 1979 earthquake that affected the Southern Pacific Coast from the Mira River delta to Guapi Bay. This event in particular caused a regional subsidence maximum of about 1.6 m along the shores of the San Juan barrier island, and its associated tsunami killed about 155 people at the

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⊡⊡ Fig. 3.3.2 Juan Chaco village at the entrance of Málaga bay, showing the 10 m terrace and cliffed coastline with caves and structurally controlled entrances cut into subhorizontal mudstones and sandstones. (Courtesy Observatorio Sismológico del Suroccidente Colombiano, Universidad del Valle.)

⊡⊡ Fig. 3.3.3 Mangrove swamp (Rhizophora mangle) along the northern border of the Guapi estuary. The mangrove trees are up to 30 m high.

San Juan de la Costa village, located on the central part of this island (Herd et al. 1981). Tumaco Bay, located between the Patía and Mira deltas, is the largest embayment of the entire Pacific Coast of Colombia. The bay is a shallow (up to 30 m deep), semiprotected area located on an active tectonic area with a N 30º E structural trend. The Tumaco Bay northern and east coasts are dominated by intertidal flats and by 10–40 m

high, vertical to sub-vertical cliffs cut into horizontal to sub-horizontal mudstones. Cliff retreat is slow in most sectors, but rock falls and slumping are caused by seismic events and occasionally strong rain. Minor pocket beaches and wave-cut platforms can be found along the eastern cliffs of Tumaco Bay. Tidal estuaries with lobe-shaped bars and tidal flats backed by extensive mangrove swamps characterise the southeastern shores of Tumaco Bay. At its

Pacific Coast of Colombia

southwestern part, highly unstable sandy barrier and bar complexes migrating from the northwestern Mira delta front the internal shores of Tumaco Bay. Most of the city of Tumaco is located on the La Viciosa and El Morro islands, the two largest sandy barriers. Both barriers are subject to rapid cycles of accretion and erosion as intertidal bars are deposited and destroyed. Offshore sandy intertidal flats protected La Viciosa and El Morro from the direct impact of the 1902 and 1979 earthquake-generated tsunamis. The Mira River delta plain is a Plio-Quaternary accretionary prism that encompasses about 1600 sq km. The delta is limited to the east by hills and terraces of semiconsolidated sediment, and to the north by the extensive mangrove and freshwater swamps bordering southern Tumaco Bay. Morphological evidence suggests that the Mira delta prograded progressively from Tumaco Bay to the southwest. The older delta is recognised by an extensive system of arcuate beach ridges landward of the more recent deposits of sandy barrier spits and mangrove swamps. South of Tumaco Bay the NW Mira delta shore is wave-dominated with narrow, erosive barrier islands backed by mangrove and transitional swamps (Martínez et al. 1995). North­ eastern barrier islands and sandy intertidal shoals and bars migrate continuously to the southwestern part of Tumaco Bay. These features represent the most active sectors of the entire Colombian Pacific coast.

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Acknowledgement The authors would like to thank Fondo BIDCOLCIENCIAS (Programa de Ciencias del Mar), grants 1216–024–99, 1216–09–153–96 and 1216–05–16911), the U.S. Geological Survey, and Universidad EAFIT for supporting the work that made this publication possible.

References Correa ID, González JL (2000) Coastal erosion and village relocation: a Colombian case study. Ocean Coast Manag 43:51–64 González JL, Correa ID (2001) Late Holocene evidence of coseismic subsidence on the San Juan Delta, Pacific Coast of Colombia. J Coastal Res 17:459–467 Herd DG, Leslie YT, Meyer H, Arango JL, Person WJ, Mendoza C (1981) The Great, Tumaco, Colombia Earthquake of 12 December 1979. Science 211:441–445 Martínez JO, González JL, Pilkey OH, Neal WJ (1995) Tropical barrier Islands of Colombia’s Pacific Coast. J Coastal Res 11:432–453 Morton RA, González JL, López GI, Correa ID (2000) Frequent nonstorm washover of barrier islands, Pacific Coast of Colombia. J Coastal Res 16:82–87 Restrepo JD, Kjerfve B, Correa ID, González JL (2002) Morphodynamics of a high discharge tropical delta, San Juan River, Pacific Coast of Colombia. Mar Geol 192:355–381

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4.0 South America – Editorial Introduction

South America consists of a major western mountain range, the Andes Cordillera, extending from Colombia southward through Ecuador and Peru to Chile and Patagonia, the broad plateau of southern Brazil, the uplands of Venezuela and the large river basins of the Amazon and the Parana (Psuty et al. 2005). The Andes are primarily the outcome of Alpine folding, while the plateaux of southern Brazil and the Venezuelan uplands are Pre-Cambrian shields.

1.  Pacific Coast In > Ecuador and northern > Peru the Andes mountains are bordered by coastal lowlands with hilly areas and deltaic plains. There are cliffs along hilly areas and bays excavated in less resistant outcrops. The coast is lined by beaches and dunes, and there are mangroves in estuaries and inlets. Further south the mountains run closer to the coast and there are many steep and cliffed sectors, with only minor, narrow coastal plains, generally at the mouths of river valleys. Earthquakes occur, and emerged beaches and shore terraces at various levels indicate a history of tectonic uplift and deformation. These landforms continue south into > Chile. In southern Chile the mountains are bordered by foothills, and coastal plains are longer and wider. Further south the coast becomes irregular, with many islands, and rias and (south of latitude 43 degrees) fiords dissect the mountain slopes. Contrasts result from climatic transitions from humid tropical Ecuador through arid Peru and northern Chile to cool temperate rainy areas in southern Chile, grading to subarctic in Patagonia. The coast bears the imprint of Pleistocene glaciations, and there are still glaciers in the mountains of southern Chile, several descending into the heads of fiords: in recent decades they have been receding. The continental shelf is narrow, and ocean swell from the SW dominates the ­coastline, with stormier conditions in southern Chile, where westerly gales are frequent. Man­groves extend south into Peru, with salt marshes in ­estuaries and inlets further south Coral reefs are poorly developed. Tide ranges are

generally less than 2 m, rising to 3.6 m in the Gulf of Guayaquil in Ecuador.

2.  Atlantic Coast In southern Argentina (>Argentina) the climate is cold. The hinterland was subject to Pleistocene glaciations. Beaches become gravelly, fed by rivers delivering fluvioglacial deposits and cliffs cut in moraines. In Tierra del Fuego there are frost-shattered cliffs, fiords, and gravel beaches derived from glacial drift. Ocean swell is accompanied by frequent storm waves, and tide ranges increase to more than 5 m, reaching 10.4 m at Puerto Gallegos. Further north, the pampas hills come to a coast with lengthy cliffs cut in Tertiary and older formations, interrupted by the Rio Colorado and Rio Negro deltas. Terraces at various levels are the outcome of tectonic uplift and sea level changes. South-easterly ocean swell has shaped sandy beaches backed by beach ridges and dunes. In northern Argentina the coast curves past the beachfringed lowlands of Buenos Aires on the southern side of the Gulf of the Rio de la Plata. The broad interior lowland drained by the Uruguay and Parana rivers opens to the Rio de la Plata through deltaic marshland. A hilly hinterland comes to the coast in > Uruguay, where headlands mostly contain granite, separate bay beaches backed by dunes and low cliffs form the northern shore of the Rio de la Plata. The coastal lowlands of southern > Brazil are fringed by long sandy beaches and barriers enclosing lagoons, notably the large Lagoa dos Patos. Lagoons and estuaries are fringed by salt marshes, and to the north of Cananeia there are also mangroves as the coastal climate becomes warmer. The coast north to Rio de Janeiro is hilly, with many irregular headlands, islands, embayments and some lagoons enclosed by sandy barriers. It is exposed to a strong SE swell from the Southern Atlantic. North of Cabo Frio there are several deltas, of which the Rio Sao Francisco is the most substantial, and the coastline is bordered by sandy beaches and biogenic reefs. The great crystalline Brazilian plateau ends in an

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escarpment that overlooks a coastal plain of varying width, fed by many rivers and interrupted by ridges that culminate in cliffy headlands. The swell from the Southern Atlantic weakens northward, particularly beyond Recife. The climate becomes subtropical as the coast swings north-westward. The vast Amazon delta then forms a long, low coastline with estuaries and rias, mangroves and occasional sandy beaches and cheniers. The Pre-Cambrian uplands of Bolivar do not extend to the coast, which consists of the wide alluvial plains of > French Guiana, > Surinam and > Guyana. Mean spring tide ranges diminish to less than 2 m, and the coast is exposed to NE waves from the Atlantic Ocean.

3.  Western Caribbean Coast In eastern > Venezuela the Orinoco valley opens to a major delta. The climate is humid tropical, with occasional

hurricanes. Dominant waves are from the NE and mean spring tide ranges are small (<1 m). A mountain range runs westward along the coast from ­northern Trinidad and the Paria Peninsula and past Caracas. It is bordered by long sectors of steep and cliffed coast, cut mainly in Tertiary sedimentary formations and fringed by narrow beaches. The coastal mountain range curves south-­ westward to become one of the three arms of the northern Andes Cordillera, separated by the Magdalena vale and the lowland around the Gulf of Maracaibo.

Reference Psuty N, Mizobe C (2005) South America, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science, Springer, Dordrecht, the Netherlands, pp 905–909

4.1  Ecuador         H. Ayon · W. Jara         

1.  Introduction The coastline of Ecuador, including the Galapagos, Puna and Jambeli islands, is about 2,500 km long. The coastal area is divided into two well defined geological regions by the northwest-southeast Chongon-Colonche Range. The northern region is characterised by ranges with horst-like blocks of pyroclastic rocks, basic lavas and dolerites which underlie Cretaceous and Tertiary sediment. The Lower and Middle Cretaceous rocks are largely volcanic deposits, whereas the Upper Cretaceous consists of sands, clays and agglomerates. The ranges are truncated by the coastline, and as a result, rocky promontories, cliffs and pocket beaches are common features along the coast. South of the Chongon-Colonche Range there is a deep basin covered by thick Tertiary and Quaternary sediment. Intensive block faulting has occurred in the region that is occupied by the Gulf of Guayaquil. The coastal region is considered to have been unstable since Pleistocene times, and continental uplift is evidenced by the occurrence of Pleistocene beach rock (locally named tablazos) as high has 200 m above sea level. The climate is equatorial, but coastal waters are cooled by the Humboldt Current and the associated upwelling of deep ocean water as this diverges from the coast. Temperatures average between 23 and 25°C, and rainfall (mainly January–May) varies between 400 and 2,000 mm with the El Niño effect: the coast becomes drier south of Bahia de Caraquez. The coast is exposed to a predominant southwesterly ocean swell, but waves are seldom higher than 3 m. The strongest and most consistent swell arrives on the coastline facing SW between Salinas and Playas, and a weaker, refracted swell comes in to the west coast, particularly between Salinas and Puerto Cayo. Between November and March there are occasional phases of northwesterly swell generated by winter storms in the North Pacific and transmitted to break heavily on NW facing shores between Valdez and Esmeraldas, and at Manta. Short, steep waves are generated locally by southwesterly trade winds on the western coastline and southeasterly trade winds in the south. Tides are ­semidiurnal, and mean spring tide range

along the ocean coastline is about 2 m (2.4 m at Manta), increasing to 3.6 m in the Gulf of Guayaquil. Beaches are generally of terrigenous sand derived from eroding cliffs and river outflow, but calcareous marine sand from the sea floor is also present, and may be locally dominant. There are some corals, and shelly organisms are abundant in the productive cool nearshore water. Mangroves and salt marshes occupy sheltered shore habitats, particularly in river estuaries and around the Gulf of Guayaquil (Cintron 1979). The coastline continues to change in response to an array of coastal processes including humid tropical weathering and biological effects (Bird 1983). On the west coast beaches are eroded by heavy northwesterly swells but rebuilt when the gentler SW to W swell revives.

2.  The Coast of Ecuador From the Colombia border along the Mataje River southward Ancon de Sardinas Bay contains many low mangrove and salt marsh islands, dissected by numerous marine and estuarine inlets in an area of about 500 sq km. Several rivers drain into the bay. Rapid sedimentation precludes deep water navigation close to the shore. Extensive offshore shoals are threaded by tidally scoured channels that run landward into more sheltered islands fringed by mangroves. To the south are the plains of the Cayapas delta, and at La Tola the beach-fringed coastline becomes smoother in outline. A hilly hinterland develops, and past Rioverdo there are cliffs and shore platforms cut into Miocene mudrocks. Seaward slumping is common in the cliffs, which descend to sandy beaches. Esmeraldas is a town beside the estuary of the Esmeraldas River, which carries a large load of silt and clay down to the coast, as does the adjacent Atacames River. The mangrove vegetation shows well-defined species zonations parallel to the coastline. Northward longshore drifting has formed mudflats to the northeast of these river mouths.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Ensenada de Atacames bay to the west curves out to Punta Galera where a ridge reaches the coast, and to the south is the bay of Mompicha, with the swampy estuary of the Repartidero River deflected north by a longshore spit at Muisne. There are a series of headlands, such as Punto Suspiro, then the large estuary of the Mache River deflected northward behind the Coijimies Peninsula. This is a barrier enclosing a mangrove-fringed bay with sand bars developing at the mouth of the bay. Shrimp (prawn) ponds are being extensively excavated in the salt marsh and mangrove areas. On the seaward side of the barrier spit a sandy beach curves south to Pedernales, where a headland is cut in Miocene mudrock. A little to the south is the Equator Monument, and the coast curves southwest to Punta Ballena. Sandy beaches alternate with cliffs and shore platforms cut in Tertiary and Cretaceous mudstone and sandstone and Cretaceous volcanic rocks. Small rivers drain into lagoons bordered by barrier spits, and where the river flows are high enough and their mouths are wide, mangrove-fringed barriers and sandy spits develop, mainly from south to north. From Punta Venado steeply-sloping cliffs in Miocene mudrocks extend down to Cabo Parada, and a steep forested slope runs along the coast to San Vicente on the northern side of the Chome estuary. This large estuary is fringed by mangroves that are tall beside the river and tidal creeks, diminishing landward to scrub and dry dusty

plains submerged only at exceptionally high tides. The mangroves have been extensively cleared and converted to shrimp (prawn) ponds. At the mouth of the estuary there is a sand bar, and on the southern shore grassy dunes on a spit that has grown northward. In Bahia de Caraquez to the south a sandy beach is backed by sloping cliffs in silty clay (> Fig. 4.1.1). Sandy beaches and sloping cliffs continue south past Punta Charapato to Bahia de Manta, where the coast curves westward past Manta to a bold promontory of PlioQuaternary sediment over Miocene mudrock. Manta is a seaport with a harbour protected by breakwaters built in 1972 and later extended, and these have intercepted sand that drifts alongshore during northwesterly swells to form a wide beach on the western side. The narrow, depleted beach to the east led to accelerated cliff recession, which was countered by the building of a sea wall. At Tarqui there has been pollution resulting from artisanal fish cleaning and gutting on the beach in a sector where cleansing wave and current scour has been diminished by nearby harbour structures. Sloping cliffs in soft silty clay run westward behind Playa Murcielego (> Fig. 4.1.2), their recession threatening cliff-top houses (> Fig. 4.1.3). The sandy beaches are highly calcareous and often shelly, and quartz is the most abundant terrigenous component. The cliffed coast swings southward to Cabo San Lorenzo, where Cretaceous volcanic rocks rise to outcrop

⊡⊡ Fig. 4.1.1 Beach in the south of Bahia de Caraquez. (Courtesy Geostudies.)

Ecuador

4.1

⊡⊡ Fig. 4.1.2 Cliffed coast west of Manta. (Courtesy Geostudies.)

⊡⊡ Fig. 4.1.3 Walled cliff sector near Manta. (Courtesy Geostudies.)

on the shore and there are cobble beaches. Upper Cretaceous sediment form hilly cliffed promontories on the coast past Punta Cayo (> Fig. 4.1.4) and there are sandy valley-mouth plains at Machalilla and Puerta Lopez, with beaches shaped by southwesterly ocean swell refracted to approach from the west.

The beaches at Machalilla and Puerto Lopez include sparse coral gravel, swept in from patchy reef gardens on rocky foundations in shallow nearshore areas. Lower Cretaceous volcanic rocks outcrop in the Colonche Mountains, just inland. Regional uplift is indicated by the occurrence of tablazos which have been mapped as far

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as 27 km inland. There is a sharp cliffy headland at Punta Lopez (> Fig. 4.1.5), where steep-sided Salanco Island stands offshore. The coastal region is dry, with areas of semi-arid scrubland. The beaches are of fine grey calcareous sand, with shelly patches and gravels near rocky outcrops in cliffs and headlands.

The sandy beach widens southward past Salanco, a resort town, and the mouth of Rio Ayambe to Mantañita Point. There are rocky stacks, encrusted with seabird guano. Manglaralto and Valdivia are in bays with sandy beaches receiving surf from southwesterly swell, and river mouths, such as the Rio Cadeate, are often blocked ⊡⊡ Fig. 4.1.4 Cliffs south of Puerto Cayo. (Courtesy Geostudies.)

⊡⊡ Fig. 4.1.5 Cliffed headland at Punta Lopez. (Courtesy Geostudies.)

Ecuador

by swash-built sand banks. To the south the sloping cliffs in soft silty clays resume. Flat-topped ridges run out to coastal headlands between bays where the sandy beaches are backed by beach ridge plains formed by intermittent progradation of the sandy coastline. There are segments of shore platform cut in indurated mudrock on the headlands. In places the coast has been cut back, threatening the coast road at San Pablo, and erosion of the beach has exposed beach rock, broken and piled up by storm waves at Punta Baruynda, to the south (> Fig. 4.1.6). The coast curves round the Bahia de Santa Elena to Salinas, a seaside town on a dry peninsula, with an esplanade behind a broad sandy beach sheltered from southwesterly swell but receiving northwesterly swell between November and March. The tabular peninsula of La Puntilla is cut in Eocene sandstone underlain by Upper Cretaceous silts and agglomerates. The coastline swings southeastward, and is thus directly exposed to southwesterly ocean swell. The relatively straight coastline crosses a hilly area, with sectors of cliff cut in MioPliocene and Quaternary sediment, some of which show landslides. These alternate with valley-mouth lowlands, where dunes have formed behind the beach, as between Chanduy and Point Piedras. The broad beaches of calcareous sand contain deposits of dark magnetite and titanium sands, which have been mined. Ocean swell breaks across sand bars in the nearshore zone, and river mouths are often dammed by barriers to form small lagoons. The coastal

⊡⊡ Fig. 4.1.6 Disintegrated beach rock south of San Pablo. (Courtesy Geostudies.)

4.1

region is dry enough for salt making, and salt water intake structures interrupt the beach west of Punta Carnero. Playas is a large seaside resort with a wide beach, popular with people from Guayaquil, and at Posorja a spit has formed as the coast curves into the Gulf of Guayaquil. The Gulf of Guayaquil is a triangular-shaped area submerged during the Late Quaternary marine transgression and partly filled with sediment brought down by the Guayas and other smaller rivers (Ayon 1976). It contains Isla Puna, a large island of Plio-Quaternary sediment (including tablazos) overlying Miocene mudrocks, which separates a narrow western entrance, Canal del Morro and a broader eastern entrance, Canal de Jambeli. There are a few low cliffs and shore platforms on the more exposed western coast, but much of the island is fringed by sandy beaches and spits enclosing tidal mudflats and mangrovefringed lagoons. Mangroves and salt marshes occupy the concave eastern coast, and there are recurved spits at the southwestern end of the island. The Canal de Morro and Canal de Jambeli lead north into branching mangrove-fringed inlets and swampy deltaic islands, and the Guayas estuary, beside which stands the city of Guayaquil. There is much floating Eichhornia, derived from the freshwater swamps north of Guayaquil. Tidal currents are strong, and there is an extensive and variable ebband-flood shoal and channel topography with megaripples. On the western side is the intricate Estero Salado, a channel about 42 km long in which tides attain 5 m. As mangroves

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spread on to accreting mudflats intervening channels are scoured more deeply (Cintron 1979). On the eastern side of the Gulf of Guayaquil a swampy coastal lowland extends south to Machala. Small intertidal deltas emerge during low tide at the mouths of the rivers running at right angles to the shore through the mangrove fringe. Shrimp (prawn) ponds have been excavated in salt marshes and saline flats, and in the mangroves. Emerged deposits of Crassostrea shells, up to 10 m above sea level, are evidence of Holocene uplift. The coastal plain ends in the mangrove-fringed Jambeli Archipelago which extends over an area of 260 sq. km, close to the Peruvian border. Sandy beaches develop as wave energy increases south of the protection of Puna Island, and sand spits are growing eastward from river mouths on the Tumbes delta.

3.  Galapagos Islands These islands form the Galapagos Province of Ecuador, and are famous for their unusual fauna, including giant turtles, described by Charles Darwin when he visited the islands in 1835 during the voyage of the Beagle. They are a group of 19 volcanic islands (and numerous islets and reefs) on the Equator in the eastern Pacific Ocean, about 965 km west of the coast of Ecuador. They formed as shield volcanoes rising 3,000 m from ocean floor fractures, and consist largely of superimposed basalt lava flows. They formed in response to a hot spot that has migrated westward, so that the older islands with extinct volcanoes are in the eastern part and the younger ones with active volcanoes to the west. Isla Fernandina is a volcanic cone rising to La Cumbre (1,463 m). It occupies a bay on the western side of Isla Isabela, a large curved island with several volcanic

peaks, the highest of which is Cerra Azul (1,689 m) in the north. Isla San Salvador, Isla Pinzon, Isla Santa Cruz, Isla Santa Maria and Isla Espando are all simple volcanic islands. Eruptions have continued on a small scale. The occurrence of submarine lavas and Pliocene sedimentary rocks indicate that some of the islands have been raised about 100 m above sea level. The climate is dry and cool as a result of the cold Humboldt Current arriving from the south, and the islands are sparsely vegetated, with extensive rocky areas. Wave action is stronger on the south and west facing coasts, which receive swell transmitted from the Southern Ocean across the South Pacific. The coasts of the volcanic Galapagos Islands show rugged cliffs, shore platforms, some dark beaches of basaltic sand and gravel and a few beaches of white calcareous (coralline and shelly) sand, notably on the older eastern islands. Perry Isthmus on Isla Isabela is a sandy tombolo. Some of the steeply plunging cliffs are associated with laterally eroded craters and calderas. Dissection along joints and planes between lava flows has guided the shaping of caves and natural arches. Shore platforms are developed on seaward protruding lava flows. There is a large blowhole at the western end of Isla Espando, and one of the islets is pierced by a natural arch.

References Ayon H (1976) Sediments of the Guayaquil estuary. Instituto Ocean­ ografico de la Armada del Ecuador Bird ECF (1983) Shore procersses in Ecuador. In: Vallejo S (ed) Ordenacion y Desarrollo Integral de las Zonas Costeras, Armada del Ecuador, pp 29–45 Cintron G (1979) Mangroves of the Ecuador coast. Departamento de Recursos Naturales, Puerto Rico

4.2 Peru

Vidal Taype Ramos · Eric Bird

1.  Introduction The coastline of Peru (>Fig. 4.2.1), just over 2,300 km long, is notably arid (Dresch 1961). Mean annual rainfall at coastal centres is low (Chiclayo 15 mm, Trujillo 28 mm, Lima 31 mm and Mollendo 22 mm) and variable; desiccation is intense between the occasional rains. This is because coast­­al waters are abnormally cool for the latitude (4°–18° S), due to the upwelling of cold water from oceanic depths as a result of trade winds blowing offshore. Sea breezes keep the coastal fringe cool, with frequent fogs but very little precipitation. Vegetation is consequently very sparse, and desert landscapes predominate. Tectonic uplift accompanied by faulting has raised the high hinterland of the Andes Mountains, the structural grain of which is followed by the coastline (Kellog and Mohriak 2001). Coast Ranges, extensive in Chile, are here confined to the sector between Pisco and Mollendo (Noble et al. 1978) with an outlier near Punta Aguja; intervening segments have subsided beneath the sea. Earthquakes still are found. In May 1970 a major earthquake centred 50 km off Chimbote caused massive avalanches in the mountains and severe flooding down the Santa River valley to its delta region. Many valleys are incised into the coastal slopes, but only streams fed by mountain rainfall and melting snow maintain their flow through to the coast, and even these show great variations in seasonal and annual discharge. The coastline is relatively simple in outline, with few estuarine inlets and only minor deltaic protrusions built in the face of oceanic wave action (swell generated in the Southern Ocean crosses the South Pacific to arrive on this coast from the south-west. Cliffs cut by marine erosion alternate with sandy beaches derived mainly from shoreward drifting of sea-floor sediment and some sand and shingle beaches nourished from cliff erosion and fluvial outwash. Beaches are backed by unvegetated dunes, or have only a sparse plant cover sustained by fog and dew (Dresch 1961). In many places sand is drifting inland to be banked against cliffs or rising slopes, or to block the mouths of coastal valleys and gullies. Ocean swell arriving from the southwest is often accompanied by small waves generated in coastal waters by southerly breezes. These waves arrive at an angle to the

coastline and cause northward drifting along beaches. Occasionally in summer there are northwesterly waves (which cause southward drifting) and penetration by the warm El Nino current to the coast north of Negritos. Mean spring tide ranges are small, generally less than a metre (0.7 m at Callao, the port of Lima), but they in­­ crease northward to the Ecuador border and the Gulf of Guayaquil.

2.  The Peruvian Coastline Close to the Ecuador border is the large, sandy Tumbes Delta on the southern shores of the Gulf of Guayaquil. Its estuarine distributaries are mangrove-fringed (the southern limit of mangroves on the west coast of South America is here at Punta Malpelo, 3° 4' S, because northward drifting carries seedlings in that direction). Parallel sandy beach, ridges, emphasised by lines of shrubs, extend southward to Caleta Le Cruz, the landing place of Fransisco Pizarro in 1527. Southwest from Zorritos the coastline is generally cliffed, cut into a fringe of uplifted Tertiary formations that dip gently seaward and are capped by a series of Quaternary marine terraces (tablazos), incised by transverse valleys (quebradas), but otherwise well preserved in this arid climate. Each terrace is a sparsely vegetated plain strewn with broken, sand-blasted pebbles, with old beach deposits to the rear, beneath bluffs that mark degraded cliffy coastlines. The oldest and highest, the Mancora terrace, stands about 360 m above sea level in the north and declines to 75 m southward. The next, the Talara terrace, falls southward from 90 to 40 m, while the youngest and lowest, the Lobitos terrace, declines from 40 to 30 m southward (Bosworth 1922). These terraces are truncated by cliffs along the present coastline: at Cabo Blanco, for example, the Mancora terrace ends in cliffs 300 m high. An early Holocene coastline is marked by gullied bluffs behind intervening sectors of coastal plain, as at Talara. The beaches are of shelly sand with some stony material, and the coastal plains include parallel beach ridges (Wells 1996), often overlain by dunes that have locally been banked against former cliffs 45–60 m high.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4. 2

Peru

⊡⊡ Fig. 4.2.1 The Peruvian coastline. (Courtesy Geostudies.)

THE PERUVIAN COASTLINE

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Near Negritos the cliffs are 80 m high, incised by valley mouths blocked by 20 m high dunes. Southeast from Parinas Point the sandy beach, backed by dunes, is 50 km long, interrupted only by the mouth of the Chira River. Behind the dunes is a flat plain inundated by the highest tides, and covered with evaporite deposits.

75°

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208

Paita is on a hilly peninsula of Mesozoic rock, but the bay to the south is sandy, fronting abandoned cliffs, and paired spits border the small estuary of the Rio Piura. Punta Agula is another hilly peninsula dissected by wadis, and followed southward by a sandy coastline fringing the Sechura Desert. The extensive dunes here include

Peru

4. 2

⊡⊡ Fig. 4.2.2 Indented coastline north of Huarmey. (Courtesy Geostudies.)

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barchans up to 6 m high and 15 m wide. Evaporation is intense, and the Rio Cascaja fades away before reaching the sea. The coastal plain narrows southward from Chiclayo to Trujillo, beyond which a steep coast predominates as the Andean foothills approach and intersect the coastline. Cliffs and bluffs are cut into Mesozoic rocks and intruded granites and there are rugged islands offshore (e.g., Isla de Macabi and Islas Guanape), many with thick guano deposits. Sectors of sandy beach (playa) also are found, backed by beach ridges, but south of Chimbote the coastline is hilly and embayed on Cenozoic volcanics, with shore platforms fronting cliffs and the harder elements persisting as rocky headlands, with caves and coves. The Peninsula de Ferrol is a tombolo, a high hill at­­ tached to the mainland by a sandy isthmus; its rocky seaward margin has bouldery clefts and a natural tunnel. A scalloped cliffy coastline (>Fig. 4.2.2), notched by wadis, near Huarmey, is fringed by many white or yellow guano-mantled islands and headlands. The Rio Pativilca has built a blunt delta, while to the south Les Salinas is a small elongated barrier lagoon, and Playa Grande has

heavy surf breaking on a shingle beach fronting bluffs cut into the seaward margin of a gravelly pediment. Callao, the port of Lima, has large breakwaters enclosing a harbour, with accretion of sand from Rio Rimac immediately to the north. A major spit (La Punta) protrudes in the lee of Isla San Lorenzo, a hilly offshore island on which Quaternary marine terraces are well developed. Southward the coast is undulating, with small plains at valley mouths, and the Paracas promontory at Pisco marks the beginning of a coastal mountain range. Punta Zarata is a cuspate foreland enclosing a lagoon and salt marshes, and a spit partly encloses Laguna Grande on the east coast  of the Carreta Peninsula, bordering Bahia de la Independencia. Asymmetrical bays, shaped by the refraction of southwesterly swell around headlands, have developed at Bahia de Lomitas, Bahia San Nicolas, and Bahia San Juan. Lo­­ cally, pebble beaches have formed from material derived from Pleistocene conglomerates deposited at old valley mouths and now exposed in cliff sections. The next part of the coastline faces southwest and receives strong surf produced by swell arriving through

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⊡⊡ Fig. 4.2.3 The blunt delta of the Rio Tambo, built on a coastline exposed to southwesterly ocean swell, which has built barrier beaches impounding small, narrow coastal lagoons. (Courtesy Geostudies.) 60 0

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⊡⊡ Fig. 4.2.4 Playa de El Palo, beach-fringed cliffs cut in gravelly pediments. (Courtesy Geostudies.)

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Peru

deep water inshore. Drifting sand blocks the mouth of the Ico River valley. To the south the coast is steep on granitic rocks, with many gullies, some of which reach the sea in small coves. Rio Ocona has a small estuary, but Rio Camana has built a blunt delta with beach ridges and swale lagoons, as has the Rio Tambo southeast of Mollendo (>Fig. 4.2.3). Rio Vitor, by contrast, reaches the coast through a deep gorge. Between Mollendo and the Chilean border are several straight beaches fronting bluffs that truncate pediments: Playa de El Palo (>Fig. 4.2.4), Playa Tacahuay and Playa Inglesa.

4. 2

References Bosworth TO (1922) Geology and palaeontology of Northwest Peru. Macmillan, London Dresch J (1961) Observations sur le désert cotier du Perou. Ann Géog 70:179–184 Kellog JN, Mohriak WU (2001) The tectonic and geological environment of coastal South America. In: Seeliger U, Kjerfve B (eds) Coastal marine ecosystems of Latin America. Ecological Studies 144:1–16 Noble DC, McKee EH, Megard F (1978) Eocene uplift and unroofing of the coastal batholith near Lima, Peru. J Geol 86:403–405 Wells LE (1996) The Santa beach ridge complex: sea-level and progradational history of an open gravel coast in central Peru. J Coastal Res 12:1–17

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4.3  Chile   José Araya-Vergara  

1.  Introduction The coast of Chile is about 6,435 km long. Tide ranges are small in the north, Arica having a mean spring tide range of 1.4 m and Iquique 1.5 m. They increase southward to Valparaiso (1.66 m) and Caleto Mansa (1.62 m), and are much larger in the Gulf of Corcovado (Puerto Montt 5.8 m). Further south they are generally between 1.2 and 1.4 m, rising to 2.06 m in Orange Bay, near Cape Horn. Wave action is dominated by the Pacific ocean swell from the southwest and storms in coastal waters that produce waves mainly from the west. The climate is hot and dry in the northern desert, mild and Mediterranean in the centre (winter rainfall and dry summers: Antofagasta has mean monthly temperatures of 20.6°C in January and 14°C in July, with an average annual rainfall of 12 mm) and cold, wet and stormy in the Patagonia to the south.

s­ upplied directly from the Copiapó River (Araya-Vergara 2001), but dune sands of the ergs located in basins many kilometres inland had a marine origin. South from the Copiapó River, all the principal rivers flow permanently to the sea through microtidal estuaries (>Fig. 4.3.4). Over more than 700 km of coastline, increasing supplies of fluvial sediment to the shore form beaches in zeta form embayments. Northward drifting within these results in wider and higher beaches and dunes at the northern ends, as in Tongoy Bay and the Los Vilos embayment. Between the estuaries of Petorca-Ligua and Maipo rivers there is a southward transition from rias to estuarine deltas, as in the Aconcagua estuary (>Fig. 4.3.5).

⊡⊡ Fig. 4.3.1 Map of northern Chile. 71°W 18°S

2.  The Chilean Coastline

Iquique

1110 km

South from the Peruvian border at Arica the coast of the Atacama desert is mainly cliffed (>Fig. 4.3.1). Near Iquique, massive transgressive dunes (>Fig. 4.3.2) have been deposited between the present beach and the local mega-cliff. Due to their position with respect to present sea level and the degree of weathering of their sands (Potter 1994) they are thought to have formed during the Pleistocene (Paskoff et al. 1998), when sea level was lower. They are a striking feature, because dunes are scarce on the cliffy coast of the desert. North of Antofagasta cliffs are cut in horizontally bedded Tertiary sandstones (>Fig. 4.3.3). In some parts the coastline has embayments cut in sandstone between protruding headlands where harder rocks outcrop near mean sea level. There are emerged terraces, as at Cobija, where a 15-m Last Interglacial terrace fronts an escarpment. At Caldera, south of Antofagasta the Copiapó River is the first river outlet with permanent runoff to the sea, after more than 1,000 km of desert coast. There are Pleistocene marine terraces with deflated beaches on their surfaces (Marquardt et al. 2000), and the ergs of the Atacama Desert. A large part of the sand of the ergs has been

Arica

Ancient transgressive dunes

La Portada: Soft cliffs and arces Antofagasta 71°W 25°S

Caldera Copiapó River

Ergs Marginal Desert of Atacama Coastal terraces isotopically correlated First river oulet with permanent runoff to the sea in the marginal desert

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 4.3.2 Transgressive dunes near Iquique, fronting a megacliff. The dune crests change in relation to wind direction.

⊡⊡ Fig. 4.3.3 Sandstone cliffs at Antofagasta, showing contrast between slope on upper softer beds and cliff on harder sandstone.

Between the Maipo River outlet and San Antonio harbour the coastline has been straightened as the result of accretion alongside a breakwater. The coastline advanced 600 m between 1908 and 1934 forming beach ridges, foredunes and a lagoon, but since 1934 the ­coastline has stabilised (Araya-Vergara 1985). At Santo Domingo sand formed beach ridges and successive parallel foredunes (>Fig.  4.3.6), four of which were added during the past 600 years (Paskoff et al. 2000). The penultimate ridge was monitored during the storm surges of 1968 and 1982, when the dune was cliffed. After 1982 a wide beach formed and microdunes developed, and by 1993 the present seaward foredune had been built (>Fig. 4.3.7). South of Santo Domingo barchans behind the beach diminished in number and area between 1955 and 1980, becoming more scattered (Araya-Vergara 1987). This has been attributed to a reduced sandy supply and abnormal heavy rain. Between Rapel and Bio-Bio Rivers ebb and flow deltas are typical (>Fig.  4.3.8). The deltaic estuarine banks become more stable south to the Maule estuary then decrease and become ephemeral down to the Bio-Bio estuary. Tidal deltas are rounded and lobate on the seaward side, where wave energy is higher, and more irregular on the landward side. Off Valparaiso are the Islas Juan Fernandez, where the Scotsman Andrew Selkirk was marooned by his captain, William Dampier, in 1704. Robert Louis Stevenson wrote the story of his survival (as Robinson Crusoe) on these remote and uninhabited Chilean islands, one of which is now called Isla Alejandro Selkirk and another Isle Robinson Crusoe.

Chile

⊡⊡ Fig. 4.3.4 Map of southern Chile.

Copiapó R.

4. 3

First river outlet with permanent runnof after the desert

Elliptic and zeta form bays with progressive surf zone change down drift Tongoy

740 km

72° 31°

Complex intermediate surf zone in front of estuary, development of beaches and dune systems

Choapa R. Los Vilos

Petorca-Ligua Aconcagua R.

Elliptic and zeta form bay

Transition from prograded rias to estuarine deltas

Valparaiso San Antonio

Dated man induced straightening

Maipo R. Sto. Domingo

Dated foredune accretion and cut and fill

⊡⊡ Fig. 4.3.5 Aconcagua estuary, with paired spits.

Zeta-form bays are common in this sector, and show increasing volumes of sand downdrift to the north, where there are three generations of dunes: early Holocene, middle Holocene and presently active (Araya-Vergara 1989). The Gulf of Arauco (>Fig.  4.3.9) is a zeta-form embay-

ment backed by a cliff cut in soft sediment. The width of the beach and surf zones increases downdrift (northward), indicating sand drift towards the Bio-Bio estuary. Similar features are seen in the embayment between Lebu and Quidico. Associated with these are the Arauco dunes, the

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Chile

⊡⊡ Fig. 4.3.6 Parallel foredunes built upon beach ridges on the prograded coast at Santo Domingo. The inner dunes were formed before 600 years bp, and a broad swale separates the inner and outer zones.

⊡⊡ Fig. 4.3.7 At Santo Domingo the foredune was cliffed by storm surges in the winter of 1982, and there is a narrow eroding beach.

largest coastal erg in Chile, with small impersistent compound barchans because the climate is humid, and there are annual floods. The Imperial River meanders across a terrace which is inset in an ancient ria. To the south is a ria coast, formed by partial sub­ mergence of river valleys incised into the schistose coast ranges. The intervening sectors are cliffy. The

­ orthernmost limit of direct influence of the Last n Glaciation is marked by a transition from rias to fiords, förde and submarine glacial piedmonts. The floors of fiords have submerged terraces formed by deposition of fine sediment during Late Glacial and Holocene times, and there are steep bordering glaciated slopes (>Fig.  4.3.10). Deltas at the heads of fiords and mouths of side valleys have streams

Chile

4. 3

⊡⊡ Fig. 4.3.8 Estuarine delta.

⊡⊡ Fig. 4.3.9 Transitional zone between estuarine deltas and rias. Zeta-form embayment

Bío-Bío estuary Gulf of Arauco Lebu

Erg of impersistent barchans

Aprox. 480 km

Quidico

Imperial estuary

Valdivia ría and estuary

Cliffs and erosioanal features

41° 75°

RIAS

Bahía Mansa

which discharge clouds of fine sediment that settle on the fiord floor. Sedimentation rates of some decimetres in the last century have been measured. To the south periglaciated coastal slopes descend to the sea, undercut by marine cliffs on exposed sectors. Active periglaciation results in the formation of screes on steep slopes. In the inner coastline of Chiloé Island the key landforms are förde, formed by marine submergence of former subglacial channels. There are many islands of soft glacial drift with coastlines smoothed by erosion, longshore drift and deposition. Tidal currents are strong because of the large tide range. To the south the coastline in front of the Patagonian ice fields is trenched by deep fiords floored by morainic banks, which were formed during the Last Glaciation and Late Pleistocene to Holocene terraces (ponding esplanades) (Araya-Vergara 1999). Above the fiords are the Patagonian ice fields, with calving glaciers at fiord heads. Glacifluvial outwash produces turbidity plumes and deposits fine to coarse sediment in the fiords. The glaciers have been receding. The San Rafael glacier was an ice lobe protruding into the sea in the 1860s, but by the 1990s it had retreated into a fiord. It is notable for its 70-m high calving ice cliffs. The Magellan region was shaped during the last glaciation, producing fiord and piedmont coastlines. The Strait of Magellan originated as a fiord in the western part, which is similar to the fiords in front to the Patagonian ice field, and a set of piedmont lobes in the eastern part with many cliffs cut in soft glacial, but the fiords are smaller.

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Chile

⊡⊡ Fig. 4.3.10 Aysen fjord, with a fjord-head delta and turbidity plume. On the left a hanging valley, and side delta, on the right roches moutonnées on the valley side and a pocket beach.

In the eastern part of the Strait of Magellan the tide range is large and intertidal zones, including well-­ developed shore platforms, are wide. Tidal currents are very strong (e.g. 400 cm/s.) in narrow straits, where they contribute to erosion of the coastline. Some valley mouths have small deltas formed in the Holocene.

References Araya-Vergara JF (1985) Sediment supply and morphogenetic response on a high wave energy west coast. Z Geomorphol Suppl. (Bd) 57:67–89 Araya-Vergara JF (1987) Significance of barchans in beach-dune system interactions in Central Chile. Thalassas 4(1):23–27 Araya-Vergara JF (1989) Remnant coastal dunes and their significance in Chilean ergs. Geoško-Plus 1:15

Araya-Vergara JF (1999) Perfiles longitudinales de fiordos de Patagonia Central. Cienc Tecnol Mar 22:3–29 Araya-Vergara JF (2001) Los ergs del desierto marginal de Atacama. Chile Invest Geogr Chile 35:27–66 Marquardt C, Lavenu A, Ortlieb L (2000). Neotectónica costera en el área de Caldera (27°–28°S), Norte de Chile. In: IX Congreso Geológico Chileno, Puerto Varas, Chile Actas 2, pp 582–588 Paskoff R, Luitiño L, Petiot R (1998) Carácter relicto de la Gran Duna de Iquique, Region de Tarapacá, Chile. Rev Geol de Chile 25(2):255–263 Paskoff R, Manriquez H, Cuitiño, L, Petiot R (2000) Caracteristicas, origen y cronologia de los cordones dunares de la playa de Santo Domingo, Region de Valpararaiso. Chile Rev Geol de Chile 37(1):121–131 Potter PE (1994) Modern sands of South America: composition, provenance and global significance. Geol Rundsch 83:212–232

4.4  Argentina

Enrique Schnack . Jorge Pousa . Germán Bértola . Federico Isla

1.  Introduction The coastline of Argentina is about 5,700 km long, with coastal landforms shown in (>Fig. 4.4.1). The coastal geology ranges from a fairly stable area in the north to a generally rising area in Patagonia and tectonically-affected coasts in Tierra del Fuego. A continental shelf, in places, over 800 km wide, is covered by terrigenous sediment accumulated during Quaternary sea level oscillations. Pebbles are a common component of Patagonian beaches, but sandy beaches are typical along the coast from Cabo San Antonio to Mar Chiquita lagoon, and along the southern Buenos Aires barrier coast. Salt marshes are present along the mostly macrotidal Patagonian coast, and on mesotidal areas, with extensive development in Bahía Anegada and Bahía Blanca, where freshwater influence is absent except for river outlets. Brackish marshes with freshwater mixing are found in the northeastern sector of Buenos Aires Province (Samborombón Bay, Mar Chiquita Lagoon). Cliffs are extensive on the Patagonian coast, but the Buenos Aires (Pampas) coastline alternates between low-lying and cliffed areas. Evidence that the sea stood at higher levels dur­­ ing  Pleistocene interglacial phases and the postglacial Holocene is seen on the Argentine coast as beach ridges, marine terraces and estuarine deposits, in most cases hosting a molluscan fauna which has enabled dating of the deposits. In many low-lying areas with marshes evidence of Quaternary emerged shorelines is found, and salt marsh development is strongly related to the postglacial sea level fluctuation. The climate of the Argentine coast is cold and humid in Tierra del Fuego, arid and semi-arid from Rio Gallegos to Bahia Blanca and temperate-humid from there to the Paraná Delta. Predominant wave directions for the Argentine coast are S, SE and E, with wave periods be­­ tween 6 and 16 s, the latter corresponding to ocean swell approaching from the south. There are stormy periods with waves up to 5 m high along the eastern Tierra del Fuego and on the Buenos Aires Atlantic coast. Tides along the Atlantic coastline and on the Rio de la  Plata are predominantly semidiurnal. Tide ranges are

large (megatidal) on the east coast of Tierra del Fuego and along most of the Patagonian coast, with maximum spring amplitudes reaching more than 10 m at San Sebastian Bay (Tierra del Fuego), and about 10 m in Magellan Strait, at Rio Gallegos and at Punta Loyola (Santa Cruz province). Puerto Gallegos has 10.4 m, Puerto Santa Cruz 9.5 m, and Bahia Oso Marino 9.2 m. Mesotidal environments are present from the area north of the Rio Negro outlet (3.3 m at Punta Redonda) to Monte Hermoso, near Bahia Blanca. From there northward the coast is typically microtidal: Mar del Plata has 1.1 m and Buenos Aires 1.0 m. Many storm surges have been recorded along the Argentine coast simultaneously with the northward travelling tidal wave. The duration of these storm surges ranges from a few hours up to 2 or 3 days. They are basically produced by the combined action of an anticyclone located to the west of Argentina (semi-permanent Pacific anticyclone) and a cyclone located over the Atlantic to the SE of Argentina, the latter moving towards the east or northeast. Because of this situation, strong winds from the south or southeast and high water levels affect the whole Argentine coast, as well as the Rio de la Plata shores. Erosion is in progress on the sandy barriers of northern Argentina, particularly when a storm surge coincides with high spring tides. Storm surges are considered the most significant natural agent for coastal erosion on the eastern coast of Buenos Aires, and the Rio de la Plata area. Historical mean sea level trends studied from tide gauge records (Lanfredi et al. 1998) show a rise of 1.6 ± 0.1 mm/ year over 70 years at Buenos Aires, 1.4 ± 0.5 mm/year for Mar del Plata and 1.6 ± 0.2 mm/year for Quequén. Recent calculations for Buenos Aires and Mar del Plata over an extended period show a coincidental rise in sea level of ~1.6 mm/year (Pousa et al. 2007; D’Onofrio et al. 2008).

2.  The Argentine Coastline The southernmost coast of Tierra del Fuego (northern Beagle Channel coast and southern Atlantic coast) extends for 220 km from west to east. It is an indented rocky coastline with pocket gravel beaches in embayments.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4. 4

Argentina

⊡⊡ Fig. 4.4.1 Predominant landforms on the Argentine coast. BUENOS AIRES Río

R

ío Sal ad

de

o

TA

36°s

N

D

IL

VE

IA

NT

BAHIA BLANCA

MAR DEL PLATA

MONTE HERMOSO

0m

gro

Ba hi Saan dAentonio Caleta Valdés

ub Ch

Río

Ne

IA

40°s

d o

Rí o

Colora

Bahía Samborombón Cabo San Antonio

Laguna Mar Chiquita

AN

Río

la Pl ata

20

220

ut

44°s

Deltas Estuaries and marshes

Bahía Bustamante

COMO DORO RIVA DAVIA

Sandy shoreline Pebbly shoreline Cliffs

48°s

Ice-fringed coast

Bahía San Julián Río Santa Cruz

60°s

Islas Malvinas

Río Gallegos

w

74°

25° w

Estrecho de

52° Magallanes

Bahia San Sebastíán

0 100 200 300 km

Isla de los Estados

70°w

66°w

62°w

58°w

54°w

50°w

Argentina

The ­channel is 5 km wide, its average depth between 100 and 450 m. Holocene raised beaches occur at various elevations along the southern coast. The Beagle Channel is a tectonic trough that was completely covered by ice during the Last Glaciation and occupied by a glacial lake at about 9,400 bp. It was invaded by the sea at about 8,200 bp and the marine environment was fully established along the channel by at least 7,900 bp, with a maximum sea level attained between 6,000 and 5,000 years bp (Rabassa et al. 2000; Bujalesky 2000). The Beagle Channel is an estuarine (fiord) system controlled by significant and seasonal freshwater input and by tidal inflow from both west (Pacific) and east (Atlantic). The channel is microtidal with a semi-diurnal regime. Mean tide range is 1.1 m at Ushuaia and the tidal wave moves from west to east. The embayments of the indented rocky coastline of the Beagle Channel are related to tectonic alignments affected by successive glaciations, and are occupied by small gravelly pocket beaches. Along the Beagle Channel Holocene raised beaches stand up to 10 m above mean sea level, and are often capped by anthropogenic shell middens. The oldest Holocene coastal deposits may be partly the result of isostatic recovery while the younger levels are due to recent tectonic uplift. The Bahía Lapataia-Lago Roca valley (20 km west of Ushuaia) is a palaeofjord that was occupied by a lateral and tributary valley glacier during the Last Glacial maximum (about 18,000–20,000 years ago). Well-rounded glacially moulded rocky hills (roches moutonnées), lateral moraines and kames are present in this area and Sphagnum peat bogs occupy the lowlands. Holocene marine depositst are scattered along Bahía Lapataia, Archipiélago Cormoranes, Río Ovando, Río Lapataia and the eastern coastline of Lago Roca, overlying glacial landforms and reaching a maximum altitude of at least 8.4 m above mean sea level (Bujalesky 2000). The eastern (Atlantic) coast of Tierra del Fuego extends for 330 km from SE to NW. It is macrotidal to megatidal, and exposed to high energy waves and strong westerly winds. Extensive wide beaches are composed of gravel and coarse sand. Pleistocene glacial drift deposits form high cliffs in its northern section, and these and other submerged glacial deposits have supplied sediment to beaches. There are gravelly beaches and a narrow Holocene gravelly beach ridge plain (250 m wide) runs between Cabo Domingo and the Río Grande inlet, attached to the base of an Upper Pleistocene marine terrace. In the northern part of the coast San Sebastian Bay is a semicircular embayment (55 km by 40 km) occupying a

4. 4

wide depression formed by glaciers during the Pleistocene and reshaped by the sea during the Holocene transgression. It includes a fossil marsh, the upper marsh, controlled by deflation and dominated by Salicornia sp. scattered on the mudflat, and Spartina grass progressively buried on the windward sides, gravel ridges and cheniers. The mudflat occupies much of the tidal flat. At low tide meandering tidal channels are seen. Península El Páramo is a 20 km long gravel spit barrier that partly shelters San Sebastián Bay. Cliffs extend 40 km from Cabo Nombre (10 m high) to Cabo Espíritu Santo (90 m high) in the northern part of Tierra del Fuego. They are cut in glacial drift deposits and Tertiary silty sands. At Punta Sinaí is an extensive erratic boulder field, and erratics can be seen on the beaches. In Patagonia most of the estuaries (with the exception of the Chubut, Negro and Colorado rivers) have low discharges, and tidal processes are very active at their outlets. Beaches are typically gravel-dominated, showing at times several ridges on the backshore and on the beach face (>Fig. 4.4.2). In some protected embayments, the coarse substrate carries marsh vegetation (>Fig. 4.4.3). Patagonian gravels extend over the whole region. They have been reworked by marine action at various higher sea levels and are typical of the Pleistocene and Holocene raised beaches as well as the modern shore deposits, which contain varying proportions of sand. Much of the coast has cliffs (>Fig. 4.4.4), cut in Tertiary continental and marine deposits and Jurassic volcanic rocks, and there are many embayments. Evidence of Pleistocene and Holocene high sea level stands extends along the coast, where six marine terraces, at least three of which are of Quaternary age, have been recognised. The most recent (Holocene) reaches an altitude of 8–12 m above mean sea level and extends from eastern Tierra del Fuego to northern Patagonia. The coast of Península Valdés has cliffs and shore platforms cut in Miocene sedimentary rocks, and the Caleta Valdez lagoon is bordered by a series of Pleistocene emerged gravelly beaches (Schellmannn and Radtke 2003). Marshes are found in some protected low energy environments. Except for the northern (Río Negro-Bahia San Blas) marshes, which are mesotidal, the Patagonian wetlands are mostly macrotidal. They are present either in estuaries (Chubut, Deseado, Santa Cruz, and Gallego Rivers, and in NE coast Tierra del Fuego) or within sheltered coastal bays. Some of the marshes are eroding, exposing peaty microcliffs. San Antonio Bay, which has a maximum spring tide range of 9 m, is bordered by pediments, hollows, Pleistocene and Holocene coastal and marine deposits and

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⊡⊡ Fig. 4.4.2 Gravel beaches and berms near Bahía San Julián (Santa Cruz, Patagonia).

⊡⊡ Fig. 4.4.3 Salt marsh vegetation on a gravelly beach near Bahía San Julián.

marshy environments (Schnack 1985). Extensive sandy shoals are exposed at low tide at the outer part of the bay. There are slumping cliffs up to 8 m high in Miocene sediment at Las Grutas.

The coast from San Antonio Bay to the mouth of the Rio Negro has alternations of beach and dune-fringed low coasts and cliffs. The Rio Negro estuary has a mesotidal regime. Two shoals (Banco Miguel and Banco La Hoya)

Argentina

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⊡⊡ Fig. 4.4.4 Cliff and abrasion platform at Peninsula Valdés.

form an ebb delta composed of well-sorted fine sand. The assymetry of the delta is mainly due to northward longshore drifting. From the Rio Negro estuary to the Rio Colorado delta the coast is predominantly low-lying, except for cliffs immediately south of San Blas. Pebbly beach ridges of Holocene and Pleistocene age are also present near San Blas. The Rio Colorado is generally regarded as the northern limit of coastal Patagonia. The south Buenos Aires (Pampas) coast extends from Bahia Blanca to the Paraná river delta. The mesotidal Bahía Blanca estuary is formed by several channels oriented NW-SE and separated by extensive tidal flats and islands. The area between Bahía Blanca and Bahía San Blas may have formerly been part of the delta of the Río Colorado, and the tidal flats are derived sediment. Mesotidal beaches, typically sandy, extend from Monte Hermoso towards Miramar, alternating with cliffs (Isla et al. 1996). In low-lying areas, such as Claromecó and Quequén there is evidence of Pleistocene and Holocene high sea levels. North and south of Mar del Plata high cliffs cut in Plio-Pleistocene sediment are retreating rapidly, at times reaching 1–2 m/year. At Mar del Plata the cliffs are in more resistant Lower Palaeozoic quartzites. From Mar Chiquita beach and tidal inlet north to Punta Rasa is a low-lying coastal plain, with dunes on a major coastal barrier. The geological substratum of this segment consists of Holocene estuarine and marine deposits, although evidence of former Pleistocene sea levels is also found.

At Mar Chiquita these deposits are well developed, with shelly beach ridges and the lagoon marginal flat built on the estuarine deposits. Following a maximum sea level stand of about 2.5 m above present level about 6,000 years ago the sea fell, and a southward-prograding barrier enclosed the Mar Chiquita lagoon. The tidal inlet has shown northward migration interrupted several times by human intervention. Severe erosion is occurring at Mar Chiquita beach immediately south of the inlet. Extensive dunes along the coast north of Mar Chiquita are in many places fixed and urbanised. Wave energy generally diminishes along the coast from Mar del Plata to San Clemente del Tuyú, but at Punta Médanos larger waves are due to the nearshore presence of linear sand ridges oriented SSE 30° to the coastline, which concentrate wave energy. These ridges continue south beyond Pinamar. At Mar de Ajó mean wave heights are 0.68 m, whilst maximum annual wave heights are 1.31 m. Erosion has lately dominated the Buenos Aires coastline northward to the Mar Chiquita lagoon. The southern beaches (Mar del Plata, Santa Clara, Mar de Cobo) are typically reflective, while the beaches at Villa Gesell, Pinamar and Partido de la Costa generally have a dissipative profile with the development of sand bars. In those areas where there is no human pressure, as in Mar Azul (south of Villa Gesell), or where sediment transport is blocked by a hydraulic barrier, as at San Clemente del Tuyú, beaches are accreting with backshore dunes (>Fig. 4.4.5). At Mar del Plata and other sites, the emplacement of sea walls, jetties and piers is responsible for erosion along the littoral drift direction.

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Dune fixation by vegetation, traditionally considered beneficial, has been claimed to be responsible for alteration of the sediment supply to the beach. A transitional zone extends for more than 100 km at Samborombón Bay, in the central part of the Salado depression and part of the Río de la Plata system. The

hinterland is a low-lying portion of the Pampas Plain, occupied by Pleistocene and Holocene coastal and marine deposits. At Samborombón Bay a brackish marsh is present behind an extensive mudflat. From Punta Piedras to Buenos Aires the Río de la Plata coast runs SE-NW and the coastal plain is up to 8 km wide. rising to about 5 m. ⊡⊡ Fig. 4.4.5 Stable or accreting beach on the NE barrier of Buenos Aires at Mar Azul, south of Villa Gesell.

⊡⊡ Fig. 4.4.6 Storm surge damage at Mar del Tuyú.

Argentina

It  formed during the mid-­Holocene maximum sea level stand around 6,000 years bp (Cavallotto et al. 2005) and the subsequent emergence. The coastal plain is composed of Holocene beach ridges, marshes and fresh-water wetlands, with a levee of estuarine sediment on the coastal fringe. It is backed by a high terrace on continental sediment, mainly loess-like silts. The inner boundary of the coastal plain is the midHolocene high sea level coastline. The Río de la Plata has an area of about 35,000 sq km and a drainage basin (Paraná and Uruguay Rivers) of more than 3 million sq km. The estuary receives large amounts of dissolved and suspended solids and organic carbon. The sediment is largely from the Paraná River, mainly as suspended load but also as bed load, and forms a delta of 18,000 sq km. The sediment in the estuary consist mainly of silty clays with some sand from the Paraná river, while the yield From the Uruguay River is sandy. Coastal plain flooding due to storm surges can be particularly destructive in areas where topographic gradients are extremely low, as on the Río de La Plata shores and in the Salado basin. Storm surges are regarded as the most important morphodynamic factor in coastal development along this area (>Fig. 4.4.6). It is widely accepted that a global sea level rise is causing impacts on coasts, and an accelerated sea level rise would certainly exacerbate coastal erosion and flooding in Argentina. The most vulnerable areas are the Río de la Plata, including the Paraná delta and the estuarine coastal plain, already exposed to frequent storm surges, and the sandy coastline extending from Mar del Plata to Punta

4. 4

Rasa, close to Samborombón Bay, where there is already severe erosion. It is likely that other areas such as Samborombón Bay and some estuaries would be affected by an increasing sea level, depending on its magnitude and the interaction with other, human and climatic variables.

References Bujalesky G (2000) Quaternary Coastal Environments of Tierra del Fuego (Argentina). Field Trip Guidebook, November 4–7. IGCP-437, IGU, INQUA, p 27 Cavallotto JL, Violante RA, Parker G (2005) Sea level fluctuations during the last 8,600 years in the La Plata river. Argentina. Quat Int 114:155–165 D’Onofrio EE, Fiore MME, Pousa JL (2008) Changes in the regime of storm surges at Buenos Aires. Argentina. J Coastal Res 24:260–265 Isla FI, Cortizo L, Schnack EJ (1996) Pleistocene and Holocene beaches and estuaries along the Southern barrier of Buenos Aires. Argentina. Quat Sci Rev 15:833–841 Lanfredi NW, Pousa JL, D’Onofrio EE (1998) Sea-level rise and related potential hazards on the Argentine coast. J Coastal Res 14(1): 47–60 Pousa JL, Kruse E, Guaraglia D, Mazzoldi A, Carbognin L, Tosi L et al (2007) Geological hazards in two sandy environments: the eastern coast of Buenos Aires (Argentina) and Vence (Italy). Env Geol 51:1307–1316 Rabassa J, Coronato A, Bujalesky G, Salemme M, Roig C, Meglioli A et al (2000) Quaternary of Tierra del Fuego, Southernmost South America: an updated review. Quat Int 68–71:217–240 Schellmann G, Radtke U (2003) Coastal terraces and Holocene sea level  changes along the Patagonian Atlantic coast. J Coastal Res 19:983–996 Schnack EJ (1985) Argentina. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 69–78

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4.5 Uruguay

1.  Introduction The coastline of Uruguay is about 600 km long, including the shores of Rio de La Plata and the Atlantic Ocean. Evolution of the coastline has been influenced by the uplift of the Ballena and Animas ranges and the Southeast and Southwest highlands, and subsidence of the structural trough that underlies the Santa Lucia River valley and Merin Lagoon basin, and lowlands occupied by Montevi­ deo Bay and East Coast lagoons and estuaries. In the eastern part of the country, the coastline truncates the geological formations of the Eastern Highlands Shield, which shows a north–south trend (N 10° E) in the folded and faulted rocks and the pattern of granite intrusions. The Pre-Devonian crystalline rocks of the Brazilian craton to the west extend to the south, in the Ingles and Rouen banks (Rio de La Plata), where they dip north into the Rio Salado syncline (Bossi 1966). North of the coastal region, the north–south trend of the geological formations of the Eastern Highlands is interrupted by a large complex of strike-slip faults, with a lateral displacement of 30 km, and a general direction of N 60° E between the tectonic basin of Merin Lagoon and the Arazati River mouth. These strike faults have led to vertical uplift of local areas between alternating areas of tension and compression. This fault zone has been filled in succession with pillow lava, rhyolitic dykes (Tertiary volcanism), and sea-floor, coastal, and continental sediment. The coast is considered to have been generally stable in late Quaternary times, indicating a higher sea level during the Holocene. Earthquakes, however, are occasionally recorded and the possibility of localized minor warping cannot be ruled out. The resultant general direction of the shoreline is northwest–southeast between Gorda Point and Colonia Point and WNW–ESE between Colonia Point and the Santa Lucia River estuary. East of Espinillo Point to Del Este Point, there is an irregular eroded rocky coast, with many bays and beaches, and a long coastal arc tending generally in a west–east direction. East of Del Este Point to Santa Maria Cape, the

general coastal direction is N 60° E, with a gradual turn to N 30° E in three sectors of the coast up to the Brazilian frontier. The variations in coastal aspect are important for coastal dynamics. East of Espinillo Point the coast is exposed to a SE ocean swell, modified by the shallow depths of Rio de La Plata and the epicontinental sea, and to storm waves mainly from the SE and SW (mouth of the Rio de La Plata). The western sector is subject only to storm waves, mainly from the SE and SW, occurring in a reduced fetch and shallow water (5 m maximum depth). Mean spring-tide ranges on the coastline are generally between 0.4 and 0.6 m, but the SE winds produce storm surges between 1.9  m (Rio de La Plata mouth), 2.7  m (Montevideo), and 3.7 m (Colonia). Cliffs are found on several geological formations and in many sections are fronted by beaches. Sandy beaches are most common, fronted by sand bars and backed by beach ridges and dunes. The sand is predominantly quartzose, but calcareous in the east and the far west. Longshore drifting on the ocean coast north of Santa Maria Cape is generally from SW to NE. There is no net drift between Santa Maria Cape and Del Este Point, but west of Del Este Point there is some westward drifting. The shore between Punta Ballena and Punta Rubia has the coarsest beach sand (1 mm); elsewhere it is finer (0.1–0.2 mm) (Jackson 1979). Most Uruguayan rivers have occasional torrential flow and drain along fault-guided valleys into estuarine lagoons, with sea entrances blocked or encumbered by inwashed sandy thresholds and flanking spits. Swampy shores, typically with Juncus and Spartina salt marshes, occupy the more sheltered parts of the coast between Montevideo and Arazati, the margins of estuaries, and the mouths of coastal lagoons.

2.  The Uruguayan Coastline The Argentine border runs N–S down the Rio Uruguay to the Rio de la Plata gulf. The Rio Uruguay widens below

Edited version of a chapter by J. M. Jackson in The World’s Coastline (1985: 79–84). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Uruguay

Fray Bentos to an estuary with numerous alluvial islands. Below Fray Bentos it is about 6 km wide, with the delta of the Rio Negro on its eastern shore, but it then narrows as it is constricted by the delta of the Parana River, coming in from the west. It widens again past Carmelo, the eastern shore passing from estuarine to marine and sandy along the Colonia coast. This coastal lowland is fringed by beaches of coarse sand brought down by the Uruguay River. Gently curved beaches between cliffed promontories of Miocene–Oligocene mudstone are backed by bluffs cut in Pleistocene sandstones and covered with dunes. The coast curves out to a granitic promontory, Punta Hornos, with outlying islands and rocks, and the port of Colonia del Sacramento. Tertiary rocks outcrop above the granite in low cliffs on Colonia Point, continuing as the coastline turns eastward, and the Rio de la Plata gulf diverges as an area of shallow fresh water some 40 km wide. The mouth of Rosario River is navigable, but there are broad sand bars offshore. To the east the beach is interrupted by headlands and shore platforms of granites with diabase dykes, then a broad lowland with fresh-water swa­ mps and dunes has developed here in front of a Pleistocene coastline, which is indicated by bluffs. There are low-lying plains with dunes and bordering cliffs in Tertiary formations near Arazati offshore sand shoals which are a source of continuing beach progradation, in contrast to the prevalent erosion in the San Gregorio sector to the SE. There is westward drifting of beach and shoal sands towards Jesus Maria Point, and the sandy beaches are backed by salt marshes with Spartina. Near the port of Juan Lacaze is the small Rio Rosario delta, and a cliffed coast runs out to Punta San Gregorio and on to Punta del Tigre. The 40 m high cliffs on Punta San Gregorio has been cut into Plio-Pleistocene limestones and sandstones, and sand derived from these has formed sand shoals that are the source of continuing beach progradation to the east, where 30 km of sandy barriers and spits formed in Pleistocene times fill an ancient bay. The Rio Santa Lucia flows by way of a small delta into a bay where beach ridge plains fringe the coast. Structurally controlled by faults, the Santa Lucia River valley borders a scarp on its eastern margin. The Santa Lucia Delta plain is developed on a swamp area, formerly a chain of coastal lagoons, that occupies a marine embayment at the southern end of the Santa Lucia tectonic depression. On the eastern side of Punta del Espinilla a succession of bays between small headlands extends to Montevideo Bay, which occupies a sunkland between the San Pedro Peninsula to the east and the Lobos Peninsula to the west.

The port of Montevideo is sheltered by large breakwaters. In the NE of the bay some inter-tidal mudflats are threatened by erosion as a consequence of dredging a navigation channel, and sandy beaches on the eastern shore (as at Capurro) were destroyed by port construction. The sand deposited in these bays came from the sea floor during and after the Late Quaternary marine transgression, but there is no longer any significant sand supply from this source. The coast to the east consists of low cliffs and coves with beaches, the largest of which is Playa Ramirez. There are steep slopes rising to Montevideo Hill, and small bays with sandy beaches are backed by partly mobile dunes. Punta Manso is the easternmost of the amphibolite and pegmatite headlands characterising the rocky coast of Montevideo. Brava Point also consists of Pre-Devonian rocks and amphibolite and pegmatite dykes dissected by marine erosion into irregular shore topography, and embayments fringed by Spartina and Juncus marshes. On either side of Brava Point, valley mouths end in sandy coves backed by dunes. East of Montevideo the southerly fetch over the South Atlantic Ocean lengthens, so that wave energy increases. Seaside resorts continue to Atlantida and round the bay to Piriapolis. There is a large beach of fine Pliocene and PlioPleistocene rocks backed by mobile quartzose sand dunes (PCYMP 1979) between Punta Manso and Punta Negro. The Animas Range, of Pre-Cambrian igneous rocks, runs south to the coast to form the promontory out to Punta Negra. The coastline has a jigsaw configuration resulting from the diagonal trend of the faults (strike-slip and oblique slip faulting) leading to movement of the coastal blocks. To the east are landslips in Tertiary formations that rise toward the Animas Range. This formation sometimes has steep coastal slopes with only limited basal cuffing behind shore platforms. The shore platforms are cut in hard Paleogene sandstones containing pebbles and cobbles, which are was­ ­hed out to form beaches in the coves. Behind the narrow Solis Grange River entrance, impeded by looped sand-and-pebble bars and a spit that has grown from west to east, there is an estuarine lagoon surrounded by salt marshes. East of this mouth the beach is backed by limestones and peaty deposits capped with dunes. The peats were formed in an estuarine ­environment. Seaward, these Holocene sediment have been cut back into steep cliffs (4–6 m) and serrated shore platforms. The same pattern characterises all the river mouths in this coastal sector. Punta Negra is formed by three syenitic ridges with diabase dykes, bordered seaward by extensive rocky

Uruguay

­ utcrops. The coves between the three ridges have beaches o of coarse sand. East of Punta Negra is an asymmetrical sandy bay with a beach of fine sand backed by dunes, which have drifted northeastward to the Laguna del Sauce, a fresh water lake that occupies a tectonic depression and is linked to the sea by a narrow channel. Ballena Point is a promontory of quartz sandstone, dipping 80 west and ending in a rocky slope that plunges into deep water, and another asymmetrical bay curves out to Punta del Este, south of Maldonado. It has a permanent natural outlet opening over a granite substratum, but this is not navigable due to a perennially recurring looped sand bar offshore. Offshore is a high island, Isla de Lobos, with a lighthouse. Punta del Este is a breached sandy tombolo in the lee of a granitic island. Settlers built a road connecting the island with the mainland around 1,900, making this a permanent isthmus. Bluffs back the tombolo, which is covered by mobile dunes. Beyond Punta del Este the coastline trends ENE following the trend of schist outcrops, and sandy beaches receive strong southeasterly ocean swell. The beach has a 10° slope and consists of coarse quartzose sand derived from weathered granite. Maldonado Inlet is a broad, shallow estuarine lagoon with extensive sandy shoals exposed at low tide, and shores that are fringed by salt marshes. There has been extensive sand deposition, with dunes surmounting barrier formations that contain Laguna José Ignacio, Laguna Garzon and Laguna de Rocha, which has a swampy deltaic northern shore. The three lagoons occupy former marine embayments in tectonic depressions, enclosed by sand barriers in Pleistocene times. The barriers were dissected during the last glacial low-sealevel phase, before the Holocene outer barrier was added to en­­close the present lagoons. The lagoons have intermittent natural outlets. There is a granite promontory and some minor granite headlands that interrupt a 75-km-long beach backed by bluffs cut into Pleistocene coarse-grained calcarenite. The barrier bears dune ridges parallel to the coastline, and parabolic dunes and blowouts advancing SW–NE. Cabo Santa Maria is a complex tombolo formed in the lee of schist outcrops and is covered by mobile dunes, which have spilled northeastward. Bluffs back the tombolo. The coastline then trends NE past La Paloma in a series of long asymmetrical bays, shaped by refracted ocean swell, with beaches backed by extensive dune ridges. The beach is in­­ ter­­rupted by outcrops of chloritic–schist phyllite running N 25° E and dipping 80° north, which form the Punta Rubia promontory.

4. 5

Rivers converge into Laguna de Castillos, which has a 22 km long meandering outlet channel through the dunes to the coast north of Punta Aguda, where granite outcrops on the shore. The lagoon is bordered by emerged Holocene beaches (2–4 m above sea level) and fresh-water swamps. Brackish water flows in through the channel from time to time. South of this beach the granitic headlands of Polonio are covered by mudstone and peat, overlain by fixed dunes composed of coarse quartzose sand capped by a recent longitudinal mobile fine dune. The deposits are up to 60 m thick above the granite. In places dune sand is spilling down on to the shore. Recent dunes also found around swales between the Pleistocene calcarenite ridges which back a 25 km long, gently curved beach. These form a barrier impounding Laguna Negra, the dunes continuing behind low cliffs and rocky shores on the granite outcrop at Punta Palmar behind the Coronilla islands. This is the Santa Teresa Granite, part of an intrusive ring pattern. Granitic headlands separate gently curved arcuate beaches of quartzose sand backed by dunes and bluffs of Pleistocene dune calcarenite. In the lee of the Coronilla Islands, wave refraction has shaped a major sandy breached tombolo. The SE Santa Teresa coast has a long sandy beach with scattered gravel. The long dune ­calcarenite ridge has been much dissected by lavakas (­gullies). Runoff playas are transformed into mobile long­­ itudinal dunes of fine sand by northerly and southwesterly winds. The gently curving sandy Chuy Beach is backed by extensive, partly mobile, quartzose-calcareous dunes, and two long Pleistocene calcarenite ridges cut by gullies and lavakas. Between the ridges are fresh-water lakes and associated swamps, underlain by rhyolite dykes and Santa Teresa granite. The ridges form a barrier on the southern side of the Merin-Patos Lagoon, which extends into Rio Grande do Sul, and the beach runs along the Brazilian border at Barra de Chuy.

References Bossi J (1966) Geologia del Uruguay, Departamento de publicaciones de la Universidad, Uruguay Jackson J (1979) Sedimentacion reciente en la costa platense-atlantica del Uruguay al este de Montevideo, Informe Ministerio de Obras Pitblicas J Uruguay

229

4.6  Brazil

Dieter Muehe

1.  Introduction The Brazilian coastline (> Fig.  4.6.1) lies between latitudes 4°N and 32°S and is about 7,400 km long, although various estimates range between 6,000 km and more than 10,000 km (Muehe 2003). It is backed by extensive continental plateaux of igneous and metamorphic PreCambrian rocks with basins of Palaeozoic sediment, and fronted by a continental shelf that widens in the North, in front of the Amazon. Sedimentary terraces, pediments of Tertiary age, of varied composition (Barreiras Group), extend from the North to the Southeast regions and represent an important element of the coastal landscape. The flat topped surface was deeply incised during Pleistocene low sea levels while the distal portions of the deposits were eroded by the sea during marine transgressions, leaving a coastline of fossil and active bluffs several metres high, preceded by narrow beaches. Rocky sectors are found mainly in the Southeast region, in places where the Pre-Cambrian basement structures of the Serra do Mar mountain range come to the coastline, as in the area around Ilha Grande bay (Rio de Janeiro) up to Caraguatatuba and São Sebastião island (São Paulo). Mangroves grow alongside the estuaries and are highly developed in the North region, between French Guiana and the Maranhense Gulf, comprising the States of Amapá, Pará and particularly Maranhão, which has about half of the area of Brazil’s mangroves. Large dune fields are found in the Northeast region where the climate is dry, and also in the Southeast and South regions in places with strong winds and a supply of fine sand. There is wide development of dunes along the coast of Rio Grande do Sul, in front of the Patos lagoon, as also locally in Espirito Santo, Rio de Janeiro and Santa Catarina States. > Fig.  4.6.1 is a general geomor­ phological chara­cterisation based on classifications of Silveira (1964); Dominiguez (2004) and Muehe (2006).

The climate varies from warm temperate in the south, where the winters are cool and rainy, to humid tropical in the north. Porto Alegre, in the south, has a mean monthly temperatures ranging from 14.3°C to 24.1°C, with precipitation well distributed through the year, amounting annually to about 1,400 mm. In Rio de Janeiro in the southeast the mean temperature of the coldest month is 20.6°C in July, rising to 25.6°C in January, and there is an average annual rainfall of 1,082 mm with a relatively dry winter. Recife, in the northeast, has a mean annual temperature of 25.4°C with monthly mean temperatures varying less than 2°C around this mean. Annual mean precipitation is slightly less than 2,500 mm with major rainfall between April and July. Belém in the north has a mean annual temperature near 26°C with little variation through the year. Annual rainfall is 2,820 mm with a wet season from January to May, when mean monthly rainfall is above 300 mm and a less moist period during the rest of the year with monthly rainfall between 100 mm and 200 mm. The west-facing coastline (north west of Cabo Cal­ canhar) receives waves generated by east and southeast trade winds, but wave energy is low in the sector of the Amazon and Pará estuaries, where the continental shelf is wide and shallow. Farther south, waves generated by the trade winds are effective, and they are accompanied by southerly swell arriving from the South Atlantic. Wave energy increases in the stormier environments of the far south. Tide ranges are very small (less than 2 m) in the south, southeast and part of the northeast, up to the state of Alagoas. Further north tide ranges of up to 12 m and tidal currents of up to 250 cm/s have been recorded in the Amazon estuary. From south to north mean spring tide ranges are: Rio Grande 0.3 m, Imbituba 0.7 m, Santos 1.1 m, Rio de Janeiro 1.1 m, Cabo Frio 0.9 m, Salvador 2.2 m, Maceió 2.1 m, Recife 2.1 m, Natal 2.3 m, Mucuripe 2.5 m, São Luis 5.2 m, Belém 2.2 m.

This is an abridged version of a chapter by O. Cruz; P.N. Coutinho; G.M. Duarte; A. Gomes and D. Muehe which was published in The World’s Coastline (edited by E.C.F. Brid and M.L. Schwartz) in 1985 pp. 85–91.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 4.6.1 The coastline of Brazil.

2.  The Brazilian Coastline From the Uruguay border the coast runs northeast (> Fig. 4.6.1A) and consists of wide sandy barrier formations enclosing major lagoon systems, Lagoa Mirim and Lagoa dos Patos, with associated salt marshes and sedge swamps (Delaney 1965). The coastal plain of Rio Grande do Sul is up to 120 km wide, and notable for extensive dunes driven by predominantly northeasterly winds. The coastal barrier fronting the Lagoa dos Patos has dunes up to 25 m high in a 10 km wide zone along its seaward fringe, disrupted by blowouts that have moved from NE to SW. This is backed by a dune ridge that rises to 35 m, and has been dissected, yielding sand for the cuspate forelands between excavated bays on the shores of Lagoa dos Patos. The sandy ocean beach extends almost continuously for 750 km from the southern border up to cape Santa Marta. North from the Patos lagoon to Cape Santa Marta, the scarped edges of the sedimentary and basaltic Paraná basin develop behind a narrow coastal plain with a sequence of lagoons dammed by the barrier beach.

North from Torres to Santa Catarina Island the PreCambrian basement comes to the coast with high promontories separated by valley mouth inlets formed by Holocene marine submergence. The hilly island of Santa Catarina is noted for its beach ridge and dune systems. The southernmost limit of mangroves in Brazil is at Lagu­­­na (28°30'S), with Avicennia ­schaueriana, Laguncula ria racemosa and Rhizophora mangle (Schaeffer-Novelli, Personal communication). North from Santa Catarina island the coast becomes simpler in outline, interrupted by inlets at the mouths of the steep valleys on the eastern flank of the Serra do Mar or by large bays such as Para­ naguá. Pleistocene and Holocene barriers enclose narrow lagoons and mangrove swamps at Cananéia, north of the hilly Ilha do Cardoso (Bigarella 1978) (> Fig. 4.6.2). Stages in the evolution of these depositional features have been traced and dated by Suguio and Martin (1978). The tide range is less than 1.5 m, but behind Ilha Comprida intertidal mangroves are extensive in the lagoonal channel system of Cananéia-Iguape, including the Ribeira de Iguape river mouth. Sandy beaches towards Santos are interrupted by steep-sided residual coastal massifs.

Brazil

4.6

⊡⊡ Fig. 4.6.1A The coastline of southern Brazil. (Courtesy Geostudies.)

IA - TRIN VITÓR DA

R PA

RIB IGUAPE R.

SANTOS

T EN YM BA M E CARDOSO I. PARANAGUÁ BAY

SÃO

PAULO

PLATEAU

Scarped and rocky coast Beach ridges Islands

O

SL

SH

EL

F

SANTA MARTA C.

PE

S.FRANCISCO I. STA.CATARINA I.

Beach ridges - spits Lagoons

TORRES JACUÍ R

N

TI N O

MIRIM L.

C

I

O

UA

SCALE OF KILOMETERS 0

100

200

350

C

UG

TI

N

UR

THE BRAZILIAN COAST

N

EN

EN

TA

L

PATOS LAKE

TA

L

Beach ridges Channel system Mangrove Beach ridges Mangrove Boys

RID GE

CABO FRIO GUANABARA BAY ILHA GRANDE BAY CARAGUATATUBA BAY de S. SEBASTIÃO BAIXADA SANTISTA

TROPIC OF CAPRICORN

Mossifs Beach ridges - delta Barreiros Form.

DE

O SUL R. AD AÍB

Basement oblique ridges Islands Beach ridges - dunes Mangrove Beach ridges lagoons salt marshes Dunes State boundary

50° Des.-Hélio

Santos stands on a peninsula between broad tidal inlets fringed by extensive marshes and mangrove swamps. A curving sandy beach faces southward to the Baia de Santos, between high, steep-sided promontories. At the western end of the beach is a high island attached to the mainland by a sandy tombolo. Sandy beaches are interrupted by rocky promontories that become larger and more frequent as the front of the mountain range of Serra do Mar approaches the coastline. A mountainous promontory runs out toward the high, steep-sided Ilha de São Sebastião, which stands beyond a deep, narrow strait, and the high hinterland dominates the coast between here and Rio de Janeiro. Embayments, some well sheltered from prevailing southeasterly wave action, have beaches and mangrove swamps. Caraguatatuba Bay faces east and is backed by a Pleistocene and Holocene beach ridge plain. The beach prograded following the 1967 downpour in the high hinterland, when catastrophic erosion caused rivers to discharge vast quantities of sediment into this bay (Cruz 1974). Smaller bays with beach ridges are found between steep irregular headlands on the coast as far as Ubatuba, and in the western part of the Baia da Ilha Grande

s­ heltered bays are backed by mangrove swamps, as at the historic town of Parati on the Rio Perequeaçu delta. The indented steep coastline of this bay has rocky headlands (> Fig. 4.6.3) fringed by numerous islands. There are steep, arcuate beaches in small coves, and longer, flatter less regular beaches in protected bays, as in the eastern part of Bahia da Ilha Grande. East of Baia de Ilha Grande the first of a series of large beach barriers, the Restinga da Marambaia, encloses Sepetiba bay. These barriers are in the form of long and straight sandy deposits up to Cabo Frio, interrupted by steep promontories of crystalline rock (> Fig. 4.6.4). Copacabana Beach (> Fig.  4.6.5) stands in front of a heavily urbanized relatively narrow coastal plain, and was artificially nourished in the early 1960s (Muehe and Neves 1995). The Sugar Loaf stands at the western entrance of Baia de Guanabara. East of Rio de Janeiro, the irregular coastline of the coastal range has been rectified by long and straight beach barriers which enclose a series of coastal lagoons, the largest of which is the Lagoa de Araruama, close to Cabo Frio. This is backed by sedimentary formations of the Barreiras Group and crystalline

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⊡⊡ Fig. 4.6.1B The coastline of Northern Brazil. (Courtesy Geostudies.)

Brazil

4.6

⊡⊡ Fig. 4.6.2 The mountainous coast south of Cananéia. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.3 Rocky headland near Bertioga, showing intertidal oyster zone. (Courtesy Geostudies.)

rocks. There is generally an inner barrier, related to the latest Pleistocene transgression, and an outer barrier developed during the Holocene sea level rise. Muehe and Corrêa (1989) have shown that the barriers migrated to his present position in association with a sea level rise. Lagoa de Araruama is segmented by large spits formed of sediment delivered from the erosive recession of the back of the Pleistocene barrier beach (> Fig. 4.6.6).

Cabo Frio is a dry and relatively cool region, due to the upwelling of cold water by the strong northeastern trade winds. The high evaporation is used for extensive salt production from pans excavated in marshlands or impounded in shallow lagoon areas. The hilly Cabo Frio peninsula is attached to the mainland by a broad sandy isthmus bordered by curving beaches and surmounted by large dunes (> Fig. 4.6.7).

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⊡⊡ Fig. 4.6.4 Rocky promontory at Saquarema, with 1630 cathedral. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.5 Copacabana Beach, Rio de Janeiro. (Courtesy Geostudies.)

North from Cabo Frio the first major feature is the deltaic protrusion at the mouth of the Paraiba do Sul river, formed largely by Holocene beach ridge progradation. The southern part, near Macaé, truncates widely spaced, subdued Pleistocene beach ridges. The large mountain range of Serra do Mar recedes inland and is breached by several river valleys that open on to the depositional coastal plain. Locally, outlying hills reach the coast alongside deep estuaries, as at Vitória (Espirito Santo). Farther

north the large coastal plain of the Doce river is formed of Pleistocene and Holocene terraces and a fringe of beach ridges. The coastal plain widens northward into Bahia (> Fig. 4.6.1B), and includes deltaic regions at the mouths of several rivers, notably the Jequitinhonha, which has a delta at Belmonte flanked by prograded beach ridge plains. A wide beach ridge system marks stages in the growth of the cuspate foreland at Ponta da Baleia in the lee of the

Brazil

4.6

⊡⊡ Fig. 4.6.6 The Punta do Acaira spit, projecting into Lagoa de Araruama. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.7 Blowout in the dunes on Cabo Frio. (Courtesy Geostudies.)

Abrolhos archipelago. The coastal plain then narrows past the mouth of the Rio Alcobaça. Long straight beaches and narrow barriers extend north to Ilhéus on a small delta at the mouth of the Rio Cachoeira. Farther north, transverse faulting has complicated coastal topography north of Itacaré, where the coastline becomes indented by estuaries and embayments, the largest of which is the Todos os Santos Bay, behind the hilly peninsula of Salvador. In an increasingly warm and wet climate, the sheltered shores of the estuaries are extensively mangrove fringed, while the

consistent trade winds produce the waves in coastal waters that have built up beach ridge plains. North from Salvador the relatively straight coast line is bordered offshore by beach rock and calcareous reef sandstones (Bigarella 1975), some of which are cemented dune sands (aeolian calcarenites). There are coastal terraces of 2–3 m and 7–8 m above present sea level. Locally there are cliffs cut into Pliocene sedimentary formations of the Barreiras Group. North of Aracajú the coastline curves out to beach ridge fringed plains around the mouth of the

237

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São Francisco river; again, there are relics of a Pleistocene delta fronted by prograded Holocene beach ridges. Humid tropical conditions, with annual rainfall between 1,000 and 2,000 mm, dominate the coastline northward, up to Recife and Natal, but with a lengthening summer dry season. The coastline remains simple, with beach ridge plains and nearshore sandstone reefs. Cape Calcanhar marks a major change in the coastline toward a north-­ facing equatorial sector. The coastline is generally smooth and depositional, beach ridges and dunes alternating with lagoons, swampy sectors and salt deposits. The dry season becomes pronounced, and vegetation is sparse in the hinterland. Near the Parnaiba delta are long sandy beaches backed by dunes. Strong onshore winds, dry climate and fine sand have formed the wide dune fields along the coast of Maranhão, west from the Maranhense gulf to Cape Calcanhar, including the whole coast of Ceará and part of the state of Rio Grande do Norte. The Lençóis Maranhenses, a wide dune field located between the Maranhense gulf and the Parnaíba delta, extends along a 20 km coastline up to 10 km inland. Large dune fields also are found along the coast of Ceará as for instance in Jericoacoara where exceptionally huge dunes are developed. Tide range increases westward along this coastline, and tidal currents are strong in the bays of Marcos and São José (Maranhense Gulf) at São Luis, the capital of Maranhão. The coastline in this equatorial sector is indented, with many inlets and estuaries bordered by mangrove swamps and backed by a low plateau of the Tertiary Barreiras Group and early Quaternary sedimentary rocks, which reach the coast in several localities as small cliffs. West from the Maranhense Gulf, the branching funnel-shaped estuary of the Marajó bay near Belém, has strong tides with intricate and variable shoal and channel topography. Numerous mangrove-fringed alluvial islands, and swampy shores are subject to rapid changes in configuration due to interactions of waves and of tidal and fluvial currents. It is separated from the similar Amazon estuary by the large island of Marajó, notable for its extensive swamps and intersecting tidal channel systems. The vast discharge of water and sediment (sand, silt and clay) from the Amazon results in

rapid accretion in and around the mouth of the river and along the adjacent coastline, especially northward into French Guiana as the result of tidal movements and the Guiana current. The low lying coast of the Amazon deltaic region consists of extensive swamps with mangroves and tropical forest, sandy beaches and cheniers, alluvial grasslands, lagoons and numerous channels. The climate is perennially hot and wet, with an annual rainfall of more than 2,000 mm, and vegetation is luxuriant. Coastal waters are turbid due to the high sediment suspension. The low lying swampy coastline continues north to Cabo Orange at the border between Brazil and French Guiana.

References Bigarella JJ (1975) Reef sandstones from Northeastern Brazil. A Ac Br Ciên 47:395–409 Bigarella JJ (1978) The Serra do Mar mountain range in the Eastern part of the State of Paraná, SEPL/ADEA, Curitiba, 248p Cruz O (1974) The Serra do Mar mountain range and the coast of Caraguatatuba (in Portuguese). Dissertation, Instituto de Geografia – University of Sao Paulo, 11. Sao Paulo Delaney PJV (1965) Physiography and surface geology of the coastal plain of Rio Grande do Sul (in Portuguese). Escola de Geologia. UFRGS, Special Edition, Porto Alegre 6, 105p Dominiguez JML (2004) The coastal zone of Brazil: an overview. J Coastal Res (Special Issue) 39:16–20 Muehe D (2003) Beach morphodynamic research in Brazil: evolution and applicability. J Coast Res (Special Issue) 35:32–42 Muehe D (2006) Erosion in the Brazilian coastal zone: an overview. J Coastal Res (Special Issue) 39:43–48 Muehe D, Corrêa CHT (1989) Beach morphodynamic and sediment transport of Massambaba beach barrier (in Portuguese). Revista Brasileira de Geociências 19(3):387–392 Muehe D, Neves C (1995) The implication of sea level rise on the Brazilian coast: a preliminary assessment. J Coast Res (Special Issue) 14:54–78 Silveira JD da (1964) Morfologia do litoral. In: AZEVEDO A de (ed) Brasil a Terra e o Homem. Companhia Editora Nacional, São Paulo, Brasil, pp 253–305 Suguio K, Martin L (1978) Quaternary marine formations of the State of São Paulo and southern Rio de Janeiro. International Symposium on Coastal Evolution in the Quaternary, I.G.C.P. Project 61, Special Publication 1, Sao Paulo Tricart J (1960) Geomorphological problems of the east coast of Brazil (in Portuguese). Bol Baiano Geogr 1:1–39

4.7 French Guiana

J. Turenne

1.  Introduction The coastline of French Guiana extends along the Atlantic Ocean between Brazil and Surinam for 370 km. It is a landscape of mangroves and salt marshes, broken here and there by outcrops of Pre-Cambrian crystalline basement (>Fig. 4.7.1). Outcrops of the Guianese Shield stand out from the coastline, and form islands in the open sea. They include the granitic islands Le Grand and Le Petit Connetable (the High and the Low Constable, 04° 50' N; 51°56' W) near the coast; Le Père (the Father, 100 m in altitude); La Mère (the Mother, rising to 111 m); L’Enfant Perdu (The Lost Child); doleritic Le Malingre (the Sickly One, rising to 55 m) and Les Iles du Salut (the Salvation Islands, 66 m) consisting of gabbros and peridotites. They also form reefs buffeted by the waves (Malmanoury Reefs). Other rocky outcrops rise from the coastal plain, which consists of marine sediment of Holocene (mainly marine clay) or Pleistocene age (fine, well-sorted sands derived from old submerged barrier sands). These include Ouanary Mountain, Montagne d’Argent (Silver Mountain, 100 m), the Cayenne Mountains (234 m, consisting of migmatites and diorites), and the granite-gneiss hills of Kourou and Organabo. The coastal region of French Guiana is a part of the large sedimentary area found in all three Guianas, and attains its maximum width in the Berbice Region (Brink­ man and Pons 1968). Clay deposits extend from the mouth of the Amazon to the mouth of the Orinoco, except for the sector between Cayenne and Organabo, where the PreCambrian basement protrudes into the sea. This tectonic partitioning, along with a slight uplift of the Pre-Cambrian basement, has influenced the morphology of the coast, the deposition of clay and the formation of sand ridges near the crystalline outcrops. There is an outer (younger) coastal plain less than 4 m above sea level, backed by an inner (older) coastal plain standing about 15 m above sea level. Sediment of the outer coastal plain are 60% clay (mainly kaolinite and some montmorillonite) and 20% fine silt, and are of Holocene age, having formed since about 8,000 years bp. They are of marine ­origin, and

include ­sediment dominated by kaolinite, which is delivered to the coast by the Amazon River and drifted alongshore to the Guianas (Dolique and Anthony 2005). The flat clay plain is varied by cheniers, ridges of coarse sand, generally oriented SE–NW, often emerging from the middle of the marine clay plain. The inner coastal plain appears to be a Late Pleistocene terrace, and consists of ridges of fine well-sorted sand, covered with savanna and forest vegetation. The Brazilian border follows the estuary of the Oyapock River, which opens to a broad gulf (Baie d’Oyapock) besides Cabo Orange, and the Approuague River opens to a similar estuary about 25 km to the west, beside Pointer Béhague. Both estuaries swing NNW, in res­­ponse to westward longshore drifting. To the west, the coastal plain is crossed by several rivers, including in westward succession the Oyac, Cayenne, Kourou, Sinnamary, Counamama, Iracoubo and Mana, all of which have small estuaries deflected to the west. The Surinam border follows the larger Maroni estuary. The coastline has prograded during the Holocene, and the river channels, banks and remnants of former river channels diversify the relief of the coastal plain. Anastomosed channels are typical of the estuaries, particularly in the Cayenne Island region, which is surrounded by such channels, as well between the Mana and Maroni rivers. The channels are vestiges of a Late Pleistocene river system, which was progressively reduced by Holocene marine sedimentation. Strong tidal currents keep the channels navigable. The remnants of a Pleistocene coastline can be observed in this region (Boyé 1963). The average range of spring tides is 3 m; that of neap tides 2.5 m. The tide wave rises through the estuaries to the waterfalls or rapids, which mark the first outcrops of the Pre-Cambrian basement. The ocean current off of the Guianas, also called the Great Southern-Equatorial Current, moves SE–NW, at between 2 and 3 knots, with a branch that runs along the coast, where sea water mingles with fresh water from the large rivers. The continental shelf, carpeted with mud and fine sand, has only a gentle slope, 0.05–1.0%; it is 92 km wide off Cabo Orange in the east, increasing to a maximum of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 4.7.1 The coast of French Guiana. A, salt marshes; B, mangroves; C, swamps.

147 km off Pointe Isère on the Surinam border, and the continental slope then drops rapidly from over 1000 m. The coastline is subject to alternations of erosion and deposition, associated with the northwestward migration of nearshore mudbanks. In a study of the port of Demerara, the Delft Hydraulics Laboratory (1962) showed that the mudbanks can form at any point on the Guianese coast as a result of refraction of incident waves, with convergence of wave energy producing zones of maximum disturbance and erosion, while divergence generates calm zones of minimum disturbance and mud deposition. The predominant

westward current along the coast drives these mudbanks from east to west, so that there is erosion and deposition as they pass along the coastline. They can be seen on satellite photographs (>Fig. 4.7.2). There is a theoretical mudbank length of 40 km, moving at 1.3 km/year, so that a mudbank arrives at 30 year intervals on any part on the coastline. Irregularities are caused by rocky outcrops on or near the coast (Turenne 1978). As the mudbanks move on, increasing wave energy brings erosion, as on the Cayenne Peninsula, a phase of erosion starting in 1963 has removed the pre-existing mangrove fringe. Landsat photographs

French Guiana  

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⊡⊡ Fig. 4.7.2 Occurrence of mudbanks and turbid waters between Cayenne and Organabo, French Guiana: 1, erosion; 2, occurrence of turbid waters and mudbanks. Based on a comparison of 1973 and 1976 Landsat satellite photographs. N

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from 1976 show that sea waters are very muddy in the Cayenne region. The alternations of erosion and deposition is accompanied by gradual progradation of the coastline. As it advances mangroves spread seaward: a pioneer vegeta­­ tion  of Laguncularia, followed by Avicennia nitida, then Rhizophora to landward and on the banks of the estuaries. Villages such as Iracoubo, Tonate and Trou Poisson are now found inside a relatively wide fringe of mangroves. Since 1956, progradation has combined the outlet from the Iracoubo and Counamama Rivers into a single channel, and muddy deposition has greatly reduced the depth of access channels to the main ports. Cheniers are extensive on the coastal plain. They are sub-parallel ridges of coarse sand 70–200 m wide, deposited

successively by wave action on a plain of marine clay, and separated by marshy depressions, particularly on the western side of estuaries.

2.  The Coastline of French Guiana From Cap Orange (04° 24' N; 51° 34' W) to the Cayenne Peninsula (4° 50' N; 52° 22' W) the low-lying coast is bordered by mangroves. West of the Oyapock estuary the coast curves out to the Montagne d’Argent Peninsula, then mangroves are extensive to Pointe Béhague. Beyond the Approague estuary, the coast is again swampy as far as Mahury Cayenne. Island is an upland rising to more than 230 m, its north and northeast coasts cut by bays between headlands

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prolonged by rocks and islets. From Cayenne to Organabo, the coast is again low and fringed with mangroves. Low waves break over the migrating mudbanks. Locally there are sand ridges on the seaward edge of mangroves or salt marshes. At Kourou, the low coast is broken by Pointe des Roches (05° 09' N; 52°38' W) and offshore are Les Iles du Salut (05° 17' N; 52° 35' W). Further west, the coast between Organabo and Pointe Isère is fringed by salt marshes, threaded by cheniers and interrupted by tidal creeks. There are intermittent and variable sandy beaches along the shore, which provide the largest egg-laying area in the world for the luth turtle.

References Boyé M (1963) La Géologie des plaines basses, entre Organabo et le Maroni (Guyane Française). Imprimerie Nationale, Paris Brinkman R, Pons CJ (1968) A pedogeomorphological classification and Map of the Holocene sediments in the coastal plain of the three Guianas. Soil Survey Paper No. 4. Soil Survey Institute, Wageningen, the Netherlands Delft Hydraulics Laboratory (1962) Demerara coastal investigation. Delft, the Netherlands Dolique F, Anthony EJ (2005) Short-term profile changes of sandy pocket beaches affected by Amazon-derived mud, Cayenne, French Guiana. J Coastal Res 21:1195–1202 Turenne JF (1978) Sédimentologie de la Plaine Côtière, Atlas des DOM, La Guyane. ORSTOM-CNRS, Paris

4.8  Surinam     Norbert Psuty     

1.  Introduction Coastal Surinam extends for 386 km from the Marowijne River in the east to the Corantijn (Corentyne) River on the western border. The low-lying coast has been formed by waves and currents acting on the large quantities of finegrained sediment discharged by the Amazon River and the sand, silt, and clay discharged by the principal rivers (Marowijne, Suriname-Saramacca, Coppename) reaching the Surinam coast (Vann 1959). Coastal processes include waves that usually arrive from the northeast or east, producing a westward longshore drift. The tides are semidiurnal with a mean spring tide range of 2.8–3.2 m. (2.1 m at Paramaribo). The westward longshore drift is important in transporting silt, clay, and sand. Vast quantities of clay are distributed NW by the general oceanic circulation and by waves in the nearshore zone. There is some sand transport along the beach where the swash eventually strikes the coastline. In areas where sand and shell particles accumulate, coarser sediment is transported primarily by beach drifting. The Holocene coastal plain has been built up as a wedge thickening seaward during the Late Quaternary marine transgression. This recent accumulation has a subaerial extent of about 20 km at the eastern margin and gradually widens to a maximum of 140 km at the western boundary. The coastal deposits consist largely of clay with some lenses of sand and shells. On the surface, the general topography is one of low relief with several series of ­chenier ridges trending east to west and separated by wide swampy swales in which peat and mud have accumulated. The cheniers are found in groupings that extend fanshaped west from the major river mouths. Each of the rivers discharges into a narrow estuary, and there is evidence that the cheniers are composed of sand and shell derived locally from streams and the reworking of nearby ­mudbanks (Geijskes 1952). A single discontinuous chenier ridge extends along much of the present coastline, related to active beachforming processes, the availability of sandy sediment, and

the presence of wide mudbanks offshore. The cheniers are low (1.5–2.0 m), with an occasional small dune cap or hummock. In general there is a westward diminution in the volume of the sandy deposits and the number, width and height of the cheniers. Sandy beaches are present primarily in the eastern sector, where accumulations of coarser sediment are locally available, probably from the erosion of older cheniers.

2.  The Surinam Coastline West from the mouth of the Marowijne River the low swampy coast is fringed by a beach about 10 km long and 75 m wide. It has low dune hummocks approaching 0.5 m in elevation and a wide sloping sandy washover extending into mangroves on its landward margin. The Cottica River flows westward behind the swamps to join the Suriname estuary at Nieuw Amsterdam. Extensive areas have been embanked by Dutch engineers and reclaimed as agricultural polders. The coast is low and swampy west to Braams Punt, a small foreland bordering the Suriname estuary. The Suriname is joined by the Cottica and the Saramacca near the port city of Paramaibo and to the west the coastline is set back in a broad swampy lobe to the estuary of the Coppename River. To the west the coast is again set back in the province of Coronie, and has a swampy fringe, extending past the small townships of Totness and Friendship. The Nickerie River turns westward to join the Corantijn River near the port of Nieuw Nickerie.

References Geijskes D (1952) On the structure and origin of the sand ridges in the coastal zone of Surinam. Tijdschrift Koninklijk Nederlands Aardrijkskundig Genootschap 69:225–238 Vann J (1959) The geomorphology of the Guiana coast. In: Proceedings of the second coastal geography conference, Baton Rouge, Office of Naval Research Washington, DC, pp 153–187

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4.9 Guyana

Eric Bird

1.  Introduction The coast of Guyana (>Fig. 4.9.1) is about 434 km long and consists of mangroves, mudflats, cheniers, and sandy beaches fronting a swampy coastal plain. The coastal plain is dominated by the Demerara Formation (Holocene), which borders the entire Atlantic Ocean shore along this portion of the Guyana Shield. It varies from 25 to 35 km in  width and, where cleared of mangrove and drained as ­polders, is the site of extensive rice and sugar cultivation. Rivers and streams are tidal across the coastal plain. Farther inland tropical rain forest is prevalent. There are two highly variable rainy seasons, April– August and November–January, with an annual average rainfall at the coast of 2,250 mm (Brooks 1982). Humidity is high, typically 80–90% in the morning and 70–80% in the afternoon, and the average temperature 26.9°C. The continental shelf is wide, and nearshore waters shallow. Tides are semidiurnal, with a spring tide range of 2.6 m at Georgetown but less than 2 m along the open coast (1.2 m at Waini Point in the northwest). The region is in an east coast swell environment with trade monsoon influences but located south of the hurricane belt (Smith 1962). For the greater part of its length the coastline consists of mudflats, with abundant mangroves, cheniers, and occa­sional sandy beaches (>Fig.  4.9.2). Sediment from the Amazon River forms migrating mudflats and associated subtidal mudbanks (Wells and Coleman 1981). The several rivers bring some quartzose sand and large quantities of silt and clay to the coast. Waves reworking the sediment concentrate the sand fraction near high tide level, and as net sediment transport is to the west there is deposition of sandy material in that direction. Sandy beaches drift along the high tide shoreline, the quartzose fluvial sand being accompanied by shelly gravel and carbonate sand derived from shelly organisms that inhabit the intertidal and nearshore mud. Waves deposit sandy material as cheniers overlying a  muddy backshore substrate. Under the westward drift regime, the cheniers tend to form on the western side of  river mouths and trail off in a westerly direction

(Psuty  and Mizobe 2005). Their seaward margins form beaches of shells and sand, fronted by tidal mudflats, but mangroves may colonise these and form strips of swamp between successively deposited cheniers. The mangroves are of the Occidental Province, and thrive on tidal mudflats bordering estuaries and along the  coast. Avicennia is often found growing seaward of Rhizophora, an unusual pattern in the Caribbean region, probably resulting from the abundance of sand along the outer shore.

2.  The Guyana Coastline The precise location of the Surinam border along the Corentyne River estuary has still to be agreed. If the high tide line is taken as the border there is the problem of whether the mangroves and intertidal zone bordering the estuary belong to Guyana or Surinam, and where the boundary runs north across the intertidal zone. There is  also the question of the ownership of shoal areas that are partly emerged at low tide, and also migrating westward. From the Corentyne estuary the low-lying coast runs northwest to New Amsterdam. The sandy coastal fringe has a string of villages, backed by drained farmland, the  Black Bush Polder. New Amsterdam stands beside the  Berbice River estuary. The coast continues past Fort Wellington, the sandy fringe interrupted by inlets at the mouths of the Abaret, Mahaicony and Mahaica Rivers, and the polders extend to Georgetown, which stands below sea level, protected by embankments. A sea wall and promenade line the coast at Georgetown on the eastern side of the Demerara River estuary. The low-lying coast runs on to the Essequibo River estuary, the funnel-shaped mouth of which contains several alluvial islands. The river diverges on either side of Hog Island, and is then split by Wakenaam and Leguan Islands. To the north Tiger Island has the form of a barrier island running parallel to the coastline and backed by a narrow strait. The mouth of Pomeroon River has been deflected northwestward behind a sandy spit through which the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4.9    

Guyana

⊡⊡ Fig. 4.9.1 The Guyana coast: Location map. (Courtesy Geostudies.)

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Cozier Canal has been cut. Beyond Dartmouth the coast is swampy, with extensive peat deposits, but there are sectors of sandy beach at Moruka Beach, Waini Beach, Point Kokali and Shell Beach.

Waini River also turns northwest as it approaches the coast, and flows through swamps behind the barrier spit that extends to Waini Point. Here the boundary with Venezuela crosses the coastline.

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⊡⊡ Fig. 4.9.2 Geomorphology of the Guyana coast: Location map. (Courtesy Geostudies.)

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References Brooks J (ed) (1982) South American handbook. Trade and Travel Publications, Bath, England Psuty NP, Mizobe C (2005) Guyana. In: Schwartz ML (ed) The Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, p 907

Smith RT (1962) British Guiana. Oxford University Press, London Wells JT, Coleman JM (1981) Periodic mudflat progradation, north­ eastern  coast of South America: a hypothesis. J Sediment Petrol 51:1069–1075

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4.10 Venezuela

L. Ellenberg

1.  Introduction The Republic of Venezuela has a coastline about 4,993 km long (>Fig. 4.10.1). The coastal climate is subtropical, with a dry season (verano: summer) between December and April and a wet season (invierno: winter) in the rest of the  year. Maracaibo has mean monthly temperatures of 27.2°C  in January and 29.4°C in July, with an average annual rainfall of 577 mm. Tide ranges are small, with mean spring tides only locally exceeding 0.5 m. Wave action produced by the steady trade winds is from the east and northeast, with moderate energy throughout the year, and there is a strong westward longshore current. Hurricanes are very rare. Coasts exposed to strong wave action are generally cliffed, while protected coasts are often mangrove-fringed. Fringing coral reefs are found at many places, the largest

being on the eastern side of the Paraguana Peninsula. Bioerosion features are common, especially on sheltered rocky coasts. Man’s direct influence on the coast of Venezuela is still very small, less than 2% of the coastline having been altered, but indirect influences are great, notably the des­ truction of vegetation and soil cover by farming and ­grazing in recent decades, which has increased fluvial ­sediment yields, especially from the Orinoco, Unare and Tucuyo Rivers, and contributed to accelerated delta growth. The coast is unstable because of tectonic uplift of about 0.5 mm/year in the Cordillera de la Costa (Schubert et al. 1977) and subsidence in the east (Fiedler 1970). Coastal outlines are influenced by the topography of the Cordillera de la Costa, the hills of Falcon and Sucre, and the mesetas of Paraguana and Guahira.

⊡⊡ Fig. 4.10.1 The coastline of Venezuela.

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.10, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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While the variety of coastal features in Venezuela is explicable mainly in terms of geology and marine processes, climatic factors are also important (Ellenberg 1978). Those related to tropical conditions include  the  exten­­ sive coral reefs, luxuriant mangrove swamps, widespread beach rock and such features as intensive supralittoral solution on rocky shores, rapid water-layer weathering and bioerosion, and a prevalence of fine sandy beaches and alluvial shore deposition, wave quarrying being relatively unimportant (Ellenberg 1980). There is a rapid transition from humid tropical conditions in the east to warm arid climate in the west, and this is reflected in coastal geomorphology. Cliffs in soft sedimentary formations are gentle and subject to landsliding after heavy rain (>Fig. 4.10.2) in the wetter eastern sectors, and steeper and dissected only by gully erosion in the drier western sectors. Sebkhas, which are swamps in rainy seasons and saline flats in dry seasons, are found on lowlying coasts throughout Venezuela, but they are larger and more frequent in the drier west. Dunes are found only in sectors where more than 4 months of the year are dry, and where precipitation is less than 900 mm/year. However, beach rock (Reuber and Ellenberg 1979), coral reefs, and mangrove swamps show no correlation with climatic variations along the Venezuelan coast.

2.  The Coastline of Venezuela From the Guyana border the coast is low-lying and swampy, with extensive mangroves (>Fig. 4.10.1, A). Isla Cor­ocoro has sandy outer shores, and there are cheniers on the coastal plain, similar to those of Guyana. The mangroves extend around the large Orinoco delta (25,000 sq. km), where about 40 distributaries open through estuaries to a 360 km deltaic coast (Warne et al. 2002). Progradation continues, about 900 sq. km of land having been added in the twentieth century. The climate is hot and humid, and river flooding is extensive in midwinter (August–September). The mangroves are backed by swamp forests with palms, and there are some sandy areas near their mouths, as at Isla Mariusa. Longshore drifting is westward, particularly on the shores south of Trinidad, where the Caño Macareo distributary is deflected westward. Along the northwest margin of the Orinoco delta are mud volcanoes, related to rapid sedimentation accompanied by ­tectonic compression along the boundary between the South American and Caribbean plates: some are active and sparsely vegetated, others inactive and densely vegetated (Aslan et al. 2001, 2003). The swampy coast comes to an end as the rugged forested mountain Peninsula de Paria (>Fig. 4.10.1, B) runs out east towards Trinidad. This consists of strongly folded

⊡⊡ Fig. 4.10.2 Slumping cliffs at Aguide, east of Puerto Cumaredo.

Venezuela

and faulted Mesozoic limestones, dissected by deep river valleys. It is bordered by steep coasts that descend to cliffs and a beach fringe, with mangroves sparse in the bays and abundant corals on the shore, although reefs are not important. On its northern coast are many bays with sandy beaches, as at Playa Pui Puy and Playa Medina, and near the eastern end is the village of Macuro, where Columbus is thought to have landed in 1498. The coastal ridge runs westward as the more arid Peninsula de Araya on the north side of the Golfo de Cariaco, and at its western end the town of Araya is notable for its extensive salinas (salt works) in artificial lagoons which become coloured pink with Artemia, a microscopic shrimp. In the Mochima National Park steep coasts extend round the Bahia de Mochima, the Manare peninsula and the Islas Caracas, and the narrow Cerro Aceite de Castilla peninsula runs out westward beside the Golfo de Santa Fe. The Islas Caracas are barren and rocky, but have beaches and fringing coral reefs, and similar high islands continue offshore: Isla Chimana, Isla Borrachas and Isla Monos. Playa Colorado has orange sand. Puerto La Cruz is a major port and resort town with a busy seafront. West of the Golfo de Cariaco the coast becomes lowlying, with active sedimentation. A sandy coastal barrier ⊡⊡ Fig. 4.10.3 Macuto, on the steep coast of Cordillera de la Costa.

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encloses lagoons on either side of the Unare delta and again at Lago de Tacarigua to the west (>Fig. 4.10.1, C). Towards Cabo Codera there are mangrove swamps in a bay sheltered from strong wave action, where sediment are trapped by the vegetation (>Fig. 4.10.1, D). Cabo Codera marks the beginning of a steep coast  (>Fig. 4.10.1, E) along the Cordillera de la Costa (>Fig. 4.10.3), which consist of strongly folded and faulted Mesozoic schists, metamorphic and intrusive rocks and Tertiary and Quaternary conglomerates. The mountains rise to the peaks of Naiguata (2,765 m) and Turimiquire (2,470 m). The steep coast descends to a narrow plain with the resort towns of Naiguata, Caraballeda and Macuto, which were severely damaged by mudslides during heavy rains in December 1999 (Perez 2001). To the west is the Parque Nacional Henri Pittier, also a steep coast with cliffs on rocky headlands, slumping slopes, beaches in small bays and mangroves in sheltered inlets. Further along the coast there are rias, such as Bahia de Turiamo. Isla Larga is fringed by beaches and reefs. Porto Cabello (>Fig. 4.10.1, F ), founded in the sixteenth century, has a sector of low-lying sandy coast that has  been  much eroded in recent times. North of Tucacas

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(>Fig. 4.10.1, G) shallow coastal water reduces wave action in the Golfo Triste, and mangrove swamps are extensive in Los Manglares (>Fig. 4.10.4) and bordering the elongated Golfete de Guare, a drowned valley south of Chichiriviche (>Fig. 4.10.5). Small islands offshore are cayos (similar to the Florida keys) with fringing reefs, ridges of coralline gravel, calcareous beaches with beach rock, and mangroves. Near San Juan de los Cayos (>Fig. 4.10.1, H) a beach ridge plain is backed by sebkhas (saline lagoons), and a fringing coral reef borders the San Juan peninsula (>Fig. 4.10.5). The reef, capped by sandy cays, extends intermittently northwest and southeast of this peninsula. EL

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⊡⊡ Fig. 4.10.6 Peninsula de Paraguana, a part of the Venezuelan coast that has been strongly influenced by tectonic movements. After the emergence of the Istmo de Médanos longshore drifting was diverted northward along the east coast, and a beach prograded along the eastern shore of the isthmus (a). This is now protected by a ledge of beach rock, which also extends along the southeast shore of Paraguana (b). In northeastern Paraguana (c) a fringing coral reef was raised above high tide level, and on the western coast (d) sand deposition and spit formation have ensued. West of the isthmus, in the now sheltered waters of the Golfete de Coro, the Rio Mitare (e) has built a large delta (450 sq. km). Sedimentation in and around this bay (f) has led to mangrove encroachment and the formation of saline clay plains, while the waters of the bay have become increasingly hypersaline. Dunes have moved across the southern part of the isthmus (g) and in the north of Paraguana (h) coastline progradation has occurred as the result of dunes moving from the land to the sea.

TODAYS COASTLINE

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wave action, where low cliffs alternate with beaches. A central hill, Cerro Santa Ana, rises abruptly to 830 m through well-defined vegetation zones. A fringing coral reef extends along the north-eastern coast. The western coast is more sheltered and low-lying, with high faultinduced cliffs only around Punto Fijo, which has an oil terminal and salt works. To the west the low-lying coast of eastern Zulia and western Falcón (>Fig. 4.10.1, L) is backed by alluvial deposits, dunes and Pleistocene gravel terraces. Coastal waters are shallow and wave action meagre. The longshore current is weak, especially in the east, where the bird-foot delta of the Rio Mitare has grown rapidly into shallow water ­during the past 3,000 years (>Fig.  4.10.6). Dunes, including some barchans, are migrating across the southern part of the isthmus towards the Golfete de Coro (>Fig. 4.10.7).

The narrow entrance to Lake Maracaibo (>Fig. 4.10.1, M) is fault-bounded, and consists of steep coastal slopes with almost no wave action, and mangrove fringes. Mean spring tide range increases from 0.4 m at Amuay on the west coast of the Paraguana peninsula to 1.1 m at Malecon in the approaches to Lake Maracaibo. Man­grove swamps on sheltered sectors of a shallow bay were the site of lake dwellings on piles, named Small Venice (i.e. Venezuela) when explorers arrived in 1499 (>Fig. 4.10.1, N). Lake Maracaibo (13,300 sq. km, 210 km N-S and 75 km E-W) is shallow with swampy shores, and is fed by many rivers, the largest of which is the Rio Catatumbo, which has built a swampy delta on the southwest coast. This is a region with an unusually high frequency of lightning, possibly due  to turbulence resulting from the interaction of

⊡⊡ Fig. 4.10.7 Dunes moving towards the Golfete de Coro. 69°45'W

500 m 11°28'N

GOLFETE DE CORO DIRECTION OF PREVAILING WINDS

ISTMO DE MÉDANOS SALT FLAT BARCHAN CROSSING THE SALT FLAT

SANDY ACCUMULATION WITH DUNES

TRANARG DEZ.

1976

MISION

0201127

FOTOS

140–141

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Venezuela

⊡⊡ Fig. 4.10.8 Beach ridges, mangrove swamps and sebkhas on the coastal plain of Sinamaica, Gulf of Venezuela.

SAN

D RID

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SINAMAICA, GOLFO de VENEZUELA

SINAMAICA

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M

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O de VENE

ZUEL

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SWAMP IN RAINY SEASON SALT FLAT IN DRY SEASON

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BAHÍA de URABA FROM SHEET 3849, SINAMAICA, VENEZUELA 1 = 100,000 (1963) REPUBLICA de VENEZUELA

Venezuela

cold descending air from the high mountains with hot ascending air from the steamy lake. The southern part of Lake Maracaibo is fresh water, grading to brackish north towards the marine entrance, which is constricted by a long sand bar. A navigable channel 10.5 m deep was dredged in 1957, bordered by a stone breakwater. The lake occupies a Tertiary sedimentary basin which in 1917 was found to contain oil. It soon became a major oilfield, with numerous derricks standing in the water and along the shore, especially in the northeast. On the western shore of the narrow entrance strait is the port of Maracaibo, with oil refineries. West of the entrance to Lake Maracaibo (>Fig. 4.10.1, O) a low-lying coast with sandy beach ridges near Sinamaica (Tanner 1971) is backed by swamps which fringe the north and east shores of the Bahia de Uraba (>Fig. 4.10.8). The longshore current is strong on the open coast. There has been strong uplift on the Venezuelan fringe of the Peninsula de Guajira (>Fig. 4.10.1, P). Cliffs of cemented gravel up to 6 m high are being cut back by strong wave action, and there are intervening shallow bays with gravel beach ridges and dunes. Outlying Venzuelan islands include the Nueva Esparta group, of which the largest is Isla Margarita. This consists of two uplands linked by a sandy isthmus, the northern coast of which has a mangrove-fringed lagoon, La Restinga, enclosed by a sandy barrier. There are numerous beaches of white sand (Alexander and Bertness 1982). To the west is Isla La Tortuga, and farther seaward is a chain of Venezuelan islands from Islas Los Hermanos and Isla Blanquilla to Isla

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Orchila, Los Roques and Isla Las Aves (Schubert and Laredo 1984). Los Roques is a small archipelago which includes a high island, Gran Roque, and a number of white sandy cays on coral reefs.

References Alexander CS, Bertness J (1982) A comparative study of modern and ancient beach morphologies: insights to the paleoclimate of Marga­ rita Island, Venezuela. J Geol 90:663–678 Aslan A, Warne AG, White WA, Guevara EH, Smyth RC, Raney JA, Gibeaut JC (2001) Mud volcanoes of the Orinoco delta, eastern Venezuela. Geomorphology 41:323–336 Aslan A, White WA, Warne AG (2003) Holocene evolution of the western Orinococ delta. Bull Geol Soc Am 115:479–498 Ellenberg L (1978) Coastal types of Venezuela – an application of coastal classifications. Z Geomorphol 22:439–456 Ellenberg L (1980) On the climatic morphology of tropical coasts (in German). Berl Geogr Stud 7:177–191 Fiedler G (1970) Seismic activity in Venezuela in association with important tectonic fault zones (in German). Geol Rundsch 59:1203–1215 Perez FL (2001) Matrix granulometry of catastrophic debris flows (December 1999) in central coastal Venezuela. Catena 45:163–183 Reuber I, Ellenberg L (1979) Beach rock in Venezuela. Acta Cient Venez 30:462–447 Schubert C, Laredo M (1984) Geology of Aves Island, Venezuela, and subsidence of Aves ridge, Caribbean sea. Mar Geol 59:305–318 Schubert C, Valastro S, Cowart JB (1977) Evidence of recent uplift on the north-central (Cordillera de la Costa) coast of Venezuela (in Spanish). Acta Cient Venez 28:363–372 Tanner WF (1971) Growth rates of Venezuelan beach ridges, Sediment Geol 6:215–220 Warne AG, Guevara EH, Asian A (2002) Late Quaternary evolution of the Orinoco delta, Venezuela. J Coast Res 18:225–253

257

4.11 Caribbean Coast of Colombia

Iván Correa . Robert Morton

1.  Introduction The Caribbean coast of Colombia is about 1,600 km long, from Castilletes on the western border of Venezu­­ ela  to Cabo Tiburón at the eastern border of Panamá (> Fig. 4.11.1). It is a relatively developed area with numerous small cities and five large commerce  centres or ports (Riohacha, Santa Marta, Barranquilla, Cartagena and Turbo). Land and air access from the interior of the country is available to all medium and large cities. Colombia’s primary Caribbean island areas are the coralline archipelagos

of San Andrés, Providencia, Santa Catalina Islands and El Rosario Islands, 100 km south of Cartagena. Located at the intersection between the Nazca, Caribbean and South American plates, the Caribbean coast of Colombia is a mosaic of geologic and physiographically varied units composed of both extensive low-relief plains and medium to high relief rocky massifs. The Caribbean coastal zone is crossed by several active faults that define its main morphostructural units. The area has been classified as an intermediate seismic risk zone.

⊡⊡ Fig. 4.11.1 Caribbean coast of Colombia.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The climate of the Caribbean coast (IGAC 2002) depends on the annual displacement of the Intertropical Convergence Zone and on particular orographic influences of the Sierra Nevada de Santa Marta massif. There are generally two rainy periods (April–May and October– November) and two dry periods (December–April and July–September). Maximum annual precipitation for the Colombian Caribbean does not exceed 2,500 mm. Minimum values are within the desert region of the Guajira Peninsula (yearly mean about 267 mm), and maximum values are at the Sierra Nevada de Santa Marta massif (yearly mean 2,000 mm). Mean air temperatures for the Caribbean coast are less than 24°C. Tides along the Caribbean coast are of mixed semidiurnal type, with maximum amplitudes of 60 cm. Trade winds predominate, mainly from the east, north and northwest, at the Guajira Peninsula, and from the NE to NW, south of the Sierra Nevada de Santa Marta. Net longshore drift along the Caribbean coast of Colombia has a dominant southward component, minor reversals to the northeast occurring during the rainy periods when south winds become dominant in some sectors. The Caribbean coast of Colombia is out of the zone of direct influence of tropical cyclones, but is affected by perimeter influences, especially along the northern Caribbean, between La Guajira and Cartagena, where high waves cause extensive shore erosion and lowland flooding.

2.  Southern Caribbean Coast The southern Caribbean coast of Colombia extends for about 550 km from the Barú Peninsula to the Gulf of Urabá (> Fig. 4.11.1). Except for the Gulf of Morrosquillo, the region is cut into sedimentary rocks of the Sinú Belt, a stratigraphic unit composed by Oligocene to Pliocene sequences of turbidities, hemipelagic and terrigenous marine deposits, strongly affected by differential tectonic movement driven in part by mud diapirism and associated phenomena (Duque-Caro 1984; Vernette et al. 1992). Along the littoral fringe, rocks of the Sinú belt are mainly claystones, mudstones and conglomerates. Coral reef limestones outcrop on the tops of coastal hills, which are up to 150 m high. Offshore and inland mud diapirs and mud volcanoes are common on the Sinú Belt, some of them with violent historical eruptions that modified the bathymetry and surface topography of the order of metres (Vernette et al. 1992). Quaternary deposits of the Sinú belt coast are mainly sandy and muddy deltas (Sinú-Tinajones deltaic complex,

Turbo and Atrato deltas) and alluvial valley fills. Along the coast, there are also Holocene terraces up to 36 m high, best developed south of the Gulf of Morrosquillo. Holocene reefs on the shallow platform of the Sinú Belt are located on positive-relief bottom features formed by mud ­diapirism (Duque-Caro 1984; Vernette et al. 1992) and are most common between Cartagena and the Gulf of Morrosquillo (El Rosario Islands). The Gulf of Morrosquillo, 100 km south of Cartagena, is a greatly modified, low wave-energy environment. Internal lagoons and mangrove swamps are fronted by protective beaches and extensive reclamation projects. Southwest of the Gulf of Morrosquillo, the coastal plain is composed of extensive detrital deposits of the Holocene Sinú River delta. The main entrance to the Bay of Cispatá changed abruptly in 1938 as the Sinu River enlarged an irrigation channel cut at Tinajones sector (Serrano 2004; Correa et al. 2007). As a result of channel widening, the Cispatá bay area became depleted of sediment and was submerged by the sea. More than 10,000 ha of cultivated rice land was lost at its margins. The lobate, symmetrical Tinajones delta has an approximate area of 26 km2; both flanks are subjected to intense beach erosion. From Punta La Rada to the Gulf of Urabá, the southern Caribbean coast has an S 50° W trend and is dominated by cliffs cut into deformed sequences of Tertiary mudstone and claystone. The cliffs, which are up to 36 m high, form the seaward limit of late Holocene wave-cut and depositional marine terraces. The coast exhibits a serrated pattern controlled by hard rocky headlands coinciding with structural axes. Shore retreat is the dominant historical trend along most of this coastline. A recent example of land loss along the southern Caribbean coast is in the Arboletes-Punta Rey area. Here the effects of natural shore and cliff erosion are exacerbated by intensive beach-sand extraction to build up the village of Arboletes. The natural processes and human activities caused complete erosion of the 1.5 km long Punta Rey Peninsula between 1957 and 2000 (Correa et al. 2007). Elimination of the wave protection provided by the peninsula triggered extensive cliff and beach erosion to the south. Accelerated erosion also affected the Arboletes mud volcano, the main tourist attraction of the city. Currently stable shores along the southern Caribbean coast are found only at few specific, unmodified stretches that receive relatively high sand supply. The Gulf of Urabá is a relatively wave protected environment with a general N-S orientation and a maximum width of 20 km at its northern end. On its NE part the Gulf shores have narrow beaches backed by cliffs cut into ­sedimentary sequences of the Sinú Belt. The southeastern

Caribbean Coast of Colombia

coast of the Gulf is dominated by erosion, some local deposition associated with recent human activities – the Turbo River delta being the most important. The SW side of the Gulf of Urabá is defined by the prograding Atrato River delta, which has 16 active distributaries protruding into the 25 m deep Gulf. The Atrato River is 730 km long and drains an area of about 36,000 km2 of the western Colombian Cordillera and the Serranía del Baudó, on the Pacific Coast. Rainfall in its drainage basin is 8 m/yr, which supplies a mean water discharge of 2,700 m3/s and a sediment yield of about 11 × 106 ton/year (Restrepo and Kjerfve 2000). Northwest of the Gulf of Urabá and the Atrato River delta, the Caribbean coast changes to basaltic cliffs alternating with wide alluvial valleys and beach ridges. Absence of sediment from the Atrato River, permit the development of fringing coral reefs which become progressively more abundant toward Panamá.

3.  Northern Caribbean Coast The northern Caribbean coast extends for about 600 km, between Castilletes and the city of Cartagena. At its northern sector, the Guajira Peninsula is a main morphostructural element consisting of tectonically raised blocks of metamorphic, granitic and sedimentary rocks (Jurassic to Tertiary in age), adjacent to sedimentary basins and

⊡⊡ Fig. 4.11.2 Aerial view of Cabo de la Vela, showing 5–10 m high cliffs cut into terraces of serpentinitic rocks covered by small dunes and a veneer of wind-transported sand. (Courtesy Diego Zapata.)

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g­ rabens filled with Tertiary limestones, cherts and claystones. The Quaternary of the Guajira Peninsula is mainly represented by extensive colluvial and alluvial deposits, and recent sandy barriers and marine lagoons. The Guajira Peninsula coastline from Castilletes (Gulf of Maracaibo) to Dibulla is about 280 km long. Spits, bars and lagoons predominate along the internal shore of the Gulf of Maracaibo, whereas narrow beaches and cliffs are dominant along the Bahia Honda-Cabo de la Vela shore (> Fig. 4.11.2). South of Cabo de La Vela, the Guajira Peninsula coast is dominated by transgressive narrow beaches, minor ­deltaic accumulations and spit-lagoon segments near the mouths of the primary coastal rivers. These coastal-plain deposits are located seaward of extensive erosional platforms cut into Tertiary mudstones. South of the Guajira Peninsula, the Sierra Nevada de Santa Marta massif is the highest coastal mountain in the world, reaching 5,800 m at Pico Bolívar, 60 km from the coastline. It is composed of Cretaceous metamorphic rocks (predominantly schists and gneisses) and quartz-dioritic intrusive rocks of Tertiary age. The Quaternary of the Sierra Nevada de Santa Marta is mainly represented by colluvial and alluvial valley fill and by recent beach deposits. The coastlines of the northwestern part of the Sierra Nevada de Santa Marta are indented, reflecting the alternation of rocky headlands and deep, NNW trending ­tectonically controlled bays. The headlands are typically

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100–150 m high plunging cliffs cut into schists and granites, whereas the embayments front alluvial valley deposits. The bay shores are typically steep, reflective beaches composed of very coarse sand to granules eroded from adjacent cliffs and alluvial deposits. Wide pocket beaches and tombolos are common along sectors with abundant rocky erosional remnants and stacks (> Fig. 4.11.3). West of the Santa Marta massif, between Santa Marta and Barranquilla, the coastal zone includes the extensive Ciénaga de Santa Marta shallow (<4 m) lagoon and the adjacent Quaternary deposits of the Magdalena River delta. The Magdalena River is the most important river of Colombia, draining a basin of 250,000 km2 on the Andes Cordillera (Restrepo and Kjerfve 2000). The Holocene Magdalena River delta and the adjacent Santa Marta lagoon occupy a low area bounded by the Sierra Nevada de Santa Marta massif and the coastal relief of the San Jacinto Belt, west of the present course of the Magdalena River. This sector is the most extensive depositional area on the Caribbean coast of Colombia. The apex of the delta is about 45 km inland from the river mouth, at the jetties of Bocas de Ceniza. The Santa Marta lagoon and mangrove system has only one connection with the sea (Tasajeras mouth) and is confined by the Salamanca bar, a transgressive sandy barrier about 65 km long that forms the present coastline of the Magdalena delta. The Salamanca bar has a maximum height of about 2 m above present sea level and has been historically subjected to intense erosion, currently reflected by the  generalised beach retreat along most of its

length and by frequent overwash events along the eastern Magdalena delta. Coastal formations between Barranquilla and Cartagena are mainly associated with the San Jacinto Belt, a deformed, Middle Eocene sequence of pelagic, hemi-pelagic and turbiditic rocks (Duque-Caro 1984). The coaastline of these formations is typically cliffs up to 50 m high, backed by hills dissected by minor Quaternary alluvial valleys. The Barranquilla-Cartagena coastline has an approximate length of 100 km and a general S 45° W orientation, controlled by the structural trend of the Tertiary rocks. The northern coastline of this area shows a serrated pattern consisting of steep cliffs (5–40 m high) cut into consolidated Tertiary sediment of the San Jacinto Belt (> Fig. 4.11.4). Most cliffs are fronted by narrow beaches supplied by erosion of Tertiary formations and sand eroded from the Magdalena River delta. Near Cartagena, the coastline is dominated by depositional features fronting inactive cliffs, mainly sandy spits and bars that modulate the coastal indentations. Historical information on this area (Vernette et al. 1992) shows large shoreline changes associated with mud diapirism along this segment of coast, including erosion of a 12 km long, E-W oriented spit reported in 1794 and the formation of up to 7 km2 tombolos. Extensive land reclamation has been carried out to enlarge the city of Cartagena. The Tierrabomba Island and the Barú Península are composed of Plio-Pleistocene mudstones and limestones that show evidence of recent relative sea level changes, including wave-cut platforms, elevated caves and stacks.

⊡⊡ Fig. 4.11.3 View southward over Cañaverales Bay at the Santa Marta massif, showing highly reflective coarse sand beaches, tombolos and a rocky NW coastline.

Caribbean Coast of Colombia

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⊡⊡ Fig. 4.11.4 Steep cliffs cut into Tertiary sandstones and mudstones and fronted by narrow beaches at El Morro Point, south of Barranquilla. (Courtesy Jesus A. Pérez.)

The features are as much as 5 m above present sea level and have been dated as late Holocene. Steep erosional cliffs and narrow calcareous pocket beaches dominate the outer coastlines of Tierrabomba Island and the Barú Peninsula. Extensive mangrove swamps are found along the margins of the Cartagena and Barbacoas Bays.

References Correa ID, Acosta S, Bedoya G (2007) Análisis de las causas y monitoreo de la erosión litoral en el departamento de Córdoba. Fondo Editorial Universidad EAFIT, Medellin, pp 1–127

Duque-Caro H (1984) Estilo estructural, diapirismo y episodios de acrecimiento del terreno Sinú – San Jacinto en el Noroccidente de Colombia. Boletín Geológico Ingeominas 27:1–29 IGAC (2002) Atlas de Colombia 5th edn. Bogotá, p 342 Restrepo JD, Kjerfve B (2000) Magdalena River: interannual variability (1975–1995) and revised water discharge and sediment load estimates. J Hydrol 235:137–149 Serrano BE (2004) The Sinú River delta on the northwestern Caribbean coast of Colombia: bay infilling associated with delta development. J South Am Earth Sci 16:623–631 Vernette G, Maufret A, Briceño L, Gayet J (1992) Mud diapirism, fan sedimentation and strike-slip faulting, Caribbean Colombian margin. Tectonophysics 202:335–349

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4.12 Caribbean Coasts, Panama to Belize

Anja Scheffers · Tony Browne

1.  Introduction The Caribbean coasts of Central America extend for about 2,200 km from the Atrato lowlands in northwestern Colombia to Tehuantepec in Mexico. The geology of Central America is complex. Two large areas of different tectonic history and geological structure have been recognized: namely the Isthmian Link from NW Colombia to southern Nicaragua and Nuclear Central America, from southern Nicaragua to SE Mexico. The Isthmian Link (Weyl 1980; Dengo 1985) has a basement of Late Jurassic to Late Cretaceous age, which consists mainly of basalts with associated radiolarites and clastic sedimentary rocks. The basement is overlain by thick volcano-clastic rocks of Late Cretaceous to Pliocene age, with some interbedded limestones. The prevailing structures are parallel to the geographical shape of the isthmus. South of the mountains of Guatemala the metamorphic basement is covered by Mesozoic sedimentary rocks: Upper Triassic (in Honduras), Jurassic sediment and Cretaceous carbonates and red beds, intruded by granitic bodies. These, in turn, are overlain in extensive areas by Tertiary red beds and ignimbrite plateaus. A Lower Palaeozoic metamorphic and igneous basement (probably including some Pre-Cambrian rocks) forms some of the high mountains in northern Nicaragua, Honduras, Central Guatemala and SE Mexico. North of the mountains of Guatemala, extending westward into Mexico, the basement is overlain by a thick sequence of upper Palaeozoic sedimentary rocks, which in turn are covered by Upper Jurassic and thick Cretaceous to Eocene strata. Upper Tertiary marine sediment are restricted to small areas. The geological structure north of the Central Cordillera is dominated by a Laramidian fold belt that extends from SE Mexico eastward across Guatemala and southern Belize, forming an open arc concave to the north and with an ENE trend as it reaches the Caribbean coast. North of the fold belt the structures vary from gentle to nearly horizontal in the large area of northern Guatemala and Belize and the Yucatan Peninsula in Mexico.

A common feature of Central America is a Quaternary volcanic mountain chain that runs NW–SE, and is superposed on the older structures. The continental divide in Central America is generally closer to the Pacific, so that the longer streams drain to the Caribbean Sea. A high rainfall (5–6 m/year), notably during hurricanes, yields a large river discharge and vast amounts of fine grained sediment are transported to the Caribbean coastline, where they are distributed alongshore and mostly form coastal lowlands with beaches and beach ridges. In Panama and Costa Rica the continental divide is more central, except in the area east of the Panama Canal, where it is very close the northern coastline and results in short steep streams flowing down to the Caribbean. The coast is microtidal (around 0.5 m spring tide range). The Caribbean coast is divided into four sectors according to their similarities in morphology and origin: Cape Tiburón on the Colombian border to Port Limon in Costa Rica, Port Limon to Cape Gracias a Dios (NicaraguaHonduras), Cape Gracias a Dios to the Sarstun River mouth (Guatemala) and from there to Chetumal Bay on the Mexican border. The geomorphology and ecology of this coast has been described by Yanez-Arancibia (2005).

2.  Panama The state of Panama has the longest Caribbean coastline of Central America, more than 600 km. West from Cape Tiburón on the Colombia border the mountains are close to the coast and the beaches are interrupted by frequent long cliffs, with some fringing coral reefs. The coast is bordered by numerous small islands (the Archipelago de San Blas), made up by a drowned part of the basic igneous rock basement of southern Central America. Towards the Canal entrance there are recently uplifted shore platforms. In many places the tropical forest almost reaches the beach. From Manzanillo Point to Chiriqui Point the north-facing coast is formed by a succession of narrow beaches separated by low cliffs where the mountains reach the sea. The coastline becomes ­intricate behind the Peninsula

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de Valiente, where the Lagoa de Chirigui and Almirante Bay are backed by Tertiary sedimentary and volcanic rocks, which also form the islands inside the bay. A part of the northern bay shows partly drowned cockpit karst.

3.  Costa Rica The coastline of the Caribbean side of Coast Rica is much shorter than that of the Pacific side, only about 200 km in length. Offshore the Miskito Keys and banks are formed by coral reefs, while the islands of Great and Little Corn consist of volcanic rocks and present small beaches separated by cliffs. At Mona Point, Cahuita Point and Port Limon there are uplifted Pliocene coral reefs and Holocene fringing reefs. North of Point Limon the coastal plain widens as the mountainous hinterland recedes. There are long gently curving beaches, some with concentrations of magnetite and ilmenite. To the NW long sandy barriers are backed by narrow coastal lagoons, notably the Tortuguero Lagoon. The beach sands are derived largely from Quaternary volcanic rocks (Nieuwenhuyse and Krooneneberg 1994). The barriers continue to the San Juan River delta.

4.  Nicaragua The Caribbean coastline of Nicaragua runs north for more than 450 km. San Juan River, which drains the Nicaragua lake basin, forms the boundary with Costa Rica. To the north is the Costa de Mosquitos between Cabo Gracias a Dios and Bluefields where mangroves, notably Rhizophora mangle, are extensive around lagoons and on swampy shores. Coastal sediment are derived mostly from low cliffs of Tertiary volcanic rock, but they include material from the metamorphic basement and Cretaceous limestones. There are several elongated lagoons bordered by sand barriers, the largest being the Perlas Lagoon. In the northern part of the coast, near Puerto Cabezas, extensive Plio-Pleistocene low terraces occur, formed largely of quartzose sand and pebbles. The coastal plain narrows towards Cape Gracias a Dios, at the mouth of the Coco River, which forms a prograding delta.

5.  Honduras The Caribbean coast of Honduras is at least 600 km long, excluding large islands (up to 50 km long) in the north.

From Cape Gracios a Dios the coast turns NW and a long sandy barrier encloses the large Caratasca Lagoon (Helbig 1959). This lagoon includes several small sandy islands, and is surrounded by a low swampy area with extensive mangroves. It was probably formed as a result of changes in the lower course of the Patuca River leaving a broad embayment across which a barrier was built by sediment delivered by the Coco River and transported alongshore to the northwest. The hinterland of the lagoon has extensive, low Plio-Pleistocene terraces. West of the Patuca delta the mountain ranges locally reach the coast. Between Cape Camaron and Port Cortes long narrow beaches extend between cliffs of Palaeozoic metamorphic rock and Tertiary ignimbrites, except near the mouths of the Aguan and Ulua Rivers, where the valleys merge into wider coastal plains, some swampy and with mangroves. The large rivers have provided a considerable amount of sediment, including sand that has been built by storm waves into beach ridges. Longshore drifting is westward, and has led to the growth of sand spits, such as the spits that border Brus Lagoon near the mouth of Patuca Ruver and Cape Camaron. Sand drifting west from the mouth of Aguan River has built a spit that partly encloses a bay at Port Trujillo. The beach sands are heterogeneous because the sediment source consists of metamorphic rocks, limestones, granitic intrusive rocks, and ignimbrites. Some of the sands have a large quartz content derived from the granitic rocks and the rhyolitic ignimbrites. In 1961 Hurricane Hattie caused extensive coastal flooding in British Honduras. damaging coral reefs and eroding and destroying sand cays (Stoddart 1962). The Bay Islands (Islas de la Bahia), north of the Honduras coast, are a continuation of the mountain system of Nuclear Central America. They are formed mostly by metamorphic rocks and one small volcanic cone (Utila). Their beaches are small and discontinuous, separated by low cliffs, and the whole island group is surrounded by a submarine coral reef.

6.  Guatemala The Sarstun River flows into Amatique Bay, which extends about 100 km along the Guatemalan coast. This bay is located on a graben, extending inland past Lake Izabal and continuing seaward as the Cayman Trough. The bay is separated from the sea by a large sand spit, Manabique Point, formed of sediment delivered by the Motagua River and transported NW by longshore drifting. The inner side of the bay is a low hilly area formed by Miocene limestones

Caribbean Coasts, Panama to Belize

and cut by the Rio Dulce gorge, the outlet from Lake Izabal. This sector has only small discontinuous beaches. The Motagua River delta is partly covered with mangroves. The beach sands include pumice carried by the Motagua River from the Quaternary volcanic deposits of the Guatemalan highlands.

7.  Belize The Belize coast has an overall length of about 300 km with some additional islands on the outer reef. The coast is dominated by carbonate deposition and reef growth (Stoddart et al. 1982; Gibson and Carter 2002). Some beaches  have been cemented by carbonate precipitation (Gischler and Lomando 1997). One of the largest barrier reefs in the world (more than 200 km long, with additional coral islands outside and atolls inside), second only to Australia’s Great Barrier Reef, forms a prominent feature in this area. It extends from the Seal Cays in southern Belize to Bahia de la Asencion in Mexico, and is interrupted by several gaps, notably a large one off the mouth of Belize River. Hurricane impacts on reefs were reported by Stoddart (1962, 1965). The reef is separated from the main coast by a shallow shelf filled with carbonate sediment (Bonis et al. 1970). The coastline and the distribution of the offshore coral reefs are structurally controlled by a series of sub-parallel normal faults trending NNE. Oil exploration drilling and marine seismic profiling have shown that the major reefs, such as Chinchorro, Turneffe, Ambergris, and Key Long, have been built since the Pliocene over basement structural highs, while the smaller linear ones are over the upthrown edges of tilted fault blocks. North from the Guatemalan border Cretaceous limestones outcrop along the coast, forming prominent cliffs separated by short, narrow beaches near Punta Gorda. North to Stann Creek Palaeozoic rocks of the Maya Moun­ tains come close to the coast, and the beach sand contains much quartz. North of Stann Creek the coastal plain widens and there are elongated lagoons behind sandy barriers.

4.12

The low-lying area around Belize City is formed by fine calcareous sediment deposited by the Mopan River, which forms a small delta on which the city was built. The coastline of northern Belize is formed by a series of calcareous sand barriers enclosing lagoons with mangrove shores, bordered inland by Plio-Pleistocene terraces, also formed predominantly of limestone fragments derived from the older carbonate rocks (Wright et al. 1959). Luxuriant mangroves are found in the shelter of the  barrier reef on carbonate platforms elevated upon mean sea level. Many of them show distinctive species ­patterns, depending on the distance to open water and the  duration of submergence, related to the level of the substrate.

References Bonis S, Bohnenberger O, Dengo G (1970) Mapa Geologico de la Republica de Guatemala (1:500,000). Instituto Geografico Nacional, Guatemala Dengo G (1985) Caribbean Central America. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 117–124 Gibson J, Carter J (2002) The reefs of Belize. In: Cortes J (ed) Latin American Coral Reefs. Elsevier, Amsterdam Gischler E, Lomando AJ (1997) Holocene cemented beach deposits in Belize. J Sediment Petrol 110:277–297 Helbig K (1959) Die Landschaften von Nordost Honduras, Geographische Kartographische Anstalt Gotha Nieuwenhuyse A, Kroonenberg SB (1994) Volcanic origin of Holocene beach ridges along the Caribbean coast of Costa Rica. Mar Geol 1201–1202:13–26 Stoddart DR (1962) Catastrophic storm effects on the British Honduras reefs and cays. Nature 196:512–515 Stoddart DR (1965) Re-Survey of Hurricane Effects on the British Honduras Reefs and Cays. Nature 207:589–592 Stoddart DR, Fosberg FR, Spellman DL (1982) Cays of the Belize Reef and Lagoon. Atoll Res Bull 256:1–76 Weyl R (1980) Geology of Central America. Bornträger, Berlin Wright ACS, Romney DH, Arbuckle RH, Vial VE (1959) Land in British Honduras. Colonial Research Publications, London Yanez-Arancibia A (2005) Middle America, Coastal Ecology and Geo­ morphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science, Springer, pp 639–645

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4.13  Mexico, Caribbean Coast    Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez     

1.  Introduction The Caribbean coast of Mexico is about 2,300 km long. The NE trade winds blow onshore from the Caribbean and produce waves on the Mexican coast. There are occasional hurricanes. Tide ranges are smaller than on the Pacific coast. On the Gulf Coast Campeche has a mean tide range of 0.3 m, Veracruz 0.5 m, Tuxpan 0.9 m and Tampico 0.4 m. The Caribbean coast is classified as a marginal sea (Inman and Nordstrom 1971; Carranza et al. 1975), and includes the low-lying Yucatan Peninsula. This passes westward to a coastal plain with broad Quaternary alluvial terraces, which varies in width from almost zero where the San Andres Tuxtla and Jalapa highlands reach the coast to about 150 km near the major rivers. It is backed by the Sierra Madre foothills. The continental shelf is wide off Tabasco out to the Campeche Banks, and narrows from Veracruz to Tampico. The eastern shore of the gently rolling, karstic Yucatan Peninsula is irregular and indented by bays backed by extensive swamps. The western shore is rather simple with minor irregularities and narrow marshes. There are extensive lagoons and barriers across the northern, eastern and western sides of the emerged peninsula. The coast declines to a plain towards the Laguna de Terminos. Towards the Tonala River are coastal marshes, swamps, and shallow lagoons (Murray 1961; Gutierrez-Estrada 1982). It is an active depositional coast. Inland there are deeply dissected Quaternary river and marine terraces. The Tabasco plain is a coastal lowland with swamps, marshes, and lakes along the coast westward from the Laguna de Terminos to the Tonala River (Murray 1961; Psuty 1966; Thom 1967). The largest lowland area is in eastern Tabasco, where the Grijalva and Usumacinta rivers have built alluvial plains. The coast consists of lagoons and barriers with sand hills and dunes. The Tabasco segment includes parts of Campeche, Tabasco and Veracruz. The coastal plain averages 90 km in width; it is narrowest in western Tabasco and eastern Veracruz, where there are

elevations of around 300 m. It widens into the basin of Coatzacoalcos River. The Isthmian embayment, from Yucatan to the Jalapa highlands, is divided into the Tabasco and Veracruz segments at the Sierra de Los Tuxtlas. Sand dunes extend from the Sierra de Los Tuxtlas to the Jalapa highlands, and are well developed between Alvarado and Veracruz, where they form the barrier between the Laguna de Alvarado and the Gulf. There is a coastal convex arc, apparently related to volcanism and uplift, and the coastline is fairly regular but for several cuspate bars and lagoons. The plain gradually narrows northward from the Papaloapan River basin. The northern part of the Veracruz segment is developed principally on volcanic material of late Tertiary and Quaternary ages. Maximum elevations are of the order of 200 m. Coral reefs are common in the vicinity of Veracruz city. North of Veracruz the coastline consists of shallow lagoons and marshes partially enclosed by spits and barrier islands. Laguna de Tamiahua and associated cuspate bars and reefs are the principal irregularities. North to the Panuco River Tertiary hills have elevations of the order of 100 m just west of Laguna de Tamiahua (> Fig. 4.13.1). Beyond the Panuco River the coastal plain is interrupted by the Sierras de Tamaulipas, Aldama. and Madre Oriental. The Tampico embayment is a physiographic unit of the coastal plain in east-central Mexico. The plain slopes gently eastward from the Sierra Madre Oriental to the Gulf, and has a maximum width of about 150 km. There are flat areas in its central part and along the Pnuco, Tamesi, and Tuxpan rivers, where marshes and swamps are extensive. Inland are the uplifted ranges of the Sierra Madre Oriental, from which gravelly fans and terraces extend towards the coast. North of Tampico the coastal plain narrows in front of the Sierra de Tamaulipas, then widens where the Laguna Madre is fronted by a long coastal barrier breached by a few marine inlets. Marshes and mangroves are common. The coastline curves gently NNE towards the lobate delta of the Rio Grande (Bravo) on the Texas border.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.13, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4.13

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⊡⊡ Fig. 4.13.1 Predominant coastal landforms of the Caribbean coast of Mexico. (Courtesy Geostudies.) Bravo R.

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Carranza-Edwards A, Gutierrez-Estrada AM, Rodriguez-Torres R (1975) Morphotectonic units of the Mexican coast (in Spanish). An Centro Cienc del Mar y Limnologia, Univ Nal Auton Mexico 2:81–88 Gutierrez-Estrada M (1982) Geomorphology and Recent sediments of the Atasa-Pom lagoon system, Campeche, Mexico (in Spanish). An Centro Cienc del Mary Limnologia, Univ Nal Auton Mexico 9:89–100

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Inman DL, Nordstrom CE (1971) On the tectonic and morphological classification of coasts. J Geol 79:1–21 Murray GE (1961) Geology of the Atlantic and Gulf coastal province of North America. Harper, New York Psuty NP (1966) The geomorphology of beach ridges in Tabasco, Mexico. Louisiana State University Coastal Studies Institute, Technical Report 30:1–51 Thom BG (1967) Mangrove ecology and deltaic geomorphology: Tabasco, Mexico. J Ecol 55:301–343

5.0 Caribbean Islands – Editorial Introduction

The Caribbean islands (Blume 1974) lie in a broad arc, starting east of the low-lying Mexican peninsula of Yucatan with > Cuba, which has extensive lowlands. The Greater Antilles include three more large and mountainous islands: > Jamaica, > Hispaniola and > Puerto Rico extending eastward. There is then a long curving island chain, the Lesser Antilles, which continues as a series of high volcanic islands southward from Anguilla to Grenada, with an outer chain of low-lying islands that includes Barbuda, Antigua, Guadeloupe and Barbados. They then extend westward as the Dutch Lesser Antilles off the Venezuelan coast. The islands formed along an arcuate zone of crustal weakness between the North and South American crustal plates. The inner Lesser Antilles originated as a chain of submarine volcanoes, several of which grew above sea level. Volcanic eruptions have continued, notably at Souifriere on St. Vincent and Chances Peak on Monserrat, and there are also hot springs and solfatara. The outer chain of the Lesser Antilles consists of islands of limestone on old volcanic foundations. Beaches on volcanic islands are mainly of dark basaltic sand, often with gravel and boulders. Coral reefs (particularly fringing reefs) are extensive and beaches derived from these are of pale coralline sand and gravel. White ­calcareous beach sands have also been derived from shelly

debris, algae and other sea floor organisms. Beach rock, formed by cementation of sand by carbonates, has been exposed by erosion of some beaches. The outer shores of the Lesser Antilles north from Grenada and Barbados receive ocean swell from the Atlantic, which also breaks on the northern shores of Puerto Rico, Hispaniola, and to a lesser extent (because of protection by the reefs, banks and islands of the Bahamas to the north) Cuba. The swell is accompanied by waves generated by the NE trade winds (with some from the SE in summer) and these are dominant within the Caribbean Sea, shaping beaches on the southern coasts of Puerto Rico, Hispaniola and Jamaica. During summer and autumn hurricanes occur, moving in from the east over the Lesser Antilles, through the Caribbean Sea, across the Greater Antilles, and into the Gulf of Mexico. They cause erosion of beaches and cliffs, and damage to coral reefs and coastal vegetation. Mean spring tide ranges are small, generally less than 0.5 m. Mangroves occur locally in sheltered situations such as reef lagoons, inlets and estuaries and seagrass beds are extensive in shallow waters.

Reference Blume H (1974) The Caribbean islands. Longman, London

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

5.1 Cuba

Ridel Rodríguez

1.  Introduction The island of Cuba is about 1,250 km long and 30–190 km wide, and has an irregular coastline about 5.800 km long (> Fig.  5.1.1). The Cuban archipelago has four island groups: the Los Colorados archipelago, the SabanaCamagüey (Jardínes del Rey) archipelago, the Jardínes de la Reina (Gardens of the Queen) archipelago and the Los Canarreos archipelago. Cuba is geologically complex, and forms the northern part of the Antilles island arc (Khudoley 1967; Fairbridge 1975), with Jurassic and Cretaceous rocks, mainly in the mountainous areas, and Palaeogene to Quaternary in the bordering lowlands. About two-thirds of the island consists of limestone, and karstic features are well developed, particularly in the eastern section. The long and narrow island of Cuba has generally short and steep rivers. The longest is the Cauto River (370 km), followed by the Sagua la Grande River (163 km) and Zaza River (155 km). Lying immediately south of the Tropic of Cancer, Cuba has a humid to subhumid tropical climate, with two clearly

defined seasons, the dry one (winter) from November to April, and the rainy one (summer) from May to October, with continuous breezes the whole year. The north coast is exposed to the NE trade winds which bring heavy rainfall in the summer months: Havana has a mean annual rainfall of 1,226 mm, of which 868 mm fall between May and October. Mean annual temperature is 24°C (20°C in winter and 27°C in summer). In general rainfall is more abundant in the west than in the east of the country. The south coast is somewhat drier, but the whole of the island coast is subject to occasional hurricanes, especially in August, September and October. Tide ranges are small: Havana has a mean spring tide range of only 0.5 m.

2.  The Cuban Coastline The Cuban coastline is extremely irregular and intersected by numerous gulf and bays, such as the Batabanó and Guacanayabo Gulfs. There are good natural harbours in the bays. Coral reefs and keys extend along the coast. Parts

⊡⊡ Fig. 5.1.1 Cuba. Location map. (Courtesy Geostudies.)

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of the coast are steep and cliffed; others have beaches or mangroves. Although the marine provinces bordering Cuba are rich in biogenic sediment, beaches on the more exposed parts of the coast have been eroding in recent decades. 86% of Cuban beaches are affected by erosion (>Fig. 5.1.2), although on the NE coast this drops to 58.6% (Juanes 1996, Rodríguez and CÓrdova 2006). Erosion is due to a deficit in sand supply, a rising sea level, the effects of hurricanes and human impacts. Progradation has continued where there is a sand supply: some river sediment yields have been augmented by the effects of open-cast nickel mining in their catchments, and this has led to local accretion near river mouths. Havana stands beside a coastal inlet formed by Late Quaternary marine submergence of a valley mouth, but there are terraces 8 and 16 m above sea level indicating Pleistocene uplift, and others ranging laterally between 15 and 33 m, and 25–51 m, indicating transverse tilting (Ducloz 1963). West of Havana as far as Punta Gobernadora the coast is fringed by low cliffs and sandy beaches with inlet and bays such as Mariel Bay. Mangroves are extensive along the NW coast (>Fig. 5.1.3), mainly behind sandy barriers and spits or fringing coral reefs, which develop west of Punta Diamante and are close inshore off Punta Gobernadora.

In  the hinterland are the Pinar del Rio Mountains, from which short, steep rivers descend to coastal inlets. The chain of coral reefs and cays becomes the Archipielago de los Colorados, sheltering the coast round to Golfo de Guanahacabibes at the western end of the island. The shores have a mangrove fringe, particularly in the Bahia de Guadiana. The limestone Peninsula de Guanahacabibes has mangroves on its northern shores, but the southern coast has cliffs and bluffs, and is beachfringed eastward from Cabo de San Antonio past Cabo Corrientes to Cabo Francés on a coast that faces deep water and receives relatively strong wave action. East of Cabo Francés there are sectors of mangrove and some enclosed lagoons: Laguna de Cortez, Laguna Cheve. There are minor headlands between beaches (playas) edging a coastal plain, Llanura Aluvial del Sur. The continental shelf widens eastward across the ­relatively shallow Golfo de Batabanó, and Cayos de San Felipe and extends towards a large island to the south, Isla de Pinos. This has hilly ranges of metamorphic rock amid broad lowlands. Nueva Gerona stands on the north coast, and the western and eastern shores are embayed, while the curved southern coast facing deep water is partly cliffed, and passes eastward into a chain of reefs and small islands - Archipielago de los Canarreos, curving east towards Cayo Largo.

⊡⊡ Fig. 5.1.2 Erosion at Guardalavaca Beach, caused by storm waves.

Cuba

5.1

⊡⊡ Fig. 5.1.3 Mangroves near Zapata.

The northern coast of the Golfo de Batabanó has sandy beaches and mangrove-edged bays, but becomes swampy as wave energy diminishes eastward into the Ensenada de  la Broa, an embayment on the northern side of the Peninsula de Zapata. This low-lying limestone peninsula has swampy northern shores and an irregular cliffy southern coast. To the south–east a deep and narrow submarine canyon, the Golfo de Cazones, curves out southeastward across the narrowing continental shelf. A wide slot-shaped marine inlet, Bahia de Cochinos (Bay of Pigs) is set into the coast to the east, and beyond that is the Bahia de Cienfuegos, with a narrow entrance strait. To the southeast the coast becomes steep below the Sierra del Escambray, a range of mountains rising to Pico San Juan (1,140 m), but remains beach-fringed. Next comes a coastal lowland, with mangroves behind small barriers and spits, running out to the beach-fringed Peninsula de Ancon (> Fig. 5.1.4). At Trinidad a recurved spit, Punta Casilda, marks the beginning of a coast less exposed to strong wave action as the continental shelf widens again eastward, with a chain of coral reefs and cays near its outer edge forming the Archipielago de Jardines de la Reina. The mainland coast remains low-lying, with fringing mangrove swamps, notably around the shallow Golfo de Ana Maria. The River Najasa flows down to the coast at Santa Cruz del Sur.

To the east wave energy again declines, this time into the Golfo de Guacanayabo, north of the Manzanillo Peninsula. The delta of the Rio Cauto is growing on the eastern shore of this gulf, north of Manzanillo, and has shores that are partly sandy and partly muddy, with mangroves. The peninsula runs out to Punta Casimba, from which low hilly ridges grow eastward into the Sierra Maestra Mountains, which rise to 1,974 m and descend abruptly to the coast, trenched by short, steep valleys. Terraces indicate intermittent uplift on this bold coast up to more than 300 m above sea level. Santiago de Cuba (> Fig. 5.1.5) stands beside an estuarine inlet as the ranges decline eastward. To the east several ­rivers converge in Guantanamo Bay, a broad ria with the delta of the Guantanamo River on the western shore. The southern part of the bay is occupied by a United States naval station, extending out to Punta Balovento. The coast steepens again eastward to Punta Caleta and the ridges decline to low-lying Cabo Maisi at the eastern end of the island. The northeast coast of Cuba is also steep, backed by the Sierra Nipe-Sagua-Baracoa, with Baracoa at the mouth of the Toa valley. There are nearshore coral reefs to the northwest. To the west are lowlands around a series of swamp-fringed bays, Bahía de Tánamo, Bahia de Levisa and Bahia de Nipe, formed by marine submergence of

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⊡⊡ Fig. 5.1.4 Sandy beach at Ancon.

⊡⊡ Fig. 5.1.5 Santiago de Cuba. (Courtesy Victor Aluija.)

v­ alley mouths and low areas, and extending to the Bahia de Banes. The coast then curves north to Cabo Lucrecia. Westward are more inlets and bays, including mangrove-fringed Bahia de Puerto Padre, Bahia de Malaguata, Bahia de Manati and Bahia de Nuevitas, which marks the beginning of a submerged lowland corridor behind an elongated peninsula that breaks westward into a chain of

reefs, cays and emerged coral limestone islands – the Archipielago de Camagüey, including Cayo Sabinal, Cayo Guajaba and the long Cayo Romano. There has ­evidently been transverse tilting along the coast. Bahia Mayanabo and Bahia La Gloria, separated by a swampy isthmus, continue westward into a long lagoon, the inner (southern) shore intricate, with mangrove swamps, inlets

Cuba

and the Laguna de Leche. Offshore is the Canal Viejo de Bahamas, and beyond it the Grand Bank of the Bahamas. Further west the Bahia de Santa Clara continues the elongated lagoon behind another chain of reefs and cays – the Archipielago de Sabana. At its western end the Bahia de Cardenas is sheltered by a large spit that has grown northeast, the Peninsula de Hicacos, with a recurved termination at Punta de Morlas and the remains of an earlier one at Punta Garda. From here past Matanzas to Havana the north-facing sandy beaches are almost continuous, with some low cliffs and a shallow sea, and there are numerous seaside resorts. Beach erosion has occurred here, particularly during stormy winters.

5.1

References Ducloz C (1963) Etude géomorphologique de la région de Matanzas. Cuba. Archives de Sci Genéve 16:351–402 Fairbridge RW (1975) Cuba. In: Fairbridge RW (ed) Encyclopaedia of world regional geology, Vol 1. Dowden, Hutchinson, & Ross, Stroudsburg, PA, pp 252–255 Juanes JL (1996) The erosion in the beaches of Cuba – alternatives for their control. Thesis in option of doctor’s scientific degree in geographical sciences. Havana Khudoley KM (1967) Principal features of Cuban geology. Bull Am Assoc Petr Geol 51:668–677 Rodríguez R, Córdova E (2006) Erosion on the beaches of the northeastern region of Cuba. Pan-American Institute of Geography and History, México

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5.2 Jamaica John Norrman · Tommy Lindell 

1.  Introduction Jamaica is located on the northern margin of the Caribbean Plate, which dips off the narrow northern Jamaican shelf into the Cayman Trench, reaching depths of 7,000 m. The island is 240 km from east to west and 60–80 km front north to south. A central east to west backbone of volcanic and metamorphic rocks forms a discontinuous mountain ridge surrounded by limestone plateaux, which cover 60% of the island. The highest part of the mountain ridge, the steep Blue Mountains, is found in the east with a peak of 2,292 m. To the west of these mountains the hilly landscape is generally about 850 m above sea level, gradually declining to 550 m in the western part (Lindell 1999). There are two major fault systems, one in an east-west direction along the straight central north coast. Raised reefs are numerous, up to a height of 180 m. The other fault system runs mainly in a NW–SE direction throughout the island (Norrman et al. 1997a, b). Along the coast these faults form step-like platforms and cliffs. Variation in geology, landscape morphology and ­climate are the major factors which generate coastal landforms. There is surprisingly large coastal variability, reflecting these factors and the associated vegetation. A series of large alluvial basins in the southern part are occupied by agricultural lowlands and almost inaccessible wetland morass with various kinds of vegetation. The island coasts have a tropical climate. Kingston has mean monthly temperatures of 24.4°C in January and 27.2°C in July, with an average annual rainfall of 800 mm. There is a major contrast between the broad landscape relief between the Blue Mountains in the east and the hills, plateaux and low plains to the west. This contrast, combined with the easterly trade winds, has a strong influence on the climate, resulting in mountain rain forests in the east (mean annual rainfall up to 5,600 mm), and dry savannah type vegetation in the south and west (down to less than 1,000 mm). Torrential rivers carry large quantities of coarse sediment (sand and gravel), derived from the steep eastern mountains and the high hilly areas down to the eastern

parts of the north and south coasts, and the sand and silt derived from slope erosion and agricultural soil degradation are transported by the sluggish streams that drain the wide basins in the south. The tide range on the Jamaican coast is less than 0.3 m; Port Royal has a mean spring tide range of 0.2 m. Coastal processes are dominated by wind-generated waves, and wave energy is unevenly distributed. Comparatively high waves, generated by the trade winds, frequently reach the E, NE and SE coasts, whereas the W coast is sheltered. Terrestrial processes produce sediment ranging from boulders to silt and clay, the coarser material deposited on beaches, while the fine sediment flows out in suspension by the waves to settle offshore or in very sheltered areas along the coast. The trade winds are rather constant, generating a dominant longshore current from east to west along the northern and southern coasts. In bays, a counter-current may occur. Climatic conditions in the Caribbean during autumn and winter favour the generation of hurricanes, which are unpredictable. The hurricane tracks that affect Jamaica usually approach from SSE–SE, south of Haiti, but some tracks pass over the eastern part of Cuba and then turn west to the north of Jamaica. A hurricane passing within about 65 km from the coast generates nearshore waves with a turbulent flow of at least 2–4 m/s (Woodley 1992). In 13 of the last 120 years, hurricane centres have passed close to Jamaica, with a median interval of 6.5 years. Much destruction is caused on the coast by the winds and waves of hurricanes. Low-lying areas are flooded, vegetation damaged and swept away and beaches  re­­moved, but features are restored by natural processes in subsequent calmer weather. However, hard, branching staghorn corals are a special problem as they are brittle, and are ­broken by storm waves, but these do not severely harm the solid reef foundation. According to Wood­ley  (1992) the branching corals grow quickly, and are restored to full maturity in about one to two decades. Man-made structures are vulnerable and expensive to repair. Coral reefs are the marine ecological environment in tropical waters (Degerfeldt and Hendry 1987). They

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form a natural shore protection, by absorbing wave energy and producing calcareous beach material such as broken ­corals, crustaceans and calcareous algae. The growth of coral reefs depends on non-turbid water and sea depth distribution. On the northern coast, the small rivers west of Annotto Bay carry little sediment. The shelf is narrow, and between Port Antonio and Montego Bay there is an almost continuous fringing reef (> Fig. 5.2.1). In Jamaica where shore-based tourism is of primary economic importance, a sustainable reef system is crucial. Emerged coral

reefs are found locally, as near Oracabessa on the north coast (> Fig. 5.2.2).

2.  The Jamaican Coast On the south coast there are several rivers with a considerable suspended sediment load. Some sections have a broad, somewhat deeper shelf. These factors make coral reefs discontinuous and patchy. Many of the reefs are in ⊡⊡ Fig. 5.2.1 A continuous fringing reef at the outer margin of a shallow platform and a narrow beach of derived coralline sand are found on parts of the northern coast.

⊡⊡ Fig. 5.2.2 Emerged, weathered coral reef east of Oracabessa.

Jamaica

poor condition because of sewage pollution, suspended sediment and excessive fishing. The distribution of fluvially-supplied sediment produces specific provinces of gravel and brown beach sand, in contrast to white beach sand produced by the coral reef systems. There is a brown sand province around the eastern part of the island from the Palisadoes (Kingston) to Annotto Bay (> Fig. 5.2.3), but the white coralline sand beaches around Morant Point form a major enclave in this province, derived from the coral and algal reefs of Morant Bay. White coralline sand derived from a fringing reef, has supplied the beach in Long Bay (> Fig. 5.2.4) on the northeast coast. Copacabana Beach has coarse sand and pebbles, as has the cove in Fairy Hill Bay, which receives gravel from a deep ravine. The fluvial sediment yield from Yallahs River has been reduced by the extraction of gravel upstream, resulting in shore erosion, while erosion of the alluvial fan at the mouth of the river has yielded boulders and gravel to adjacent beaches (> Fig.  5.2.5). Longshore drifting of  sediment supplied by the Morant River has formed a ­barrier beach. The north coast west of Annotto Bay and the west coast are other major provinces of white coralline sand, which reflect the presence of fringing coral reefs  and  only ephemeral rivers. The south-western and ­southern coasts have beaches with various mixtures of white  and  brown sand from rivers carrying sand and silt  and the less luxuriant coral reefs than on the north coast. Some beaches are backed by dunes, as at Thatch­­ field,  and sand has blown inland from Treasure Beach. ⊡⊡ Fig. 5.2.3 In Annotto Bay sand and gravel have been supplied by the Wag Water River, the mouth of which has been deflected towards the southeast by the growth of the barrier spit.

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Longshore drift has supplied sand and gravel to spits and tombolos, including the barrier spit that protects Kingston Harbour (> Fig.  5.2.6). Groynes have been inserted in an attempt to retain beaches, particularly at seaside resorts. In Jamaica lowland coasts are occupied by wetlands bordering the sea. There are four large coastal wetland areas called the Great Morasses: on the west coast the Negril Morass, on the south coast the Black River and the Portland Bight Morasses, and on the eastern end of the  island the Cape Morant Morass. On the south coast there are also many smaller wetlands, some lagoons and swamps occupying narrow swales between multiple beach ridges. On the north coast there are fewer, smaller wetlands often, degraded by exploitation. The wetland vegetation varies from thick forest to open swamp, and from fresh water to salt water communities, the latter dominated by mangroves. In the outer parts of the Black River and Negril Morasses the swamps are underlain by sedge and mangrove peat for many metres down to the substratum. Radiocarbon dating of peat from different levels has shown that that the swamp surface grew upward at the same rate as the global Holocene sea level rise (Digerfeldt and Hendry 1987), indicating isostatic stability inl these areas for at least the past 8,000 years. Irregular cliffs cut in the rock formations of the Blue Mountain foothills dominate the north-eastern Portland coast, with beaches between cliffy headlands. Another very irregular cliffed coast is found near Hanover, west of

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Montego Bay. Steep limestone cliffs protrude on several headlands along the northern and south-western coast, and there are prominent cliffs along Portland Ridge and the Hellshire Hills. There are abrasion shore platforms on the Negril cliffed coast (> Fig. 5.2.7). About 100 km SSW of mainland Jamaica, Pedro Bank extends over 2,400 sq. km within the 20 m depth contour.

Pedro Cays in the eastern part of the bank area are a group of four tiny coralline islets. The oldest reefs on the Bank form a peripheral strip along the northern edge and there is a swarm of patch reefs at the western end. These reefs are covered by at least 25 m of water, which suggests that they are drowned and support very little active coral growth. At the eastern end of the Bank a third set is

⊡⊡ Fig. 5.2.4 Long Bay on the northeast coast, has a white sand beach with a high berm, behind a fringing reef exposed to dominant easterly waves. (Courtesy L-U. Bergström.)

⊡⊡ Fig. 5.2.5 A wave-cut cliff on the Yallahs River alluvial fan, with a steep shingle beach and residual boulders.

Jamaica

­ ourishing. The present day shoals are covered with extenfl sive sheets of carbonate sediment, mainly composed of reef detritus and skeletal debris. The low cays are characterised by scrub and grass vegetation growing on coralline sand and gravel, and there are some small beaches with pebbles and sand. Pedro Bank is the largest and most productive fishing ground in Jamaica. The commercial categories include

⊡⊡ Fig. 5.2.6 The southern coast of the Palisadoes. The barrier spit protects Kingston harbour. (Courtesy L-U. Bergström.)

⊡⊡ Fig. 5.2.7 High terrace above steep cliffs. (Courtesy L-U. Bergström.)

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scale fish, lobster and conch. Parts of the cays are permanently occupied by fishermen, who usually stay for a month and return to the mainland for two weeks. Note: All material in this chapter was collected from our project on Integrated Coastal Planning and Management in Jamaica (1992– 1998), a co-operation between the National Resources Conservancy Authority, Jamaica and Uppsala University, Uppsala, Sweden.

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References Degerfeldt G, Hendry MD (1987) An 8,000 year Holocene sea level record from Jamaica: implications for interpretation of Caribbean reef and coastal history. Coral Reefs 5:165–69 Lindell LT (1999) Coastal zone mapping of Jamaica for planning and management. Proceedings, Pecora 14/Land Satellite Information III, Denver, CO

Norrman JO, Lindell LT, Bergström L-U, Mohlund ö, Nisell J (1997a) Manual for integrated coastal planning and management in Jamaica. CBA Internal Report No. 7 Norrman JO, Lindell LT, Bergström L-U, Mohlund ö, Nisell J (1997b) Coastal Atlas of Jamaica. CBA Internal Report No. 10 Woodley JD (1992) The incidence of hurricanes on the north coast of  Jamaica since 1870: are the classic reef descriptions atypical? Hydrobiologia 247:133–138

5.3  Hispaniola

(Haiti and the Dominican Republic)

  Anja Scheffers · Tony Browne  

1.  Introduction Hispaniola is a mountainous island about 660 km long and up to 260 km wide, with the Dominican Republic in the east and Haiti in the west. It has many features in common with the islands of Cuba and Jamaica to the west and Puerto Rico to the east, the island group forming the Greater Antilles. Four nearly parallel WNW-trending mountain ranges are separated by relatively narrow, alluvial-filled, longitudinal structural depressions (> Fig. 5.3.1). Coastal plains are also found, and are most extensive on the southeast coast of the Dominican Republic. In Haiti, where the mountains frequently come close to the shore, the area of coastal plain is relatively small. This account is based on a review by Alexander (1985). The island is located on the northern boundary of the Caribbean plate, and has been subjected to folding, faulting, uplift, and depression since the early Cretaceous (Bowen 1975). Raised coral reef terraces, all of Quaternary age, are found at a number of places along the coast, indicating that local uplift of up to several hundred metres continued at least well into the Pleistocene. Although geological studies of Hispaniola are relatively numerous there have been few accounts of the coastal geomorphology (Barrett 1962). As in Cuba there are sectors of steep, sometimes cliffed, coastline alternating with sandy beaches (some white and calcareous from reefs, others dark and of volcanic origin), and mangroves are extensive on low energy shores, particularly behind shallow coralline seas that are bordered by reefs and island chains. Much of the coast is fringed by marine Quaternary deposits, including emerged coral limestone and associated calcarenites. There is widespread evidence of terraces indicating phases of higher relative sea levels and tectonic uplift in Quaternary times. Hispaniola has a humid to subhumid tropical climate. The north coast is exposed to northeasterly trade winds that bring heavy rainfall in the summer months, but the

south coast is drier. Wave action is strong on the north coast when the trade winds blow onshore, and moderate on the south coast which receives waves generated across the Caribbean Sea. The island coast is also subject to occasional hurricanes, especially in August, September and October. Tide ranges are small: Jacmel on the south coast of Haiti has a mean spring tide range of 0.6 m. The marine provinces bordering Hispaniola are rich in calcareous biogenic sediment, notably coralline sand and gravel, but beaches on the more exposed parts of the coast have been eroding in recent decades.

2.  The Coast of Haiti This discussion of coastal features will start at the international boundary on the north coast and work anticlockwise around the island, first in Haiti, then the Dominican Republic. From the Dominican border to Cap Haitien the coast is backed by a broad alluvial plain, the landward limit of which is a bevelled rock platform possibly formed by marine erosion. The platform ranges in elevation from 150 m in the east to 75 m in the west. Between Cap Haitien and the Baie de 1’Acul the coast is steeply cliffed on Tertiary sedimentary formations. Coral reefs are generally abundant in sheltered bays and bay mouths. Coastal inlets have formed where rivers incised through an emerged fringing reef during Pleistocene low sea level phases. East of Cap Haitien a 20 km long reef stands a little offshore, enclosing a mangrove forest. West of Baie de 1’Acul is a 10 km section of tombolo coast with long, curving beaches. Offshore is Tortue Island, 37 km long and up to 7 km wide. Its entire north coast is dominated by high sea cliffs and is inaccessible. There are two terraces, best preserved at each end of the island, where the slope is relatively gentle. On the southern side the terraces have been virtually destroyed by marine erosion.

This is a revised version of a chapter by C.S. Alexander in The World’s Coastline (1985:181–185). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Hispaniola (Haiti and the Dominican Republic)

⊡⊡ Fig. 5.3.1 Coastal features of Hispaniola. (Courtesy Geostudies.)

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A remarkable set of well-preserved Quaternary terraces occur around the western end of the northwest ­peninsula. They increase in number and size westward from Port-de-Paix, eventually extending to a limestone plateau 640 m high. Successively higher terraces are veneered with progressively older limestones. There are at least 20 terraces between Bombardopplis and Baie de Henne, and they diminish again eastward to Port-a-Piement. There are at least eight high, prominent terraces leading up to the top of the plateau. Uranium series dating of corals from the crests of the three lowest terraces of this series yielded mean dates of 81,000 ± 3,000 years bp for the Mole terrace (16 m), 108,000 ± 5,000 years bp for the Saint terrace (28 m), and 130,000 ± 6,000 years bp for the Nicolas terrace (52 m). Assuming sea level was 6 m higher than at present during the formation of the Nicholas reef, the local uplift rate is calculated to be 0.35 m/1,000 years (Dodge et al. 1983). On the southeastern side of Cap St. Marc there are four prominent limestone terraces which show evidence of crustal tilting. On the northern side of the promontory, two terraces skirt the shore of St. Marc Bay. Except at the head of the bay, the shore is bordered by an almost continuous cliff 20–30 m high. The alluvial lowlands along the west coast of Haiti have flat, marshy shores with fringing mangroves. One of these lowlands, the Cul-de-Sac, has several coral limestone terraces along its inland margins, deposited when a Quaternary sea extended more deeply into the lowland. Subsequently, the terraces appear to have been tilted. The

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Cul-de-Sac is a rift valley between parallel faults, with Port-au-Prince on its southern side. Gonave Island is about 57 km long and up to 15 km wide. Its shores are almost entirely bounded by coral reefs except for the high cliffed section along the northwestern end. Steep cliffs are also found on the southeastern coast of the island and low mangrove shores lie between the cliffed sections. Ten or more well-preserved limestone terraces occur on the north and west coast. West of Port-au-Prince there are cliffs which give place near Grand-Goave to a zone where rapid alluvial deposition dominates the coast, and much of the shore is lined by mangroves and mudflats. A single, slightly arched terrace occurs from Miragoane to slightly west of Jeremie. This north coast of the Jacmel peninsula is cliffed for long sections, the cliffs ranging from 10–18 m in height, ­interrupted by submerged karstic features. Coral reefs are relatively rare and, along with mangroves, are mainly confined to sheltered locations. From Pointe-Dame Marie round to Tiburon a number of small bays are separated by cliffed headlands. The cliffs are occasionally more than 100 m high. At Grante Pointe there is a 10 km long gravel beach 3–5 m high, behind which is a low gravel terrace. The southwest coast of Haiti coast curves out to Point l’Abacou, beyond which are a number of small bays with islands with some detached coral reefs. Between Aquin and Jacmel there are at least two marine terraces, one of which shows evidence of differential uplift. There is also evidence at Jacmel of a third high sea level stand.

Hispaniola (Haiti and the Dominican Republic)

The south coast of Haiti is mostly bordered by steep cliffs. Between Jacmel and the Dominican Republic boundary the cliffs are occasionally 200 m high as the mountains drop to the sea. Coral reefs are relatively rare and ­mangrove growth is confined to a few protected bays.

3.  The Coast of the Dominican Republic From the Haitian border on the south coast of the island cliffs cut in Tertiary limestone run south-east to Cape Beata. There are three terraces in the vicinity of Cabo Beata. The lowest has an elevation of about 40 m, the second lies between 90 and 250 m, and the third has an elevation of about 400 m. Weyl considered that the terraces indicate the extent of Pleistocene uplift in the area. East of Cape Beata the coast trends northeastward, and passes into cliffs in Tertiary limestone. There are low sectors where Quaternary deposits occupy basins east to Punta Palenque. From Punta Palenque to Cabo Engano the coast is almost continuously cliffs of disintegrating limestone, the cliff height varying from less than a metre to about 18 m ⊡⊡ Fig. 5.3.2 Crumbling cliffs of limestone at Boca de Yuma on the southeastern coast of the Dominican Republic near Cabo Espada. (Courtesy W. Barrett.)

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(> Fig.  5.3.2). Nearly everywhere the cliff is topped by a low ridge of calcareous debris that may obscure the cliff in low sectors. Where prominent, the cliff can be divided into two general levels, a submerged level at about 4 m and an emerged level or secondary cliff at about 3–4 m (Barrett 1962). In places the mountains come close to the sea. Barrett (1962) studied the marine terraces along the southeast coast of Haiti between Punta Palenque and Cabo Engano. The inner edge of the coastal plain lies at an elevation of about 120 m, but clearly defined emerged coast features are absent above 78–82 m. Below 78–82 m are numerous marine-eroded scarps and benches but none are continuous for long distances. However, the scarps and benches can be grouped into a number of ­composite terraces, the most important of which is between 29 and 32 m, extending from Punta Palenque to the bay east of La Romana. The terrace surfaces do not appear to have been warped or tilted and are thought to be of eustatic origin. A variety of evidence from overdeepened river valleys, sand dune remnants, and submerged marine terraces suggests a lowering of sea level by at least 150 m (Barrett 1962). Uranium series dating of corals from the 3–6 m and 8–9 m terraces indicate the age of the lower to range between 95,000 and 120,000 years bp,

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whereas one sample from the upper gave an age of 190,000–210,000 years bp. From Cabo Engano to Cabo San Rafael the coast is low and is bordered by an intermittent coral reef. Much of the shore consists of a calcareous sand beach backed by a low dune, both of which rest on Cretaceous limestone (Barrett 1962). Marine terraces border the south coast of the Bahia de Semana and near Cabo Samana, and there are coral reefs in the bays. The north coast of the Samaria peninsula is steep and rocky with coral reefs, and towards Cabo Francis Viejo the coast is low and bordered by a fringing reef. The Cape has a series of raised coral reefs up to at least 360 m above sea level. In the broad bight between Cabo Francis Viejo and Puerto Plata coral reefs extend as much as 12 km offshore. There are some marine terraces the vicinity of Cabo Francis Viejo and along the coast to Puerto Plata. Sandy and pebble beaches occur at intervals along the north coast of the Dominican Republic west of Cabo Francis Viejo, where the coast is backed by mountainous terrain. At Punta Mangle the coast curves southward and is low, with an almost uninterrupted mangrove fringe as

far as the Haitian border. There is a 1–2 km wide reef platform, separated from the mangrove belt by a reef lagoon several hundred metres wide. Beach monitoring in the Dominican Republic has shown depletion of beaches by storm waves during Hurricane Hugo in 1989 followed by a phase of accretion, then further erosion during later hurricanes (Cambers and James 1994).

References Alexander CS (1985) Hispaniola. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, Stroudsburg, pp 181–185 Barrett W (1962) Emerged and submerged shorelines of the Dominican Republic. Instituto Pan Americano de Geographie e Historie. Geogr Rev 30:51–77 Bowen C (1975) The geology of Hispaniola. In: Nairn EM, Stehli EG (eds) The Ocean Basins and Margin, 3: the Gulf of Mexico and the Caribbean. Plenum, New York, pp 501–522 Cambers G, James A (1994) Sandy coast monitoring: the Dominican example 1987–1992 UNESCO Reports in Marine Sciences 63 Dodge RE, Fairbanks RG, Benninger LK, Maurrasse E (1983) Pleistocene sea levels from raised coral reefs of Haiti. Science 219:1423–1425

5.4 Puerto Rico

Jack Morelock ∙ Wilson Ramirez ∙ Maritza Barreto

1.  Introduction Puerto Rico is 62  km by 178 km, with an area of 9,104  sq.  km and a 501 km coastline (Morelock 1978). There are more than 150 beaches with a total length of 208 km, or 41% of the coastline. The beaches are relatively short, and separated by headlands into distinct beach systems that are closed or semi-closed units receiving sediment from limited local sources and transmitting little sand to other beaches. The geology of the Island is complex, as is to be expected from its formation at the leading edge of the Caribbean Plate. Puerto Rico occupies part of a platform that includes the Virgin Islands. The basic pattern is a volcanic and plutonic central mountain core trending west to east with thick limestones to the north and south (Kaye 1959). The volcanic core has steep, highly eroded ridges separated by deep valleys. The limestone flanks form a series of ridges and valleys that are related to fault trends; they have karst topography and end in steep cliffs along the coast. The coastal plains are at the mouths of river valleys and below colluvial fans from the mountains; they are fringed along the coast by beaches, dunes, beach rock and dune calcarenites. River discharge, rainfall and vegetation cover are greater on the north coast, which has most of the major rivers. The offshore sediment of the narrow (2–4 km wide) north and northwest insular shelf are dominantly terrigenous. The major protection from wave action is provided by remnants of former dune calcarenite lines. Off the semi-arid south coast, the shelf varies in width from 1 to 10 km and has numerous coral reefs. The southwest and eastern areas of the shelf are relatively wide and are predominantly coral and carbonate sediment areas (Morelock 1978). The west and east coasts of Puerto Rico are cut across the pattern of folding and faulting, and the geomorphology reflects both structural trends and rock types. Rocky limestone and volcanic coasts alternate with alluvial deposits,  beaches and mangrove shorelines. The north coast is formed by Tertiary limestones with extensive karst topography and narrow alluvial plains in front of the mountains, where beaches and dunes have developed, in places ­lithified

to beach rock and dune calcarenites. Large areas of mangrove are present, but they are fronted by beaches, spits and barriers. The south coast has similar rock types except for intrusive igneous rocks at the eastern end, and a large area of alluvial fan deposits. Because of lower rainfall, karst erosion on the limestones in the western half of the south coast is much less than on the north coast. The beaches of Puerto Rico contain sand derived from several major sources (Morelock 1978). Transport of rock material by rivers to the beach has supplied terrigenous sand, the composition of which depends on catchment geology. Outcrops of volcanic rocks, notably basalt, have supplied dark minerals and dark rock fragments; weathered granite has yielded quartz and feldspar sand. Beach sediment derived from the erosion of cliffs and shore outcrops are also influenced by the nature of rock outcrops. Sand and gravel swept in by wave action from the sea floor include calcium carbonate supplied to the beach by the shoreward transport of the shells of marine organisms, as well as material eroded from limestones and dune calcarenites. The composition of particular beaches depends on the proportions of these various source materials: near river mouths there is an increase in the terrigenous content of beach sands, and there is also a shift toward terrigenous beaches where an offshore carbonate source is lacking. Storms and hurricanes are the major agent of geomorphic change on the coastline of Puerto Rico (Morelock et al. 2000; Donnelly 2005). A day of storm activity can shape features that would take years to form under normal conditions. The distribution of wave regimes varies around the Island, with changes more rapid on exposed sectors of coastline. Although periods of more than 2 m/year ­erosion have occurred locally, most of the coastline shows alternations of erosion and accretion, with a trend toward moderate erosion of unconsolidated deposits (Liegel 1972).

2.  The Coastline of Puerto Rico The northwest coast near Punta Borinquen is a series of rocky karstic limestone cliffs receding as the result of wave erosion. Where there is a narrow coastal plain in front of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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the limestones, sand from the beach is being incorporated into backing dunes. The west coast of Puerto Rico has the terminations of east-west mountain ridges separated by the broad alluvial valleys of the Culebrinas and Guanajibo rivers (Kaye 1959). The ridges form rocky coastline, and the coastlines of the alluvial valleys are fringed by sandy beaches. The insular shelf is very narrow from Aguadilla to Mayaguez. Much of this coastline receives NW waves refracted from swell generated by north Atlantic storm activity. To the south, headlands alternate with beaches, as at Punta Higuero. Here the beach has sands with almost equal parts of carbonate sand from the shelf and quartz and feldspar, dark minerals and dark volcanic rock fragments from the adjacent coast and river sediment. The sand composition is similar south to Mayaguez, although there is an increase in dark minerals and rock fragments southward. The largest estuary on the west coast is Río Anasco. At Mayaguez the submarine canyon of the Río de Yaguez separates the beaches to the south and north; the  composition changes drastically. The south beach is ­composed of igneous rock fragments and dark minerals, mainly serpentine. The absence of carbonate shell material indicates almost no onshore transport of sediment. South of Mayaguez, the shelf is relatively broad and reefs and shoals with depths of less than 10 m extending 20 km offshore. The pattern of ridges ending in rocky headlands and valley-mouth bays with sandy beach continues southward along the coast, but there are also mangrove shores and fringing reefs because of the protection

offered by the broad shelf and irregular shoal reef bathymetry, which results in reduced wave energy. The beaches south of Punta Guanajibo are of carbonate with quartz and some igneous rock fragments. Punta Arenas has a short beach of carbonate sand (>Fig. 5.4.1) near to the Laguna Joyuda on a coastal plain fronting volcanic hills. From Punta Ostiones south, the sands are largely carbonate, as on bay-head Boqueron Beach, with some quartz derived from rivers draining the alluvial plains east of Boqueron. At the western end of the south coast of Puerto Rico is the Cabo Rojo tombolo, where a low sandy isthmus with a wide beach on its western shore runs out to a headland of Cretaceous limestone (>Fig. 5.4.2). On the south coast of Puerto Rico east of Punta Cabo Rojo the coastline has a narrow fringe of mangroves to Punta Montalva, with low tidal flats and salt ponds to the rear. The irregular cliffy coast east of Punta Montalva is cut in limestone, with small pocket beaches of shelly sand and gravel at the base of the cliffs. Karstic weathering is present, but inhibited by the low rainfall. East from Punta Verraco wave erosion has smoothed sectors of low depositional coast, which are beach-fringed with some low cliffs of alluvial sediment. From Ponce eastward, the coast is a low-lying alluvial plain. There are almost continuous beaches, interrupted by short sectors of mangroves, eroding alluvial plain and rock riprap placed to counter severe erosion. The beach sediment is primarily igneous rock fragment and dark mineral. Many of these beaches have gravelly material in the surf zone, as at Manzanilla (>Fig. 5.4.3). ⊡⊡ Fig. 5.4.1 Carbonate sand beaches at Punta Arenas.

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⊡⊡ Fig. 5.4.2 Punta Cabo Rojo is a block of Cretaceous limestone connected to the mainland by a low sandy isthmus to form a tombolo.

⊡⊡ Fig. 5.4.3 The beach at Manzanillo is lag gravel derived from an alluvial fan.

From Punta Las Marias to the east, the coastline is a low-lying wave-trimmed alluvial plain south of the central mountains. There are extensive narrow beaches of sand and gravel at the base of low cliffs. Mangroves occupy the shore in Bahia de Jobos, where wave energy is low behind a wide zone of coral reefs. At Cabo Mala Pascua granodiorite comes to the coast, and there are small valleys with alluvial fans between headlands. The beaches near the mouth of Río Maunabo are rich in quartz sand.

East of Arroyo is a rocky coastline with small pocket beaches. To the east, the beach sands show an increase in the amount of quartz and feldspar, as well as magnetite and other heavy minerals derived from the large southeast granodiorite batholith. From Arroyo to Naguabo rocky headlands alternate with a few small and narrow beaches. The east coast of Puerto Rico is bordered by a shallow shelf with abundant coral and marine organisms which generate carbonate shelf sands. As a consequence, the

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beaches are again a mixture of carbonate sand from the sea floor and of quartz and feldspar with igneous rock material from alongshore and hinterland sources. All the beaches have relatively high quartz content, reflecting drainage from the weathered igneous rocks of the eastern part of Puerto Rico, but there are varying proportions of calcareous sediment from the sea floor. Morphologically, the east coast of Puerto Rico differs from the west coast in that it is less indented and has broader headlands. Large rocky headlands of plutonic and volcanic rock are located between broad alluvial plains. Small pocket beaches are seen at the base of the cliff in some places. Extensive alluvial plains have been formed by  the major river systems at Naguabo, Humacao and Yabucoa, and there are long beaches between Punta Fraile and Morro de Humacao. Bahia Las Cabezas and Las Croabas have narrow carbonate beaches lying behind a fringing reef which supplies most of the beach material. There are igneous rock fragments in the Las Croabas beach, derived from local sources. Fringing reefs line the coast between Fajardo and Cabo San Juan. The north coast of Puerto Rico is exposed to Atlantic swell and storms arriving across a narrow insular shelf, and has higher wave energy than elsewhere in Puerto Rico. From Cabo San Juan (>Fig.  5.4.4) in the northeast, the coast westward to San Juan is generally low-lying, with beaches and barrier spits at the mouth of the Río de Loiza. The coastal area is low and sandy except for occasional

dune calcarenite bluffs, as at Puerto Vacia Talega. The lowland extends inland for up to 7 km. Several large mangrove areas lie several hundred metres inland behind sandy beaches and barriers. The beach sediment is predominantly carbonate sand from the sea floor, except between Punta Uvero and Punta Vacía Talega, where the beach receives a significant input of river sediment. Quartz content is especially high because the upper drainage basin of the Río Loiza is an area of weathered igneous rock (Gillou and Glass 1957). At the port of San Juan the beaches are thin coverings of sand over a rocky shore. West of San Juan are dune calcarenite cliffs at Punta Maldonado and a large tombolo at Punta Salinas (>Fig. 5.4.5). Much of the north coast is a low-lying coastal plain. There are sandy beaches and dunes with occasionally rocky dune calcarenite coast and beach rock coast west from San Juan. There is a sand spit at the mouth of the Río de la Plata, and to the west is an irregular rocky dune calcarenite coast, which has been breached in many places to form small lunate bays bordered with sand and beach rock (>Fig.  5.4.6). At Porto Nueva a breached headland has been reunited by sand deposition to form a small tombolo. Except for the locally terrigenous sediment supplied by the Bayamon and La Plata rivers, the beach sands are largely carbonate from San Juan to the Río Manatí. West from the mouth of the Río Manatí, the beach sands are igneous rock fragments, dark minerals, quartz,

⊡⊡ Fig. 5.4.4 Cabo San Juan is a headland of volcanic rock with small sand beaches.

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⊡⊡ Fig. 5.4.5 Punta Salinas is a tombolo that has been eroding for the past five decades. Riprap has been placed to counter erosion.

⊡⊡ Fig. 5.4.6 A segment of beach rock at Cibuco indicates the former coastline, and thus the extent of subsequent beach retreat.

and some feldspar. There is a slow decrease in terrigenous material westward until, near the port of Arecibo, the sediment is again carbonate. Sand from these beaches is being delivered to dunes behind the beaches by onshore winds.

The dunes near Arecibo are well vegetated. The beach was sandy into the surf zone until 1999, when storm waves exposed beach rock. Frequent winter storms overwash the dunes carrying gravel on to them. There are karstic cliffs

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west of Camuy, and the dunes west of San Jacinto have been reduced or removed by sand quarrying. Cliffs of Tertiary limestone extend along to the headland at Punta Borinquen.

3.  The Islands of Puerto Rico More than 45 small islands and islets are scattered around Puerto Rico. Some like Isla Magueyes are less than 50 m from the island of Puerto Rico others are 10 (Caja de Muertos) to 100 km (Mona) offshore. The largest three of these are discussed in this section. Vieques and Culebra are populated, Mona is under the supervision of the Department of Natural and Environmental Resources and open to visitors. Isla de Mona, to the west of Puerto Rico (18° N, 68° W) is 7 km by 11 km, and Isla Monito, about 5 km northwest of Mona, is comparatively tiny, only 40 acres in area. Mona and Monito are different from the rest of Puerto Rico. They appear as great flat white slabs of limestone floating on the sea and bounded by high cliffs rising vertically out of the water. They rest on two upthrown structural blocks delineated by several near-vertical faults. On the west coast of Mona are limestone cliffs and fringing reefs. Only on the more protected sides of these islands is the sea sufficiently shallow for the development of coral reefs. On the southwest coast of Isla de Mona they protect a narrow coastal plain partially fringed with sandy beaches. The coastal plain narrows eastward to limestone cliffs, interrupted by low plains east and west of Pajaros, where there is a beach ridge plain and a Pleistocene marine terrace. Reef terraces of late Pleistocene age form the coastline along 14 km of the south and west coasts of Isla de Mona. Beach sands and beach rock extend from Playa Sardinera SE to Playa del Uvero, and from Punta los Ingleses NE to Playa de Pajaros, and slightly beyond. The rest of the coastline is dominated by limestone cliffs. There are caves high in the cliffs at the contact between the grey Lirio Limestone and the underlying buff-colored Isla de Mona dolomite.

The cliffs continue underwater for more than 25 m and shallow sea caves have formed at water level. Isla Vieques (18° 10' N, 65° 20' W) is 13 km off the east coast of Puerto Rico, and measures 5 km by 29 km. The island has Tertiary limestones, Cretaceous volcanics and plutonic intrusive rocks with alluvial plains, beaches and mangrove swamps. Green Beach at the western end is located on a sandy point from which a large shoal submarine sand body extends NW. On the southwest coast, rocky headlands separate pocket beaches and on the south coast curving beaches front mangrove-fringed lagoons. The north coast is rocky, with bordering beaches, some of which are partially protected by fringing reefs. West of Isabel Segunda, rocky headlands separate pocket beaches on the narrow coastal plain. The island of Culebra (18° 20' N, 65° 20' W) is a dissected, partly submerged volcanic island with rocky and beach-fringed coasts and fringing coral reefs. Dakity has a fringing reef on an island at the entrance to Ensenada Honda, a mangrove and rock-lined lagoon on the south coast of Culebra. The north coast is rocky with a few beaches, but on the northern peninsula Flamingo Beach is a striking beach of fine white carbonate sand.

References Donnelly JP (2005) Evidence of past intense tropical cyclones from backbarrier salt pond sediments: a case study from Isla de Culebrita, Puerto Rico. J Coastal Res (Special Issue) 142:201–210 Gillou RB, Glass JJ (1957) A reconnaissance study of beach sands of Puerto Rico. Bull US Geol Surv 1052–1:273–303 Kaye CA (1959) Shoreline features and quaternary shoreline changes, Puerto Rico. US Geological Survey Professional Paper 317-B Liegel LH (1972) NCERI beach-offshore study of mainland Puerto Rico. Puerto Rico Area of Natural Resources, Department of Public Works Morelock J (1978) Shoreline of Puerto Rico. Puerto Rico Coastal Zone Management Program, Department of Natural Resources Morelock J, Capella J, Garcia JR, Barreto M (2000) Puerto Rico – Seas at the millennium. In: Sheppard CRC (ed) Seas at the millennium. Oxford University Press, London

5.4.1 Virgin Islands

Maurice Schwartz

1.  Introduction The Virgin Islands lie at the eastern end of the Greater Antilles. Some, including St. Thomas, St. John and St. Croix, belong to the United States; others, such as Tortola, Virgin Gorda and Anegada, are British. There are high islands with bay beaches and coral reefs rising from the Virgin Island Bank, a submarine platform that runs out eastward from Puerto Rico. St. Croix, an outlier to the south, is located on a submarine ridge south of the Anegada Trough (Mattson et al. 1990). There are also low coralline islands and reefs. Details of one island, St Thomas, are given here.

2.  St. Thomas St. Thomas has an area of 83 sq. km. Rock formations ranging from Cretaceous to Eocene include submarine lava flows, andesite breccias and tuffs, greywackes with blocks of limestone and volcanic rocks, and lenses of foraminiferal limestone. A large batholith intruded during the mid- to late-Eocene was accompanied by tectonic deformation and metamorphism (Weaver 1975; Mattson et al. 1990). There has been faulting and northward tilting, followed by dissection resulting in high ridges with steep slopes, trending mainly NW–SE. There is a wet season from May to November, with NE trade winds prevalent in June and July. This is generally followed by a dry season during the rest of the year, with SE trade winds prevalent from December through March. The trade winds produce high wave energy on eastern coasts, while western coasts remain sheltered. Occasionally, easterly swell generated by Atlantic winter storms reaches the shores of the island. Tropical low-pressure systems form off the west coast of Africa and move westward across the Atlantic Ocean to become tropical storms and hurricanes as they approach the Caribbean Islands. The hurricane season lasts from early June to late November, most occurring in September. The North Equatorial Current coming westward from the Atlantic gives water temperatures of at least 25°C. Mean spring tides are small (0.3–0.5 m) and semidiurnal.

With 85 km of coastline, fringing and barrier reefs on St. Thomas are located predominantly along the northern and eastern shores of the island. A mid-shelf reef complex lies to the south. There are over 40 species of scleractinian corals and 3 species of Millepora in the region; with welldeveloped reefs dominated by Montastraea annularis oc­­ curring at mid-shelf depths of 33–47 m. Surveys have shown that coral reef growth around St. Thomas improves progressing from the inner fringing reefs to the outer mid-shelf reefs. Complicit in this trend are runoff from the land bringing increased pollution and sediment to the shore regions, the Charlotte Amalie harbour being the most impacted of all (Hubbard et al. 2008; Rogers et al. 2008). Added to this around the island is the damage to reefs done by dragging of  anchors and grounding of vessels, careless snorkelers and scuba divers, and overfishing. On a larger scale, hurricanes  and higher sea water temperatures are an added stress  to living coral. Bleaching and disease (black band, white  band, white plague, white pox) also take a toll (Drayton et al. 2005).

3.  Beaches St. Thomas has a number of fine beaches. The sequence described here starts on the southwest coast of St. Thomas and progresses in a counterclockwise fashion around the east end, along the north coast and then ends up at the western tip of the island. Brewers Bay, located near the University of the Virgin Islands, has clear, calm water along a lengthy sand beach. Emerald Beach fronting Airport Road is also known as the Lindberg Bay Beach. Though broad and sandy, in stormy times the beach face is cut by low scarps. To the east, across Charlotte Amalie, the Morningstar Bay beach has fine sand at the upper foreshore but coarse sand and pebbles in the sloping surf zone. Limetree Beach is located in a cove within Frenchman Bay. The sandy beach gives way to a rocky substrate in shallow water. A small stream crosses the beach at one end. While the Bolongo Bay beach is composed mostly of sand and pebbles, beach rock outcrops in places on the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.4.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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foreshore. Along one side of the bay, the shore is rocky. Nestled between Breverhout Point and an adjoining headland, Secret Harbour has a sandy beach and a calm wave climate. The shoreline is fairly close to the base of a restaurant building and palm trees, hinting at some beach erosion over the years in this protected embayment. The beach at Cowpet Bay is sandy with some pebbles and coral fragments. While the backshore is lined with palm trees, rocky areas bracket the ends of the beach. The water is known to be calm and clear. Bluebeards Beach and Turtle Cove are sandy with calm water just offshore. Sea grass is common at both sites. Vessup Beach is at the SW end of a long embayment. Ranging from narrow to wide, the sandy beach has outcrops of beach rock exposed on the lower foreshore (>Fig. 5.4.1.1). Sapphire Beach is sandy with some pebbles, and driftwood occasionally accumulates at the backshore. Shallow reefs on one side and an artificial pond on the other bracket the beach. Lindquist Beach is also sandy, with outcrops of beach rock at intervals (>Fig.  5.4.1.2). A dense growth of mangroves backs some sections of the beach. The narrow beach at Sugar Bay has trees growing at the  upper swash limit and slight scarping. Patches of ­different  sand sizes along its length indicate possible replenishment. The beach at Water Bay, also known as Pineapple Bay Beach and Renaissance Beach, is composed mostly of sand with some small pebbles. Sea grass is abundant offshore

and often litters the beach. Coki Beach is small, with clear calm surf and a nearshore reef. The Mandahl Bay beach is composed of sand, gravel, and coral fragments; there are small tide pools, and some rocky shore areas. Two rock jetties have been constructed at the east end of Mandahl Beach beside a channel into the lagoon. The most noted beach on St. Thomas is that at Magens Bay, deeply embayed between the long, NW–SE trending, wooded headlands of Tropaco Point on the west and Peterborg Point on the east. Magens Beach is a 1.5 km long pocket beach, with fine sand, a gently sloping nearshore bottom: there is no longshore current. Somewhat to the west of Magens Bay there is a small beach at Hull Bay. There is a boat ramp crossing the foreshore at the center of the beach, where sediment has accumulated on the east side and slight erosion has occurred on the west. The beach at Dorothea Bay, embayed between two small, rocky headlands, has a long northeasterly fetch that results in strong wind and wave action (>Fig. 5.4.1.3). The beach sediment mixture of sand and gravel is often formed into large beach cusps in the pounding surf. Neltjeberg Beach a long, broad, expanse of sand with gentle surf. Extensive reefs lie offshore. The west end of St. Thomas has beaches in two coves separated by a promontory, Botany Bay and Sandy Bay respectively. These sandy beaches at the western end of St. Thomas are sheltered from the prevailing northeasterly and southeasterly trade winds, providing a gentle wave environment.

⊡⊡ Fig. 5.4.1.1 Beach rock on the lower foreshore at Vessup Beach.

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⊡⊡ Fig. 5.4.1.2 Lindquist beach.

⊡⊡ Fig. 5.4.1.3 Beach cusps at Dorothea Bay beach.

References Drayton N, Rogers C, Devine B (2005) The state of coral reefs of the U.S. Virgin Islands. The Ocean Conservancy, Washington DC, 58 p Hubbard DK, Burke RB, Gill IP, Ramirez WR, Sherman C (2008) Coralreef Geology: Puerto Rico and the US Virgin Islands. In: Riegl, Dodge RE (eds) Coral reefs of the USA: coral reefs of the world, Vol 1. Springer, Berlin, pp 263–302 Mattson P, Draper G, Lewis JF (1990) Puerto Rico and the Virgin Islands. In: Dengo G, Case JE (eds) The Caribbean region: the geology of

North America, Vol H. Geological Society of America, Boulder, CO, pp 112–120 Rogers CS et al (2008) Ecology of coral reefs in the US Virgin islands. In Riegl BM, Dodge RE (eds) Coral reefs of the USA: coral reefs of the world, Vol 1. Springer, Berlin, p 303–373 Weaver JD (1975) Virgin Islands. In: Fairbridge RW (ed) Encyclopedia of world regional geology part 1: western hemisphere. Dowden, Hutch­ inson and Ross, Stroudsburg, pp 654–656

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5.5  Lesser Antilles

Gillian Cambers

1.  Introduction The Lesser Antilles represent an arc of island systems separating the Caribbean Sea from the North Atlantic Ocean. They extend from the Virgin Islands (latitude 18° 30' N) in the north to Trinidad (latitude 10°N) in the south (Deane 1985). The arc of islands lies along the eastern boundary of the Caribbean Tectonic Plate, which lies adjacent to the North American Plate, which in turn is moving westwards and subducting beneath the Caribbean Plate. The Lesser Antilles consist of two main arcs: an inner active volcanic arc and an outer calcareous arc. The inner volcanic arc is of Eocene origin and extends from Saba in the north to Grenada in the south. Many of these islands have experienced recent volcanic eruptions, as at Mount Pelée in Martinique in 1902, Soufrière in St. Vincent in 1979, and Soufrière Hills in Montserrat in 1995 where the volcanic activity is continuing into the 21st century (>Fig.  5.5.1). The submarine volcano, Kick ‘em Jenny, 8 km north of Grenada, rose from a depth of 200 m

to 140 m below sea level between 1974 and 1989 and may emerge as the next Caribbean island (Hendry 1996). The outer arc, which has been in existence since the early Tertiary, stretches from Sombrero past Anguilla to Barbados in the south. The islands of this arc are much lower and have less rugged topography than those of the inner volcanic arc. Coral reefs have developed around many of these islands and the extent and type of reef is influenced by aspect, substrate and hurricane influence (Wells 1988). Reef development is usually more extensive on the small low islands with low rainfall and little sedimentary runoff, as in Anegada, Antigua, Barbuda and the southern Grenadines. St. Vincent and the Grenadines and the US and British Virgin Islands are archipelagos, and many of the other territories consist of several islands and cays, e.g., Anguilla. The > Table 5.5.1 lists the island territories, located in (>Fig. 5.5.2). The climate is tropical, with waves generated by the Northeast Trade Winds, which are strongest from June to July and from December to March. Waves generally

⊡⊡ Table 5.5.1 Island Territories of the Lesser Antilles Political grouping

Outer chain (Coralline)

Inner chain (Volcanic)

Independent Island States

Antigua and Barbuda Barbados

Dominica Grenada St. Kitts and Nevis St. Lucia St. Vincent and the Grenadines

Departments of France

St. Martin

Guadeloupe Martinique St. Barthélemy

Netherlands Antilles

St. Maarten

Saba St. Eustatius

UK Dependent Territories

Anguilla

Montserrat

UK Overseas Territories

Other Trinidad and Tobago

British Virgin Islands U.S. Virgin Islands

A revised version of a chapter by Compton Deane in The World’s Coastline (1985):193–199. With contributions by Maurice Schwartz (St. Martin) and Pierre Saffache et al. (Guadeloupe). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 5.5.1 Since the 1995 eruptions of the Soufrière Hills volcano, the southern part of Montserrat, including the former capital Plymouth, has been abandoned. (Courtesy Angela Greenaway.)

approach from directions between northeast and southeast. The east facing coasts are known as windward coasts and experience higher wave energy than the west-facing leeward coasts, which are more sheltered. Between the months of October and April, the Lesser Antilles are periodically affected by swell waves generated by intense midlatitude storms in the North Atlantic Ocean. During each swell season, there may be five to ten swell events each lasting one to eight days and these bring high swell waves to east-, north- and west-facing coasts. One such event in March 2008 resulted in swells more than 10 m in height and caused at least one drowning death in the Lesser Antilles. Between June and November, the Lesser Antilles are affected by tropical storms and hurricanes. In general, these low pressure systems form towards east of the island arc and travel in a westerly or northwesterly direction, although there have been exceptions: Hurricane Lenny in November 1999 formed off the coast of Nicaragua and moved eastwards across the Caribbean Sea to impact the Lesser Antilles. These low-pressure systems bring intense rainfall, high winds and storm surges to the islands of the Lesser Antilles causing significant coastal and beach changes, environmental damage, flooding, loss of life and damage to island infrastructure. During the last century, there have been multi-decadal fluctuations in hurricane

activity in the Atlantic Basin and Caribbean Sea, with the period 1970–1994 showing low hurricane activity, and the period since 1995 showing increased activity. The record is insufficient to indicate whether these fluctuations are a natural climate variation or linked to climate change (IPCC 2007). The period 1995–1999 saw intense hurricane activity in the northern islands of the Lesser Antilles, while hurricane activity moved to the southern islands (Grenada and St. Vincent and the Grenadines) in the years 2004–2005. Tides are semidiurnal and the range is small, around 0.3 m and 0.6 m at Spring tides. Coastal landforms include cliffs, rocky slopes, dunes, beaches, wetlands and salt ponds, beach rock ledges; offshore there are coral reefs, seagrass beds and sandy bottoms. Beach composition varies from large granite boulders to giant coral heads, and from black volcanic sand to white coralline sand. Most of the ports and tourism infrastructure are located close to the west coasts of the islands in the Lesser Antilles. These are the more sheltered coasts and favoured for tourism development. In many places, this infrastructure interrupts coastal processes and results in localised erosion and coastal changes. Selected islands are considered in more detail in this chapter to illustrate the range and variety of coastal geomorphology in the Lesser Antilles.

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⊡⊡ Fig. 5.5.2 Map of the Lesser Antilles. (Courtesy Caribbean Conservation Association.)

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2.  Anguilla Anguilla, one of the most northerly of the outer arc islands, lies on the extensive Anguilla Bank. The island is low-lying and composed of a Lower Miocene limestone and marl series overlying volcanic rocks (Christman 1953). There are several small uninhabited cays. There are no surface streams but several salt ponds resulting from the closure of former bays. The beaches for the most part consist of calcareous sand and shells from the reefs. The island trends northeast/southwest and the predominant littoral drift is from east to west, although several of the bays form partially closed beach cells. Stretches of the north coast are characterised by high cliffs and there are also pronounced embayments. At Shoal Bay, there is a depositional promontory that is eroding, especially at the northern point. Along the north coast, west of Road Point, beaches have formed between limestone outcrops at Road Bay, Long Bay, Meads Bay and Barnes Bays, those at Road Bay and Meads Bay being barriers that have enclosed salt ponds. Along the western part of the south coast, beaches have formed barriers that enclosed salt ponds at Ren­dezvous, Cove, Maundays and Shoal Bays. Extensive dunes, up to 10 m high, developed behind these sandy barrier beaches. However, many of the dunes have disappeared because of alteration for tourism development, past mining for construction sand, and damage in recent hurricanes, particularly between 1995 and 1999. As a result, these barrier beaches are now being breached even during minor weather

events, as at Cove Bay. At Maunday’s Bay successive attempts to dredge and renourish the beach after hurricanes in 1995, 1998 and 1999 were largely un­­successful (>Fig. 5.5.3), and a line of tourism resort villas built on top of a dune system were  eventually protected with a sea wall (>Fig.  5.5.4). Accretionary features such as a tombolo west of Blowing Point Ferry Terminal on the south coast have all but disappeared in recent years. Monitoring of beach profiles over the period 1991–2000 showed an average erosion rate in excess of 1 m/yr (UNESCO 2003). During this period Anguilla was impacted by three hurricanes, which have damaged its extensive coral reefs; after Hurricane Luis in 1995 surveys showed large expanses of Elkhorn (Acropora palmata) reefs reduced to coral rubble (Bythell et al. 1996).

3. St. Martin

(Contributed by M.L. Schwartz) Department of Geology, Western Washington University, Bellingham, Washington, USA

St. Martin, south of Anguilla at the northern end of the  Lesser Antilles, has an area of 95 km2. The world’s smallest bi-national island, it is divided into Saint Martin (France) in the north (capital Marigot) and Sint Maarten (Netherlands Antilles) in the south (capital Philipsburg). The two divisions are known locally as the Dutch Side and the French Side. The island stands in the west-central part of the Anguilla Bank and forms part of the non-volcanic arc of ⊡⊡ Fig. 5.5.3 Beach nourishment in progress in 1996 at Maunday’s Bay, Anguilla, after Hurricane Luis (1995).

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5.5

⊡⊡ Fig. 5.5.4 Sea wall under construction in 2000 at Maunday’s Bay, Anguilla, after Hurricane Lenny (1999).

the Lesser Antilles, which curves east of the volcanic inner arc (Brouwer 1975). It has a basement of silicified tuffs with coarse pyroclastic material including Eocene carbonate fragments, intruded by Oligocene granodiorite. The hard tuffs persist as mountain areas, bordered by hilly country on less resistant igneous rocks, and there are intrusive dykes, stocks and plugs (Christman 1953). An unconformity separates the Upper Oligocene to Upper Miocene limestones, which form the lowlands. Similar to Anguilla the coastline is embayed, but has  been shortened by the formation of barrier beaches enclosing lagoons. There is a double tombolo impounding Simpson Bay Lagoon in the southwest. Pleistocene coral reef terraces occur 5–6 m above sea level, and there are nearshore reefs. Limestone headlands separate sandy beaches. The Dutch Side begins south of the western Terres Basses Peninsula on the isthmus south of the Simpson Bay Lagoon. Cupecoy has low (3–6 m) limestone cliffs facing southwest with narrow beach segments (>Fig.  5.5.5), where a formerly wide beach was depleted by Hurricane Luis in 1995 and Hurricane Jose in 1999. Mullet Bay to the east has low karstic limestone shores. At Maho an emerged limestone terrace is fronted by cliffs with caves, declining eastward to a curving sandy beach with segments of grassy backshore terrace. South-facing Simpson Bay is backed by a 1.5 km long curving sandy barrier, with a bridge across an inlet to the Simpson Bay Lagoon. Promontories and bays continue eastward to Great Bay, where Philipsburg is built on a barrier enclosing a lagoon, the Great Salt Pond, where salt was extracted formerly. In 1995 and 1999 Hurricanes Luis and Lenny damaged beaches and coral

reefs, and the Great Bay Beach was subsequently replenished with dredged sand. Pointe Blanche is a wide limestone peninsula on the southeast coast. The east coast, more exposed to the Northeast Trade winds and strong waves, has limestone cliffs and sandy beaches. A sand spit almost encloses a small lagoon at Oyster Pond. North of Oyster Pond, the boundary between the Dutch and French sides crosses the coast. At Orleans a sandy barrier encloses another lagoon, and on its seaward side is Galion Beach, often strewn with seagrass. Orient Bay, to the north, has a broad sandy beach. Offshore are extensive coral reefs in a marine reserve. On the northwest coast, Grand Case has another barrier beach enclosing a lagoon, and in Friars Bay to the south there are beach gravels (shingle) beneath the sand. South of the town of Marigot the barrier beach in Baie Nettle, on the northern side of Simpson Bay Lagoon, has been extensively armoured against erosion. To the west is the Terres Basses Peninsula, with Baie Longue on its southern coast. This is a wide sandy beach backed by a small lagoon and fronted by a shallow bay where wave energy is generally low.

4. Guadeloupe

( Contributed by Dr. Pascal Saffache and Garry Ramedine) Université des Antilles et de la Guyane

The Guadeloupe Archipelago has 680 km of coastline and is composed of eight islands: Basse-Terre, Grande-Terre,

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⊡⊡ Fig. 5.5.5 Limestone cliffs at Cupecoy Beach, St Martin.

Terre-de-Haut, Terre-de-Bas, Marie-Galante, La Desirade, Saint-Martin and Saint-Barts. While some of the islands are exclusively volcanic (Basse-Terre and the Saintes ­archipelago) in others the volcanic bedrock is covered with calcareous sediment (La Desirade, Grande-Terre, Saint-Martin, Saint-Barths, Marie-Galante). Emergence, due to tectonic uplift, results in cliffs over 30 m high on eastern coasts with sandy coves containing emerged beach rock. Beach rock also outcrops extensively in many other islands of the Lesser Antilles indicating the former position of beaches as coastlines have retreated inland. The Gates of Hell in the north-east is a limestone calanque bordered by flat ledges and backed by a sandy beach washed by Atlantic waves. The western coasts have low cliffs cut in pyroclastic tuff backing beaches and swamps. A tidal channel, Rivière Salée, runs across the northern end of Grande-Terre from the Atlantic to the Caribbean. There are alternating black volcanic sediment and white coralline sediment beaches. Bays and lagoons, e.g., the Grand-Cul-de-Sac Marin and Petit-Cul-de-Sac Marin, the Saint-Martin and Saint-Barths lagoons, are occupied by mangrove swamps. The inner wetlands area is composed of red mangrove (Rhizophora mangle), the middle section by black mangroves (Avicennia germinans), and the outer portion by red mangroves. Sandy beaches in the Guadeloupe archipelago occupy 19% of the coastline, mainly in bays or behind reefs. Most of the sandy beaches are along the south coast of Grand-Terre (between Gosier and Pointe des Chateaux),

in the north-west of Basse-Terre (between Bouillante and Deshaie) and on the south-west coast of Marie-Galante and Desirade. Gravel beaches with black pebbles of volcanic origin are present in bays along the southern coast of Basse-Terre, while white pebbles, formed from calcareous rock, are found only at Anse Pistolette in the north of Grande-Terre. In nearshore water up to 30 m deep, there are fring­ ing  reefs, barrier reefs (such as the Grand-Cul-de-Sac Marin) and volcanic rock outcrops covered with coral. In Guadeloupe, the fringing reef is limited to 90 km in length. The reefs are located mostly along the southern coast of Grande-Terre, in the Petit-Cul-de-Sac Marin, to the south of Marie-Galante, and on the Atlantic coast between Pointe des Chateaux and Port-Louis as well as surrounding Petite Terre and La Desirade. The Grand-Cul-De-Sac Marin barrier reef is about 30 km long and up to 10 km wide, enclosing a 78 km2 lagoon with a maximum depth of 20 m. Behind the solid coral reef is a zone of seagrass (Thalassia testudinum) on sandy shoals, then mangroves along the coast. All these features have been modified by human activities.

5.  Dominica Dominica, with an area of 751 km2 is a rugged, steep volcanic island with a mountainous interior and many rivers. The island began to form during the Miocene period and the volcanic rocks are mainly andesites with subordinate

Lesser Antilles

dacites and basalts. Active volcanism is evident in the Sulphur Springs region and the area of the Boiling Lake and the Valley of Desolation. Most of the flat or gently sloping land exists near the coasts and this is where most of the urban and agricultural developments have taken place. Most of the 153 km of coastline is cliffed with beaches along 24 km. With a very narrow coastal shelf, there are few true coral reefs in Dominica. Along the west coast, there are a number of narrow beaches made up of a mixture of volcanic sand and stones. Many of the beaches are located near the place where rivers discharge into the sea. They often exhibit a seasonal change in composition with sand during the lower wave energy months from May to September, and stones during the high wave energy months of November through to March. At the SW tip of the island the small islet of Scotts Head is joined to the mainland by a narrow gravelly tombolo (>Fig. 5.5.6). Prior to Hurricane David in 1979, this tombolo was composed of sand and pebbles; it was wide enough to have a football pitch and there were many large trees. Now the tombolo is low and narrow and often breached during storms and hurricanes. Along much of this west coast, the coastal highway runs near to the sea and had to be protected with extensive seawalls, usually resulting in the loss of any narrow beaches that may have existed in the past. This has created problems for land crabs that annually migrate to the sea to deposit their eggs. Monitoring of beach profiles along the coasts of Dominica has shown erosion rates of 1 m/yr over the period 1987–2000 (UNESCO 2003). The capital city of Roseau has been protected with a massive sea wall combined with boulder protection. ⊡⊡ Fig. 5.5.6 Tombolo joining Scotts Head Islet to the main island of Dominica.

5.5

At the northern end of the west coast there is a ­ ronounced embayment, Prince Rupert Bay, where the p second largest town, Portsmouth is located. A once wide, continuous sandy beach now exists only in patches. Behind the beach there is the extensive Indian River wetland area. Further to the north, in the Cabrits Peninsula, once sandy beaches at Belle Hall and Toucarie, which in the past were favoured for recreation by Dominicans, are now eroded and reduced to stones and gravel following several hurricanes between 1989 and 1999. The north coast from Hampstead to Pagua has several sandy embayments and in places such as Pointe Baptiste there are very scenic cliffs and pockets of brown sand beaches. Much of the east coast is cliffed with long stretches of black sand beaches at Londonderry, Castle Bruce and Bout Sable, to mention a few. Most of this coast experiences high wave energy for much of the year. At Marigot a man-made fishing harbour has been created.

6.  Barbados Barbados has an area of 430 km2 and a coastline 92 km long. It consists mainly of Tertiary limestone, marl and flysch deposits, with a Pleistocene coral limestone cap up to 100 m thick. The island has a regular coastline with no deep indentations or offshore islands. The coast consists mainly of long stretches of sandy beach, except in the north and southeast where cliffs rise straight from the sea. On the east coast the Pleistocene coral limestone that covers the rest of the island has been eroded away ­exposing

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the less resistant underlying strata, mainly marls and clays, in a region known as the Scotland District. A 10 km long, wide quartz sand beach backed by low sand dunes occupies most of this coast, which is fully exposed to the Atlantic waves and has minimal protective coral reefs. Blocks of limestone from the eroded cap are visible in the offshore area. This beach ends in cliffs at both ends. At the southern end, the coast becomes more indented with bays separated by cliffs. At Bathsheba there are several stacks (>Fig. 5.5.7) and between Bath and Consett Bay, remains of the old railway line can be seen near low water mark, indicating the erosion taking place along this coast. The SE coast from the Ragged Point lighthouse to South Point is characterised by coral limestone cliffs with several significant bayhead beaches, of which Foul Bay and Long Bay are the longest. The beaches are made up of sand and coral, shell, and algal fragments. The Cobblers Reef, a barrier reef, lies between 1 and 3 km offshore, and encloses a partially sheltered lagoon. Along the south coast from South Point to Bridgetown the coastline is low and rocky with intermittent coralline sand beaches. Offshore there are rubble banks and scattered coral patch reefs. At Oistins a long jetty originally planned for use by the Coastguard trapped a wide sandy beach, known locally as “Miami Beach”. In several places, groynes have been built to trap the westward moving sediment and at Accra Beach there is a submerged offshore breakwater. Near St. Lawrence Gap, the one remaining

significant wetland in Barbados, Graeme Hall Swamp, is connected to the sea via a sluice gate. The coastline is densely developed with tourism properties and associated infrastructure. At the western end of the south coast, a former accretionary feature has been stabilised with a breakwater at Needhams Point. The main town, Bridgetown and the deep water harbour are located at the SW end of the island, and significant accretion has taken place north of the deep water harbour at Pelican Beach. The west coast consists of a low sand terrace varying in width from a few hundred metres to 1 km. There is a nearly continuous coralline sand beach, consisting of some 19 beach cells that are anchored by fringing reefs at the sandy headlands separating the cells (Bird et al. 1979). These fringing reefs are poorly developed and have been stressed for decades by runoff and pollution from densely developed west coast tourism. Some of the beach cells are well developed with wide beaches while others have only narrow strips of sand, which disappear at high tide. Many beachfront property owners have resorted to the use of backshore protection, with varying degrees of success. Despite these pressures, the west and south coasts are important sea turtle nesting sites, particularly for Hawksbill turtles (Eretmochelys imbricata). The beaches of Barbados have been under pressure from erosion (>Fig.  5.5.8) and intensive use since the 1980s and as a result, a Coastal Zone Management Unit has been established to monitor and manage these issues. ⊡⊡ Fig. 5.5.7 Stacks at Bathsheba, Barbados.

Lesser Antilles

5.5

⊡⊡ Fig. 5.5.8 Eroded west coast beach at Heywoods, Barbados.

7.  St. Vincent and the Grenadines St. Vincent is one of the youngest of the islands of the inner volcanic arc. The northern end of the island is formed by the active volcano, Soufrière, which is a recent volcano built up within the last half million years, and was last active in 1979. This volcano, like many of the others in the Lesser Antilles, is monitored by the Seismic Research Unit of the University of the West Indies in Trinidad and Tobago. The island has a central chain of mountains, which are the eroded remnants of a series of volcanoes, and a coastal plain of varying width. Cliffs have been cut in the volcanic formations, and locally there are outlying stacks (>Fig. 5.5.9). The chain of Grenadine islands to the south of mainland St. Vincent has a more complex geological history and has rock types of both sedimentary and volcanic origin (Caribbean Conservation Association 1991). Shore shelf around the St. Vincent mainland is very narrow except at the southeast side. Most of St. Vincent’s beaches are made up of black sand, stones and boulders; the exception is on the south coast where there are coral reefs and white sand beaches. The east coast consists of long stretches of beaches separated by headlands of more resistant rock. The beach and nearshore sediment along the windward (eastern) coast of St. Vincent are dominated by volcanic ash fragments and

heavy minerals (Adams 1968). There are several ­significant rivers, bringing sediment to the coast, e.g. the Rabacca River north of Georgetown, where the river bed is still mined for construction material (>Fig. 5.5.10). Settlements on the coastal plain, near the mouths of these rivers, are vulnerable to flooding during storms and hurricanes (>Fig.  5.5.11). There used to be extensive black sand dunes at Brighton and Diamond Bays at the south end of the east coast. The Brighton dunes, with maximum elevations of 60 m, were formed when sand accumulated north of a tombolo opposite Milligan Cay, and were blown landward by the Trade Winds. Dunes 6 m high at Diamond Bay were removed by quarrying in the 1970s and 1980s. The south coast consists of several beach segments with a mixture of coralline (white) and terrestrial (brown) sand. This coast has been substantially modified with several sea defence structures. The west coast consists of a series of indented bays with volcanic sediment (sand, gravel and boulders). Mining from the beaches and river mouths, although usually small scale, is still a significant issue in St. Vincent. The Grenadine Islands, to the south of mainland St. Vincent, are located on a platform, 40 m deep and are an important yachting destination. There are extensive reef systems, which have been impacted by anthropogenic impacts, hurricanes and coral diseases. Several of the

307

308

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Lesser Antilles

⊡⊡ Fig. 5.5.9 South Stack, a volcanic neck with an abrasion notch at present sea level. (Courtesy J.R.V. Prescott.)

⊡⊡ Fig. 5.5.10 River mouth on the east coast of St. Vincent. The river is delivering sand and gravel to beaches, but the yield has been diminished by extraction. (Courtesy J.R.V. Prescott.)

islands have significant wetlands, such as The Lagoon on the SW coast of Mustique. The islands are rocky and vary in size and elevation. The coastlines are indented, with white sand beaches separated by cliffed headlands. There are few significant sand dune features. Bequia, one of the largest islands, has a large embayment on the west coast, Port Elizabeth. This is an important yacht anchorage, and the bay is fringed with a narrow white sand beach.

8.  Trinidad and Tobago Trinidad and Tobago with a combined area of 5128 km2 is situated on the continental shelf off the north coast of Venezuela. Trinidad has three mountain ranges orientated east–west and numerous rivers drain the upland areas ending in coastal swamps and lagoons. Tobago is rugged with an internal mountain range and a narrow coastal plain.

Lesser Antilles

5.5

⊡⊡ Fig. 5.5.11 Villages behind this black sand beach on the east coast of St. Vincent are vulnerable to flooding by seawater and land runoff.

⊡⊡ Fig. 5.5.12 Caroni Swamp, Trinidad, home of the Scarlet Ibis.

The east coast of Trinidad is dominated by three long bays: Matura, Cocos and Mayaro Bays. The beaches at Matura and Mayaro Bays are composed of quartz sand eroded from cliffs cut in Miocene and Pliocene formations and moved by rivers to the coast. At Cacos

Bay, a 20 km barrier beach has formed, impound­‑ ing three rivers to make the large freshwater Nariva Swamp, 70 km2 in extent. The barrier beach is broken only at the centre, where the Nariva River reaches the sea.

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Lesser Antilles

Longshore movement of sand along the south coast of Trinidad ends in the large Icacos spit at the SW tip of the island. The major source of beach sediment is from erosion of cliffs composed of sands, clays and weakly cemented sandstones of Miocene, Pliocene, and Pleistocene age on the western part of the south coast. Mangroves are extensive. On the SW coast of Trinidad cliffs in Miocene sands and clays are being eroded. This sector of coastline comes to an end at Los Gallos, where there are stacks of slightly more resistant Pleistocene sandstone. The near-estuarine conditions of the Gulf of Paria restrict coral development and the coast is dominated by mangroves fringing the extensive Caroni Swamp (>Fig. 5.5.12). Along the north coast, there are several small embayments and considerable coral development. Tobago has a more indented coastline, with several sandy embayments separated by lengths of cliff coastline. The Buccoo Reef area on the SW coast has extensive seagrass and associated mangrove stands.

References Adams RD (1968) Distribution of littoral sediment on the Windward Coast St. Vincent, West Indies. Transactions of the 5th Caribbean Geological Conference, pp 55–59

Bird JB, Richards A, Wong PP (1979) Coastal subsystems of western Barbados, West Indies. Geogr Ann 61A:221–236 Brouwer GC (1975) Netherlands Antilles. In: Fairbridge RW (ed) Encyclopedia of world regional geology, part 1: western hemisphere. Dowden Hutchinson and Ross, Stroudsburg, pp 362–365 Bythell JC, Cambers G, Hendry MD (1996) Impact of Hurricane Luis  on  the coastal and marine resources of Anguilla. Summary report prepared for the UK Dependent Territories Regional Secretariat, 13 p Caribbean Conservation Association (1991) St. Vincent and the Grenadines country environmental profile, 222 p Christman RA (1953) Geology of St. Bartholomew, St. Martin and Anguilla, Lesser Antilles. Geol Soc Am Bull 64:65–96 Deane C (1985) Lesser Antilles: Virgin Islands to Trinidad. In: Bird ECF and Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 193–199 Hendry MD (1996) The geological legacy of small islands at the Caribbean–Atlantic boundary. In: Maul GA (ed) Small islands marine science and sustainable development. American Geophysical Union, Coastal and Estuarine Studies, vol 51, pp 205–224 IPCC (Intergovernmental Panel on Climate Change) (2007) Climate change 2007: the physical science basis. Summary for policy makers UNESCO (2003) Wise practices for coping with beach erosion in the  islands of the eastern Caribbean. 10 booklets, Environment and Development in Coastal Regions and Small Islands. UNESCO, Paris Wells SM (ed) (1988) Coral reefs of the world. Atlantic and Eastern Pacific. UNEP Regional Seas. IUCN Monitoring Centre, 373 p

6.0 Atlantic Ocean Islands – Editorial Introduction

The Atlantic Ocean was initiated about 140 million years ago in the Late Jurassic when the disruption of the ancient megacontinent of Pangaea was followed by westward movement of the American Plate and eastward movement of the Eurasian and African Plates. It is still widening by up to 10 cm/year. The divergence was marked by the formation of the mid-Atlantic Ridge by the extrusion of basalt, resulting in the formation of a series of volcanic  islands extending from Jan Mayen Island in the north  to  Bouvet Island in the south: the series includes Iceland, the Azores, St Peter and Paul Rocks, Ascension, St Helena, Tristan da Cunha and Gough Island. In addition, Greenland is a largely ice-covered island of continental Pre-Cambrian and younger rocks projecting into the NW Atlantic and the Falkland Islands are also of continental rocks, with a Pre-Cambrian basement overlain by Mesozoic formations on the South American shelf. The >Bahamas and >Bermuda are islands of limestone and associated calcareous deposits rising from the American shelf. East of the mid-Atlantic ridge the Faeroe Islands, St Kilda, Madeira, the Selvagens archipelago, the Canary Islands, the Cape Verde Islands and the islands in the Gulf of Guinea all surmount the Eurafrican shelf (>Fig. 6.0.1). The Atlantic islands are generally high islands, mostly of volcanic origin, with cliffed, rocky and boulder-strewn coasts (Hansom 2005). Coral reefs are poorly developed (except in the Bahamas and Bermuda), and the low coralline islands and  cays found in the Indian and Pacific Oceans are not present. There are active volcanoes in Iceland, in the Azores and on Tristan da Cunha. Volcanic structures have influenced coastal evolution, with rapidly receding cliffs in ­scoria and ash while lava flows, dykes and necks form persistent bold cliffs, rocky headlands and stacks. Steep coasts and plunging cliffs are common. Encrustations of algae are found on rocky shores, notably in Ascension, where coralline algae coat bare rock and boulders. Emerged beaches and shore terraces are found at various levels, indicating episodes of tectonic uplift as well as sea level changes. Many of the islands have beaches of dark volcanic sand, as well as gravel and boulders derived from the

v­ olcanic rocks. Locally there is pale calcareous sand, mainly from shelly organisms, as on Ascension. The Bahamas and Bermuda have coral, limestones and biogenic carbonate sands, and exposures of beach rock. Sable Island, 300 km SE of Nova Scotia, consists of ridges and shoals of quartzose sand derived from glacial drift deposits. Quartzose sands in the eastern Canary Islands may have come from the Sahara. The Atlantic Ocean encompasses climatic zones from Arctic in the north through the mid-latitude westerly wind belt to a calm high pressure zone around 30° north, then the NE trade winds, the equatorial calms, the SE trade winds, another calm zone around 30° south, and the southern mid-latitude westerlies extending into the Antarctic. There is ice on Jan Mayen Island and in Iceland, and on  Bouvet Island in the far south, and frost shattering occurs on rock outcrops on these islands and in the South Shetlands. There are sectors of peat cliff in Iceland and the Falkland Islands. Sea ice forms on the shores of islands in high latitudes. Within the tropics the islands are generally wet and well vegetated, but in the eastern Atlantic the Cape Verde Islands, like the subtropical Canary Islands to the north, are relatively dry. Mangroves are found on the Bahamas and at their northern limit in Bermuda, but there are few niches for mangroves or salt marshes on the rocky volcanic islands. Ocean swell, initiated in stormy areas in the South­ ern  Ocean moves across the Atlantic from the south, arrives as a SW swell on the west coasts of Europe and  Africa, and as a SE swell on the east coasts of the Americas. Storms in the North Atlantic generate waves that arrive as  a NW swell on the west coast of Europe, occasionally  penetrating to NW Africa, and as a NE swell  on the  east  coast of North America, occasionally penetrating  to  Brazil. There are contrasts between the bold coasts of  islands exposed to strong wave action over  long fetches  and those that are more sheltered, ­sometimes intricate and  low-lying. Mean spring tide ranges are ­relatively large  in >Greenland and >Iceland but generally small (<2 m) on >Other Atlantic Ocean Islands.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

6.0

Atlantic Ocean Islands – Editorial Introduction

⊡⊡ Fig. 6.0.1 Atlantic Ocean Islands. (Courtesy Geostudies.) Iceland Greenland

Faeroe Is. St. Kilda Rockall

British Is.

EUROPE NORTH AMERICA

RI

DG

E

Sable I.

Azores

Bermuda

IC

Madeira Selvagens

NT

Bahamas

- ATLA

Canary Is.

MID

Cape Verde Is.

AFRICA

Fernando Po Principe Sao Tome Annobon

St. Peter and St. Paul Rocks Atol das Rocas

Ascension

SOUTH

AMERICA

RIDGE

Fernando de Noronha

St. Helena

C

Martin Vaz

TLA

NTI

Trindade

M I D- A

312

Tristan da Cunha

N

Gough I.

4000m. Isobath Falkland Is.

SCOTIA

ATLANTIC OCEAN ISLANDS 0

1000

South Shetland Is.

South Georgia

SEA

Bouvet I.

South Sandwich Is. South Orkney Is.

2000 km

Reference Hansom JD (2005) Atlantic Ocean islands, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 88–95

ANTARCTICA

6.1 Bahamas

Alan K. Craig

1.  Introduction

there are larger fluctuations in sea level related to variations in atmospheric pressure and strong winds. The islands are situated upon the Bahama Banks (>Fig. 6.1.1), large sedimentary shoals typically no deeper than 4 m at low tide, but plunging at their margins to oceanic depths. The four principal Banks (Little Bahama, Great Bahama, Cay Sal, and Acklins) constitute a group that extends to the southeast where, although not politically part of the Bahamas, the Turks and Caicos islands and three more shallow banks (Mouchoir, Silver, and Navidad) to the southeast are geologically similar. They consist of a thick sequence of Cretaceous, Tertiary and Quaternary sediment including calcareous sandstones, limestones, coralline and other biogenic material, accumulated since

The Commonwealth of the Bahamas, east of Florida and north of Cuba, consists of 29 substantial islands and over a thousand cays, islets and rocks with a total area of over 13,000 sq. km and a 3,542 km coastline. They straddle the Tropic of Cancer, and have a subtropical climate with warm summers and mild winters; Nassau has a mean monthly temperature of 21.7°C in January, rising to 27.2°C in July, and an average annual rainfall of 1,179 mm. The prevailing winds are the northeast trade winds, and hurricanes may occur between July and November, passing westward into the Caribbean. Atlantic ocean swell and storm waves arrive from N to NE. The tide range is small, about a metre, but

⊡⊡ Fig. 6.1.1 The Bahamas, Turks and Caicos Islands. 80°

73°

27°

27°

BAHAMA GRAND BAHAMA ISLAND

GREAT ABACO ISLAND

I S L A N D S

ELEUTHERA ISLAND 22°

NEW PROVIDENCE ISLAND TONGUE ANDROS ISLAND

69°

CAICOS ISLAND

EXUMA

OF THE

SOUND

CAT ISLAND

TURKS ISLAND

OCEAN

MOUCHOIR BANK SAN SALVADOR

21° SILVER BANK

CAY SAL LONG ISLAND

GREAT BAHAMA BANK

NAVIDAD BANK 20° MAYAGUANA ISLAND

72°

ACKLINS ISLAND

GREAT INAGUA ISLAND 21° 80°

0

5

73°

10 kms

21°

OIB. L BEFORE JUN. 31

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

69°

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Bahamas

Cretaceous times and resting upon a Jurassic basement which has subsided to a depth of 6,000–7,000 m (Newell and Rigby 1957). They are not coral reefs, but sedimentary structures (notably dune calcarenites that formed when sea level was relatively low) bearing extensive corals in patterns that show variations related to wave energy. They are areas of oolite formation, together with grapestones (aggregates), biogenic sands and silts, and oozes. The sea floor has planed-off coral limestone and dune calcarenite, and ridges and ripples of varying dimensions and mobility, with extensive seagrass growth. The Bahama Banks are trenched by some of the world’s deepest canyons, with steep walls and floors that descend to abyssal plains that may be more than 6,000 m deep (Andrews et al. 1970). These include the wide Tongue of the Ocean, beside Andros Island, and Great Abaco Canyon. There are also circular blue holes, submerged dolines up to 80 m wide, which in some cases emit fresh water from submarine springs. The islands lie near the northeastern margins of the Bahama Banks, the largest being Andros (168 km by 64 km), the others generally long and narrow, aligned at right angles to the dominant northeast waves and often recurved at the ends. They have relatively subdued, riverless landscapes of stabilised calcarenite dunes that rise to a maximum height of 67 m at Mt. Alverina on Cat Island (>Fig. 6.1.1). In general, the northeastern coasts rise relatively steep to a dune calcarenite ridge, followed by a gentler southwestern slope down to lagoons and mangrove swamps. Pleisto­ cene dune calcarenites, generally well indurated (Bahamas Limestone) and having conspicuous karstic features, are common in the interior and underlie dunes up to 20 m high (Hearty and Kindler 1997). The coastlines are dominated by low cliffs and shore platforms cut in Pleistocene aeolian calcarenites that formed as successive, superimposed dune formations during interglacial phases, when the sea was at or close to its present level. These dunes were subsequently stabilised and lithified during glacial phases when the sea was relatively low and the Bahama Banks emerged, and soils (now seen as interbedded palaeosols in  cliff sections) formed, incorporating airborne dust. The dune calcarenites were then trimmed back as cliffs during the Holocene marine transgression and ensuing stillstand. The cliffs and shore outcrops show strong karstic weathering, with jagged pinnacles and rough pitted rock surfaces, as well as notches and visors at the cliff base. Locally there are vertical soil pipes with indurated rims of travertine, which protrude, or have been excavated as banana holes. The islands have windward (NE) and leeward (SW) coastlines, the strongest wave action being from the

ocean to the northeast. Coastlines facing the open Atlantic are subject to large swell and occasional storm waves that sweep in over the narrow coral-covered shelf slope and break on the barrier reef rim. Between the barrier reef rim and the beach there is usually a comparatively broad, shallow lagoon with varying amounts of biogenic sand being transported toward the coastline (Rankey 2002). Microa­tolls occur here, oval or circular structures with diameters of up to 30 m, rimmed by vigorously growing corals and algae. There are also algal mats (stromatolites), thin sheets of blue-green algae, diatoms and other organisms, which trap and bind sediment and induce carbonate precipitation. The beaches are of calcareous sand and gravel, including coralline, algal and oolitic sediment, with occasional outcrops of beach rock. On Cat Island there are sub-parallel Holo­ cene sandy beach ridges up to 5 m high, occurring in truncated sets on successively higher terraces that are occasionally capped with stabilised dunes. Leeward coastlines have wide shelves (e.g., Great Abaco, Andros, Eleuthera, Aklins and Caicos islands) on which highly mobile, fine-grained calcareous sediment occur as shoals. The north coast of Acklins Island has a chain of low narrow barrier islands of caronate sand banked up along the outer edge of shoals. Oolitic sands and grapestone are forming. Large sand shoals are often dissected by tidal channels. Surf reaching leeward beaches is seldom more than a metre high, so that both depositional and erosional features tend to be subdued. Fossil beaches and hypersaline lagoons with abandoned tidal channels are common. Sets of parallel and truncated beach ridges occupy elevations equivalent to the more extensive series found on the windward shore. Emerged Pleistocene marine terraces have been described from the island of Eleuthera (Kindler and Hearty 2000). Some have explained the beach ridges as the outcome of eustatic fluctuations of sea level, each ridge formed during a phase of emergence and trimmed back when the sea returned to its former level (Lind 1968). Others consider that the beach ridges are due to ll major storms alternating with calmer phases (Perkins and Enos 1968; Shinn et al. 1969), particularly on Andros Island. The effects of severe storms on the Great Bahama Bank and shore of the Bahamas in general appear to depend in large measure on the local conditions associated with each event. Some hurricanes have crossed the Bahamas leaving little trace, but most have had major effects on coastlines, eroding beaches and cutting back cliffs. Deep, sub-parallel channels cut through low-lying islands are known as bogues (as on the east end of Grand

Bahamas

Bahama and the centre of Andros), and are believed to be storm-related. Erosion has been a problem on several islands in the Bahamas. In the Turks and Caicos Islands, the former Back Street of Cockburn Town is now Front Street because of coastline recession. Throughout the Bahamas, vegetation has been disturbed by human activities. Nearshore species are halophytic and the Australian pine (Casuarina equisetifolia) is ubiquitous. The interior of Andros Island supports native pine stands, but a complex, dense coppice is typical of other islands. Red mangroves (Rhizophora mangle) are often found along leeward shores.

References Andrews JE, Shepard RP, Hurley RJ (1970) Great Bahama Canyon. Geol Soc Am Bull 81:1061–1108

6.1

Hearty PJ, Kindler P (1997) The stratigraphy and surficial geology of New  Providence and surrounding islands, Bahamas. J Coastal Res 13:798–812 Kindler P, Hearty PJ (2000) Elevated marine terraces from Eleuthera (Bahamas). Glob Planet Change 24(1):41–58 Lind AO (1968) Coastal Landforms of Cat Island, Bahamas: A Study of Holocene Accretionary Topography and Sea-Level Change. University of Wisconsin Newell ND, Rigby JK (1957) Geological Studies on the Great Bahama Bank. Society of Economic Paleontologists and Mineralogists, Special Publication 5:15–72 Perkins RD, Enos P (1968) Hurricane Betsy in the Florida-Bahama Areageologic effects and comparison with Hurricane Donna. J Geol 76:710–717 Rankey EC (2002) Spatial patterns of sediment accumulation on a Holocene carbonate tidal flat, NW Andros Island, Bahamas. J Sediment Res A 72:591–601 Shinn EA, Lloyd RM, Ginsburg RN (1969) Anatomy of a modern carbonate tidal-flat, Andros Island, Bahamas. J Sed Petrol 39:1202–1228

315

6.2 Bermuda

Eric Bird

Introduction Bermuda consists of a group of about 150 small limestone islands in the west Atlantic (32°20' N, 64° 45' W), just over 1,000 km east of Cape Hatteras. They are situated near the south-eastern margin of the submerged Bermuda Platform on part of a narrow submerged reef that encircles a lagoon up to 16 m deep. They have a total area of 54 km2 and are in the form of an italic J, with a 103 km coastline, the longest segment of which faces SSE. The Main Island is 36 km long and generally 2–3 km wide, running WSW-ENE, with a curving southwestern extension round to the Ireland islands, sheltering Great Sound and Little Sound, a recurved inner ridge on which the little city of Hamilton stands, and narrow strips enclosing three large lagoons, Harrington Sound, Castle Harbour and St George’s Har­ bour, at the northeastern end. Fort St. Catherine stands on the northernmost point. There are many sites of scientific interest around the island coasts, but visitors should be warned that some are on private land and inaccessible. The climate is subtropical, under the influence of the warm Gulf Stream, with an average annual rainfall of 1,450 mm. The prevailing winds are westerly, sometimes stormy in winter. In addition to waves generated by westerly winds there is a southerly Atlantic Ocean swell which becomes southeasterly as it breaks on the coast of Main Island. The tide range is small, about a metre. The islands are hilly, rising to summits 60–70 m above sea level, and consist of Pleistocene dune calcarenites (aeolianites) and associated beach calcarenites (Sayles 1931) on the rim of the Bermuda Platform, a submerged seamount truncating a volcano that was last active in the Oligocene. Borings have shown that the base of the calcarenites rests upon an eroded foundation of volcanic lava undulating between 20 and 170 m below sea level. The rock sequence indicates that there has been tectonic subsidence, but the Bermuda Platform is thought to have become stable in Pleistocene times. The dune calcarenites formed from the deposition of successive, superimposed dunes deposited during interglacial phases behind and over Pleistocene beaches that existed when the sea stood at or near its present level (Vacher 1973). These dunes were subsequently stabilised

and lithified during glacial phases, when extensive areas of the Bermuda Platform emerged. Terra rossa soils then formed on the dunes, incorporating airborne dust presumably from North America, and these are now seen as interbedded palaeosols, sometimes associated with indurated calcrete horizons. The rock formations are highly permeable, and there are no rivers. The Pleistocene dune calcarenites were trimmed back as low cliffs and shore platforms during the Holocene marine transgression and the ensuing stillstand. The cliffs and shore outcrops show strong karstic weathering, with jagged pinnacles and rough pitted rock surfaces, as well as notches and visors at the cliff base. Locally there are vertical soil pipes with indurated rims of travertine. The cliffs are fronted by erosional shore platforms that locally carry corals and algae, and associated calcareous sandy beaches. There is evidence of Pleistocene marine terraces up to 22 m above present sea level, but no indication that sea level rose significantly above present datum in the Holocene (Hearty and Kindler 1995). Beaches are calcareous, consisting of calcarenite sand and gravel with some coralline debris. They have a pink colouration, attributed to the presence of red foraminifera (Homotrema rubrum). Sand and gravel from the bordering reefs is still being delivered to beaches, but on the northern (lagoon) shore the supply has diminished and the coast is cliffed with only a few pocket beaches. The cliffed coastline has numerous small bays and coves formed by marine submergence of low-lying parts of the dune calcarenite landscape. Sinky Bay is an almost circular cove. There are caves, natural arches and stacks. Mangrove Lake is the largest of five shallow brackish (anchialine) ponds with limited tidal flushing through fissures and small caves. Coral reefs are extensive off the island coasts, those on the northwest side being the northernmost reefs in the Atlantic Ocean. They have been damaged by the dredging of navigation channels and the grounding of ships. Dundonald Channel is a straight navigation channel cut across a broad shoal to access Great Sound and Hamilton Harbour. Microatolls are found, oval or circular structures with diameters of up to 30 m, rimmed by vigorously growing corals and algae, and there are also serpulid ridges on

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Bermuda

calcarenite shore platforms and cup reefs built by vermetid gastropods and algae. Sinkholes formed by solution during low sea level phases persist in coral reefs and platforms of calcarenite. Mangroves grow at their northernmost limit (32° N) in  Bermuda, and include the red mangrove (Rhizophora mangle), black mangrove (Avicennia germinans) and button mangrove (Conocarpus erectus). They have responded to a sea level rise of 28 cm over the past century by aggrading a peaty substrate, but there has been recession of their seaward margins (Ellison 1993). 

References Ellison JC (1993) Mangrove retreat with rising sea level, Bermuda. Estuar Coast Shelf Sci 37:75–87 Hearty PJ, Kindler P (1995) Sea level highstand chronology from stable carbonate platforms (Bermuda and the Bahamas). J Coastal Res 11:675–689 Sayles RW (1931) Bermuda during the Ice Age. Proc Am Acad Arts Sci 66:381–468 Vacher L (1973) Coastal dunes of younger Bermuda. In: Coates DR (ed) Coastal geomorphology. State University of New York, Binghampton, USA, pp 355–391

6.3 Greenland

Introduction The coast of Greenland is related to the variety of rock types, and the activity of glacier ice during the Pleistocene and Holocene. Much of the western and southeastern coast consists of bedrock, often fringed by sea ice. The coast facing the Polar Sea is primarily built up of PreCambrian sedimentary rock, whereas the northeast coast is dominated by Palaeozoic and Mesozoic formations. Younger geological strata, such as erodible Cretaceous sandstone and plateau basalts of Tertiary age are found both on the east and the west coast, but only between latitudes 69° and 72° N (Escher and Watt 1976). Most of the rocks outcropping along the Greenland coasts are very resistant to marine erosion, and the coastal landforms bear the strong imprint of glaciation. The ice cap has been much more extensive, covering nearly all of Greenland until about 12,000 years ago. The effects of glaciation are most obvious in the numerous large fiords, where the ice cap, through selective erosion along preexisting (Tertiary) valleys, created deep troughs that were invaded by the rising sea as the ice cap partially melted. More broadly, the spreading ice smoothed rock ridges and dissected coastal outcrops into scattered hills, which with postglacial marine submergence have become skerries on archipelagic coasts. Greenland covers an area of about 2.2 million km2, of which only 342,700 km2 are ice-free. The precise length of the coastline is unknown, but is estimated to be at least 40,000 km. The range of variation in coastal geomorphology is nevertheless relatively limited. About 90% of the coast is rocky, with more than 100 long, deep and much branched fiords. In Scoresby Sound on the east coast there are 11 fiords with lengths as great as 325 km and depths of up to 1,450 m. Present-day marine erosion and accumulation are not conspicuous, but there are gravelly pocket beaches (pebbles to boulders), and on basalt and limestone outcrops marine processes have shaped coastal landforms. On the west coast island of Disko, for example, there are high

cliffs along the south and west (> Fig. 6.3.1) and long shingle barriers with lagoons and marshes on the east (Nielsen 1969). On parts of the Greenland coast the ice cap still reaches the sea. In Melville Bay, northwest Greenland, an ice front forms cliffs up to 40 m high almost 500 km and farther north, at Kane Basin, ice cliffs extend for about 100 km. Many valley glaciers discharge into the sea, some of them producing icebergs. When annual mean wave energy levels are calculated for the outer coast on the basis of wind statistics and fetch, far higher values are found than those that actually occur. This is partly due to the freezing of the sea, and partly to the arrival of ice floes from the Polar Sea. These pass down the east coast, round the southernmost point, then northward, carried by the Irminger Current along the west coast to about 65° N latitude. There are also many icebergs from the calving glaciers. All these types of ice hamper the transmission of waves to the coast. The sheltering effect of the ice favours various accretion forms shaped under low wave energy conditions. Where there is occasionally no ice, the full energy of the ocean waves reaches the shore. Nanok (ice bear) hunters on the northeast coast have reported that a single gale completely changed the configuration of the coast, which had been stable as long as could be remembered. In this case the Storis, the constantly south-flowing ice-laden current from the Polar Sea, had been blocked for some time by heavy ice farther north, and thus could not exert its usual subduing effect on the waves. Although this situation is rare, such events must also be considered when interpreting coastal geomorphology. Mean spring tide range varies from about 3 m in the southwest at Nanortalik (60° N), 4.5 m at Godthab (64° N), 2.5 m at Disko Island (70° N), to 3.5 m at North Star Bay (77° N). The coasts facing the Polar Sea have almost no tide, only 0.3 m being recorded at Kap Morris Jesup (83° 40' N). Along the east coast the mean spring tide is about 3.0 m in the south, decreasing northward, to less than 1.0 m at Nordostrundingen, the northeastern corner of Greenland (81° N).

Edited version of a chapter published by Niels Nielsen in The World’s Coastline (1985: 261–265). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 6.3.1 Cliffs cut in a raised marine terrace. Cirques with rock glaciers are common along the west, north, and northeast coasts of Disko Island, and in many places the rock glacier front forms an advancing cliff. The upper edge of the plateau basalt lies about 900 m above sea level.

When the sea freezes in winter the tide is of great significance because of ice foot formation. On sedimentary coasts this leads to the formation of thermokarst on the backshore, and accumulation or erosion when parts of the ice foot, containing gravelly material, are driven up on the shore or rafted away from it. These processes result in the displacement structures characteristic of arctic beaches. Erratic rocks and alien sediment incorporated in the ice foot are stranded on the shore when it melts (Nielsen 1978). The ice foot may also influence rocky coasts by breaking off rock fragments in the springtime, when the warmer sea melts the ice foot from below, so that it collapses into the sea (Nielsen 1979). Occasional large waves are produced suddenly by calving ice cliffs or capsizing icebergs. Though rare, this type of wave is both high and long, and can result in the deposition of wave-transported sediment up to 15 m above sea level in a fiord where the fetch is very limited (Nielsen 1969). Throughout most of the Quaternary period, the ice cap covered Greenland, though varying in extent. In Late Pleistocene times, about 12,000 years ago, when the climate began to become warmer, the ice margin receded, sea levels were rising, and the reduction in ice mass resulted in rapid isostatic uplift in coastal regions. This relative upheaval of the land has led to the emergence of Late Pleistocene and Holocene coastlines, seen as raised marine terraces along many parts of the Greenland coast (> Fig. 6.3.2). In East and West Greenland the isostatic upheaval of land started about 9,000 years ago and apparently ended 4,000–5,000 years later (Hjort 1979). A similar development seems to have taken place in North Greenland,

although somewhat delayed by the later melting of the ice cap in this part of the country. Thus the formation of coastlines that are now about 10 m above sea level, has been dated to 3,000–4,000 years ago (Weidick 1972; Stablein 1975). Emerged coastal features include marine terraces and beach ridges that reach elevations of about 60 m in South Greenland, compared with 100–200 m in the central part of West Greenland, and up to 216 m above present sea level in East Greenland, indicating the pattern of the highest Holocene marine limit. The upheaval of land has varied from one region to the other because of local variations during the melting phase. Minor fluctuations in sea level relative to the land are primarily related to eustatic changes in sea level (Humlum 1980). After the Little Ice Age, the climatic deterioration that set in during the seventeenth century, a general sea level rise of about 1 m occurred, resulting in the submergence of ruined buildings. During the past 50 years there appears to have been a slight lowering of sea level relative to the land, but this may now have reversed. Ammassalik Island off SE Greenland is dominated by high ground, presumably upwarped along the contintental margin, with a narrow coastal plain. During the Quaternary the island was covered by the Greenland Ice Sheet, with the exception of a few outcrops on the escarpment edge. Periglacial processes have also been active, intensifying during the late Holocene (Humlum and Christtansen 2008). In Ikkafjord in southwestern Greenland are the so-called Viking Warriors, towers of tufa that grow up from the shallow floor of the fiord to a height of 20 m,

Greenland

6. 3

⊡⊡ Fig. 6.3.2 Raised marine terraces in Mellemfiord, west Disko. The border between the scree from the basalt scarp and the upper terrace edge lies 53 m above sea level, and was formed about 9,000 years ago.

occasionally protruding from the water on calm days. They can grow upward 0.5 m in a year, and have formed as the result of exudations of alkaline groundwater from sea floor fissures. Minerals dissolved in the groundwater crystallise into a mineral called ikaite, which is added to the top of the growing towers. The feature is geologically unique.

References Escher A, Watt WS (eds) (1976) Geology of Greenland. Geological Survey of Greenland Copenhagen, Denmark Hjort C (1979) Glaciation in northern East Greenland during the late Weichselian and Early Flandrian. Boreas 8:281–296

Humlum O (1980) Arctic geomorphology, Enoks havn, Disko (in Danish). Geogr Noter 5:1–142 Humlum O, Christtansen HH (2008) Geomorphology of Ammassalik Island, SE Greenland. Geografisk Tidsskrift 1089:50–20 Nielsen N (1969) Morphological studies on the eastern coast of Disko, West Greenland. Geografiska Tldsskrift 68:1–35 Nielsen N (1978) Coastal morphology in the Arctic region (in Danish). GO. Geograforlaget Brenderup:242–247 Nielsen N (1979) Ice-foot processes. Observations of erosion on a rocky coast, Disko, West Greenland. Zeitschrift für Geomorphologie, N.F. 23:321–331 Stablein G (1975) Ice-margin deposits and coastal evolution in West Greenland (in German). Polarforschung 45(2) Weidick A (1972) Holocene shorelines and glacial stages in Greenland – an attempt at correlation. Geological Survey of Greenland 41

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6.4 Iceland

Eggert Larusson

1.  Introduction Situated in the North Atlantic Ocean between 63° 30' N and the Arctic Circle (66° 30' N), Iceland is a volcanic island with an area of 103,100 sq km. It is the outcome of Tertiary and Quaternary extrusions and intrusions of volcanic rocks (lavas, tuffs, pyroclastic sediment) from the mid-Atlantic Ridge, a zone of crustal plate divergence. It has been glaciated, and there are five residual ice areas. Glacial drift deposits are found widely, particularly bordering the residual glaciers, where terminal moraines in­­ dicate recession of ice margins in recent decades. Much of the coastline is irregular, with fiords (known in Iceland as fjördur) between mountain ranges, but on the southern coast between the Olfusa estuary and Djupivogur, large volumes of sand and gravel deposited by powerful streams fed by melting glaciers have formed a wide lowland (sandur), with a smooth coastline of wide bays and broad lobate promontories (>Fig. 6.4.1). The lowland is backed by cliffs and bluffs marking a 9,000 year old cliffy coastline that existed before this extensive deposition occurred. The landscape of Iceland is dominated by basalt plateaux and glaciated ridges, with 10% still covered by ice. Emerged beaches are found on many parts of the Icelandic coast (Jonsson 1957; Norddahl and Petursson 2005). The height of the Postglacial marine transgression limit differs from place to place: 130–140 m in the west, but only 35–50 m in the north, the east, and the south. The highest Holocene coastlines have been dated at 12,300–9,600 bp. The sea reached about its present level by 9,000 bp and the isostatic uplift was complete 3,000 years later. In the interim, the Holocene marine transgression was more or less matched by the isostatic recovery. Before 3,000 bp, sea level seems to have stood about 2–4 m below the present level, and a small further transgression may then have taken place (Thorarinsson 1956). Roches moutonnées, striations and some till cover indicate that the Icelandic strandflats existed before the last glaciation. The climate is cool temperate oceanic. Reykjavik, the capital city on the southwest coast, has mean monthly temperatures of 0°C in January and 11°C in July, with an average annual rainfall of 800 mm. Affected by the ­numerous

arctic (polar front) cyclones, Iceland is windy. Wind speed and direction are very variable, but the most frequent winds blow between southeast and northeast, and the main track of the depressions passing close to the southern coast. Gale force winds are common. There are extensive areas of bare, unvegetated landscape, rocky or gravelly, as the result of strong winds, low temperatures and desiccation. Peat deposits are of limited extent, mainly around lakes and lagoons. Some shore ice forms in cold winters, especially in the northern fiords. The coasts of NW Iceland are less stormy than those of the south and east coasts. Prevailing winds and waves in coastal waters are from SE and NE rather than SW, while the southern coasts receive an Atlantic ocean swell from the south. Mean spring tide ranges are generally between 1.5 m and 3 m, and a little more along some parts of the western coasts, where Reykjavik has 3.8 m and Grindavik in the SW has 3.2 m. Bakkagerdi on the NE coast has 1.5 m and Bolungarvik in the NW has 2.4 m.

2.  The Icelandic Coastline The NW peninsula consists of much-dissected Tertiary and volcanic formations (the oldest in Iceland), and presents an indented rocky coastline of deep, steep-sided fjördur. There are very steep coastal slopes and megacliffs (300–600 m) on which landslides sometimes occur. Deep inlets branch from the fiord of Jokulfjördir, Leirufjördur on the southern side receiving glacifluvial sand and gravel from a stream draining out of the Drangajökull glacier. There are several large snow-filled corries and waterfalls below hanging valleys. On the southern side of Isafjadardjup are several parallel deep fiords (fjördur) running from south to north. The Isafjördur is the easternmost, with a grey sandy beach in the bay at the mouth of the Laugadalsa, but much of the fiord coast is rocky with angular gravels in coves. On the western side of Isafjördur is a cuspate spit, flanked by shingle beaches that fade out along the shore (>Fig.  6.4.2). Beaches and spits in these north-western fjördur are related to streams delivering shingle, or to eroding sectors

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

6.4    

Iceland

⊡⊡ Fig. 6.4.1 Some coastal features of Iceland. 1 – glaciers, 2 – subglacial volcanoes and/or hot zones, 3 – outwash plains (sandurs) near the coastline, 4 – track of main glacier bursts (jökulhlaups), 5 – beaches and barrier spits, 6 – predominant longshore drift, 7 – strandflats. (Courtesy Geostudies.) 1 4°

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⊡⊡ Fig. 6.4.2 Cuspate spit on the western shore of the Isafjördur, NW Iceland. (Courtesy Geostudies.)

Iceland  

of glacial drift. There are several cuspate spits, shaped by waves that approach diagonally across these long, narrow inlets. There are sandy beaches and flanking river mouths at the head of each fjördur, but in several sectors the coastline is artificial because of road construction and protective boulder walls. Cliffs are found towards the more exposed northern ends of the fjördur, and are high and bold towards Gardstadir. Skötufjördur, Hestfjördur, Seydisfjördur and Alfdafjör­ dur are parallel fjördur on the north coast. Sudavik stands on a morainic slope behind a beach and a lobate foreland. In January 1995, 22 houses were swept away by an avalanche, and a new town was then built to the south. The next deep inlet is Skufulsfjördur, with high volcanic cliffs and slopes, and an incipient corrie on its eastern slopes. The town of Isafjördur stands on its western side, a port built on a former recurved spit, protected from northerly storms by a large boulder wall. West of Isafjördur, the coast is steep on glacial drift, with screes and slumping slopes. The beach is of dark volcanic sand and gravel with splays of pale brown calcareous sand, backed by low grassy dune ridges, then a scatter of small dune hummocks (>Fig. 6.4.3). Bolungarvik harbour is protected by a large boulder breakwater. Round a high promontory is Sugandafjördur (>Fig. 6.4.4), with the Stadur valley on its southern side. Steep slopes rise to a high plateau, and at the valley mouth is the cove at Skeravik (>Fig. 6.4.5). To the east is Sudureyri, a fishing port. The upper part of Sugandafjördur is shallow, with gravelly shoals exposed at low tide. ⊡⊡ Fig. 6.4.3 The shore at Bolungarvik, NW Iceland. The dark beach is of basaltic sand and gravel, the paler material calcareous sand from the sea floor. (Courtesy Geostudies.)

   6.4

On the north coast of Onundarnfjördur, the town of Flateyri has been built on a spit that curves upfiord. At the head of this fjördur the beach has a covering of brown comminuted shell sand. At Holt is an inner spit, concave seaward, which has a brown calcareous sandy beach bac­ ked by grassy dunes and partly eroded hummocks. Opening to the west coast is Dyrafjördur, with Pingeyri on a sandy cuspate spit, with a brown sandy beach on the southern side. Upstream the beaches and spits are of grey volcanic sand and gravel. Arnarfjördur is steep-sided, with some minor spits, and cliffs over 500 m high at Selardalshidar face WSW into the open ocean. It branches headward into several smaller fjördur, one of which is Dynjandisvogur, where steep coasts in glacial drift slope down to beaches of grey sand and gravel with several cuspate spits. A spit has grown half way across Talknafjördur. In Patreksfjördur, the town of this name stands on two sand spits, Vatneyri and Geirseyri on the northern shore, while Sandoddi is a recurved spit on the southern shore. At the western end of the Bardastrandar Peninsula Blakknes is a 281 m headland and Breidavik has a curved beach. High cliffs run along Bjarnarnupur, and on the southern side of the peninsula the cliffs of Latrabjarg rise to 440 m. Raudisandur is a long barrier enclosing a marshy area, and Baejarvadall is a coastal lagoon. To the east, high cliffy peninsulas separate a succession of fiörden on the northern coast of the broad Breidafjördur. Breidafjördur is a submerged strandflat, with numerous islands, skerries, and shoals. One of the islands is

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⊡⊡ Fig. 6.4.4 Sugandafjördur, NW Iceland. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.5 Stadur, west of Skeravik, looking towards Saudanes. The steep coastal slopes descend from the coastal plateau to an emerged terrace. (Courtesy Geostudies.)

Flatey, with a low cliffy shore where columnar basalts are disintegrating into aprons of basalt boulders. On the east coast of Breidafjördur, a strait narrows to Gilsfjördur, and to the south is the mountainous Klofningsvegur Peninsula, then the Hvammsfjördur strait.

On the Snaefellsnes Peninsula to the south the width of the emerged strandflat is generally 10–20 km, backed by steep bluffs. South of the Hvammsfjördur is the Skog­ arströnd coast, which has a terrace incised by rocky gorges, such as those cut by the Dunkur and Skrauma rivers.

Iceland  

Gravelly watercourses cross the coastal plain, and there are seaweed-strewn shores of rocky lava, and few beaches. An archipelago includes the large embayed island of Brokey. Eyrarfjall is a high volcanic ridge beyond which two valleys converge to the Osa. Alfafjördur is a major inlet, backed by a salt marsh and an emerged fjördur floor ringed by rocky cliffs. Stykkisholmur is a ferry port on a peninsula on the south coast of Breidafjördur. To the west is the coastal district of Helgafellssveit, with a terrace that descends northward to low cliffs and rocky shores with gravel at river mouths. Incised streams tumble over waterfalls to the sea on this emerged coast. There is a wide lowland with an extensive spiky lava field, and inlets with strong tides. Pass Eidi leads westward through to the bay of Grundarfjördur. West of Grundarfjördur is the mountainous peninsula of Kirkjufell (Church Mountain), 463 m high. At Buland­ shöfdi, there are said to be marine shells 40–55 m above sea level. Gamlavik is a broad bay backed by two barrier lagoons separated by a delta. At Ondverdarnes, the coast turns southward along the dark Svörtuloft cliffs, cut in Holocene lava flows. Similar cliffs fringe the western part of the Snaefellsnes Peninsula, which is dominated by blocky lava deposited within the past 11,000 years. On this coast, beaches are found only where there is a nearby source of sand or gravel, either a cliff or a river mouth.

⊡⊡ Fig. 6.4.6 A boulder beach derived from lava outcrops in the cliffs at Hellnar. (Courtesy Geostudies.)

   6.4

At Hellnar, the cliffed coast turns north-eastward. There is a sharp contrast between smooth pastoral land on tuff at Hellnar and the irregular bouldery lava flow to the north. Cliffs cut in tuff come to an end at the harbour, ­giving place to irregular lava cliffs dissected by inlets and a natural arch. The harbour contains a sandy beach, derived from tuff in the adjacent cliffs and on the sea floor. A boulder beach of dark and pale volcanic rock occupies a small bay to the north of the breakwater (>Fig. 6.4.6). East of Arnarstapi is the Midhusavatn lagoon, separated from the sea by the wide sandy Hraunlandarif barrier. Further east a hummocky lava flow descends from the escarpment at Axlarholar, out across the gently sloping coastal lowland to form the lobate foreland of Budarhaun. This is a fine example of a lobate foreland produced by volcanic deposition, and includes a cinder cone. Budir stands at the eastern end of the rugged Budahraun lava field. There is a tidal lagoon, Budaos, and a curved beach of light brown calcareous sand backed by grassy dunes extends eastward past Bardastadir, where an inlet is bordered by low cliffs exposing stratified lagoon deposits that had been overrun by dunes on a receding coastline (>Fig. 6.4.7). Stakkhamar is a gravelly plain, and a chain of sandy barrier islands extends round to Hitarnes. The coastline then turns southward. There are extensive beaches and dune-capped barriers fringing lagoons, some of which are behind wide grassy strips that may include glacial drift. South of the Hafjardara River, the

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⊡⊡ Fig. 6.4.7 Stratified lagoon deposits exposed on the shore at Bardastadir. (Courtesy Geostudies.)

coast is backed by the wide Myrar coastal plain that is like a strandflat, but includes areas of glacial and fluvial deposition. It has recent knobbly or bouldery lava flows, and the Eldborg volcano rises prominently. The coastal plain is interrupted southward by the Borgarfjördur inlet at Borgarnes. On its northern shore, the Langa River cascades over lava ledges, as it drops to the sea. To the south, skerries fringe the low-lying coast past Höfn, and a bold volcanic escarpment overlooks the coastal plain. Hvalfjördur is fringed by strandflats, with steep headwalls. South of Hvalfjördur the Kjalarnes Peninsula is backed by a volcanic escarpment with screes and gullies that feed into glacifluvial channels incised across a coastal lowland that ends in a low cliffy coastline. The hilly embayed Reykjavik Peninsula is urbanised and bordered by sea walls and boulder ramparts, and a large harbour. There are outlying grassy low islands. South of Reykjavik is the Keflavin Peninsula, a grey mossy and grassy Holocene plain with volcanic cones and dykes, and low rocky shores. This is a continuation of the mid-Atlantic Ridge and is consequently a young volcanic area, with lavas deposited in the past 11,000 years. The southern coast is exposed to strong Atlantic swell and storm waves. Low cliffs cut in Holocene lava flows predominate, and high level storm beaches of loose boulders and pebbles that have been piled up on top of some of the cliffs by the very large waves that break upon this coast in winter storms. Beaches of well-rounded cobbles and

boulders dominate the shore. Sandy beaches are uncommon, but small isolated dune fields can be found. To the east, a coastal plain backed by a line of cliffs and bluffs marking an early Holocene coastline. Several rivers, some fed by hinterland glaciers, flow southward across a sand and gravel outwash coastal plain (Hine and Boothroyd 1978). From the Pjorsa River a straight sandur coast extends eastward to the large lobate foreland in the lee of the Vestmannaeyjar islands. Southwest of this archipelago a new volcanic island, Surtsey, began to grow in 1963. After the volcanic activity came to an end the island was re-shaped by storm waves which cut back cliffs and deposited beaches of derived sand and gravel. Recession of the cliffs was most rapid on the exposed southwest coast, whereas a cuspate foreland was built on the more sheltered northern coast by convergent longshore drifting (Norrman 1980). The coastal plain, backed by lava cliffs, continues eastward, and the old coastline becomes a grassy slope up to the edge of the Eyjafjallajökull glacier, which lies over a volcano that erupted in 1612 and 1821–1822, causing massive ‘glacier burst’ flooding and the deposition of the wide sandur plain. The 60 m high Skogafoss waterfall pours over the cliff into the River Skoga. The coast is fringed by gravel beaches with some dunes (Mountney and Russell 2006). At Dyrholaey, a spur of high ground runs out to form a 110 m high cliffed headland, with a natural arch. Behind Dyrholaey, a lagoon is impounded by a barrier spit, and a

Iceland  

high rocky ridge of Upper Pleistocene pyroclastic and tuffaceous sediment (Reynisfjall) runs south to a prominent 200 m high headland. On the next headland there is a large cave in columnar lava flows (>Fig. 6.4.8) and The Needles (Reynisdrangar) are rocky stacks 66 m high close inshore (>Fig. 6.4.9). A high scarp then runs inland besides Vik.

⊡⊡ Fig. 6.4.8 Columnar lava on the coast at Dyrholaey. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.9 The Needles (Reynisdrangar), stacks off the cliffed coast near Vik. (Courtesy Geostudies.)

   6.4

Vik is not a coastal town, and there are no harbours along this coast. At Kerlingardalur there are hummocky dunes on the coastal plain, behind a wide black sandy beach (>Fig. 6.4.10). East of Vik is the wide sand and gravel plain of Myrdalssandur, with lava blisters, extending to the braided

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⊡⊡ Fig. 6.4.10 Coastal dunes east of Vik. (Courtesy Geostudies.)

estuary of the Kudafljot River, which is fed by meltwater from the Myrdalsjökull glacier. Eruptions of Katla have occurred every 40–80 years, melting the Myrdalsjökul glacier and sending catastrophic floods southward over the Myrdalsandur. The last of this glacier burst floods (jökulhlaups), occurred in 1918, so another glacier burst is overdue. The bluffs behind the coastal plain decline very gradually beneath lagoons behind barriers, as at the 76 m high island of Ingolfshöfdi (King 1956), passing through broad wet sandy flats with a wide zone of low mounds of windblown sand. The lagoon is more than 6 km wide but only up to half a metre deep. Beaches of dark volcanic sand alternate with zones or ridges of cobbles. Predominant longshore drifting diverges westward and north-eastward from Ingolfshöfdi. On the east coast near the Gljufura canyon there are terraces and barrier-enclosed lagoons, and to landward, is a scarp capped by the glacier lobe of Breidamerkurjökull, part of the huge Vatnajökull ice cap. To the north, the coastal lowland has ridges of glacial drift backed by a glacier lake, Fjallslon, with a clear ice front of the Fjallsjökull lobe, several icebergs and a steep sided cliffed moraine of gravel and boulders. At the Jökulsa River, another morainic ridge encloses Jökulsarlön, a lake crowded with icebergs. The Breidamerkurjökull glacier is melting into a coastal lagoon (>Fig.  6.4.11), which has an outlet to the sea through the sand and gravel barrier. Comparison of maps made in 1945, 1965 and 1998 show that the glacier front has receded by up to 3 km (Evans and Twigg 2000).

The coast of Skeidararsandur has few beach ridges, no tidal inlets, no estuaries, no true lagoons, and no marshes, because of the high rates of fluvial and shore sedimentation. The so called Hornafjördur, Skardsfjördur, Papafjördur, Lonsfjördur, Alftafjördur and Amarsfjördur are shallow lagoons, not true fiords, surrounded by low lying land (strandflats and sandurs). Large volumes of material are deposited in the lagoons, and glacifluvial deltas are presently growing. Tidal inlets are open, and narrow marshes and tidal flats fringe parts of the lagoons. Between Skardsfjördur and Papafjördur, a high rocky promontory of Tertiary lava interrupts the barrier beaches. Another such mountainous sector of steep rocky coast with a narrow strandflat extends between Lonsfjördur and Alftafjördur. North of Djupivogur is a typical fiord coast, with some irregular strandflats, as at erufjördur. Stodvarfjördur, Fastrudsfjördur and Reydarfjördur are fiords at the mouths of glaciated valleys (>Fig. 6.4.12), with intervening ridges ending in cliffy promontories. More fiords are seen north to Glettinganes, where the steep coast turns NW. Heradsfloi, a bay on the NE coast, is a typical sandur fed by rivers draining north from the Vatnajökull glacier. Beyond Vopnafjördur, the northeast coast has lower promontories and wider bays and the narrow Langanes Peninsula declines to Fontur, the northeast point of Iceland. The northern coast of Melrakkasletta is very irregular, with a succession of promontories and bays with lagoons and lakes separated from the open sea by wide barrier embankments of boulders.

Iceland  

   6.4

⊡⊡ Fig. 6.4.11 Breidamerkurjökull glacier calving into a coastal lagoon near Jökulsa. Breidamerkurjökull. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.12 Glaciated valley ­descending to Mjoifjördur on the east coast of Iceland. (Courtesy Tony Stutterd.)

South of the wide bay of Axarfjördur, a broad sandur is occasionally reached by glacier-burst floods, and some old barrier beaches indicate that there has been coastal progradation. Eyjafjördur is a deep fiord and on the east coast of Skagafjördur is the double tombolo of Thordarhöfdi, linking an island over 200 m high (Bodéré 1973). At the southern end of Skagafjördur barrier beaches and spits block river mouths on either side of the Hegraxes ridge, which protrudes northward.

Beyond the next promontory is the wide bay of Hunafjördur, the southern part of which is split by a protruding steep-sided ridge into Hunafjördur to the east and Hrutafjördur to the west. Hunafjördur has a narrow 7 km long and up to 5 km wide sand barrier (Pingeyrasandur) that divides the Hop lagoon (Hunafloi). At the western end of this beach is the strange stack of Hvitserkur, the remains of a volcanic neck pierced by natural arches and daubed with white gauno.

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Hrutafjördur, by contrast, is a long fiord that narrows southward to a steep-sided valley at Melar. On its west coast are several large tributary fiords, notably Stein­ grimsfjördur and Reykjarfjördur, then the mountainous promontory at Krossnesfjall. To the north–west the fiords are smaller between sharp, narrow peninsulas along the coast to Hornbjarg, where the north-facing cliffs rise to more than 550 m.

References Bodéré JC (1973) Le tombolo double de Thordarhöfdi. Norois 78:213–235 Evans DJA, Twigg DR (2000) Breidamerkurjökull 1998 map (1: 30,000). University of Glasgow and Loughborough University, UK

Hine AC, Boothroyd JC (1978) Morphology, processes, and recent sedimentary history of a glacial outwash plain shoreline, southern Iceland. J Sediment Petrol 48:901–920 Jonsson J (1957) Notes on changes of sea level in Iceland. Geogr Ann 39:143–212 King CAM (1956) The coast of south east Iceland near Ingolfshöfdi. Geogr J 122:241–246 Mountney NP, Russell AJ (2006) Coastal aeolian dune development, Sólheimasandur, southern Iceland. Sediment Geol 192:167–181 Norddahl H, Petursson HG (2005) Relative sea level changes in Iceland: new aspects of the Weichselian glaciation of Iceland. In: Caseldine C et al (eds) Iceland: modern processes and past environments. Elsevier, Amsterdam, the Netherlands, pp 25–78 Norrman JO (1980) Coastal erosion and slope development in Surtsey Island, Iceland. Z Geomorphol 34:20–38 Thorarinsson S (1956) The thousand years struggle against ice and fire. Bokautgafa Menningarsjods, Reykjavik, Iceland

6.5 Other Atlantic Ocean Islands

Jim Hansom

1.  Introduction

2.1.  The Azores

A succinct review of the Geomorphology of Atlantic Ocean Islands, including > Bahamas, > Bermuda, > Greenland, > Iceland can be found in Hansom (2005). The Atlantic Ocean islands fall into three groups (>Fig.  6.5.1): those that lie on the Mid-Atlantic Ridge, from Iceland south to Bouvet Island; those that lie on the American shelves to the west, such as Bermuda and the  Falkland Islands; and those that lie on rises and the Eurafrican Shelf to the east, from the Faeroes south to the islands in the Gulf of Guinea. In stark contrast with the low coral islands and cays of both Indian and Pacific oceans, the Atlantic is mainly characterised by high islands, many of them of volcanic origin. Although extensive in Caribbean waters, in general, coral reefs are poorly developed in the Atlantic Ocean and even the Atol das Rocas, off Brazil, is not a true coral atoll. The Atlantic islands show variations in size and shape, and features that reflect contrasts in their lithology, in their vulcanicity (whether active, dormant or extinct), and in climate and vegetation (the lush tropical Azores to the glaciers of Bouvet Island). Details of the geology and coastal geomorphology of the Atlantic islands are ­discussed in Mitchell-Thome (1970, 1976) and Hansom (2005).

The Azores (Portuguese) (38° 30' N, 28° 00' W – 1,500 km west of Lisbon) consist of nine widely separated mountainous, but fertile islands (total area, 2,344 sq km). The temperate oceanic climate is mild (January 14°C, July 22–33°C), but with abundant rainfall (about 950 mm/year, winter maximum), vegetation cover is extensive. The steep slopes of the islands are dissected by deep and narrow gulleys, but relics of emerged marine beaches and platforms are found at various levels of up to 60 m (Guilcher and Battistini 1982). The coasts are generally steep, with cliffs up to 600 m high, cut into lavas, tuffs, and scoria, sometimes delivering gravel to the shore via scree slopes, but there are also many small beaches of dark volcanic sand. Some white calcareous beaches are also found. The sea is frequently stormy, particularly on north-facing coasts, and the harbour at Ponta Delgada on Såo Miguel is artificially protected by tetrapods. With deep water close inshore (the 100 m isobath is mainly within 3 km of the coastline), the power of the waves reaching the Azores is such that Pico was chosen as the site of a pilot wave-power plant. The coast of the largest island, Sao Miguel, is characteristic of the Azores coastline, in general (>Fig.  6.5.2). The coast is backed by higher volcanic peaks, where steep cliffs fronted by caves, patchy low shore platforms, offshore stacks and islets are found, such as at Mosteiros. Between the cliffed indentations, short sandy beaches occur. Longer sandy beaches occur where the hinterland is lower, such as at Praia dos Moinhos and at Populo, where small pine-clad sand dunes occur. Pico Island is 42 km long and dominated by the cone-shaped stratovolcano of Ponta do Pico whose smoking crater lies at 2,351 m. Various historical lava flows have moved the coastline seaward but have now been eroded back to form low cliffs. There are 600 m high cliffs on the west coast of Flores, which has seven volcanic craters. in sheltered places from the dominant westerly waves smaller volcanic forms survive on the coast, such as at Barca on the island of Graciosa, where a volcanic cone has been eroded into bays fringed by crumbling cliffs of ash that cascade onto a boulder beaches below. The 1957/1958 eruption at Capelhino on Faial showered the adjacent coast

2.  Islands of the Mid-Atlantic Ridge All of the islands of the mid-Atlantic Ridge are of volcanic origin, having been produced by the upwelling of basaltic magma into the crustal gap at the mid-ocean spreading centre in the mid-Atlantic. Many remain active volcanoes, rising from deep water, their coastlines characterised by steep cliffs cut into lava, ash or hyaloclastite. The volcanoes are all youthful and the island coastlines are in a state of continuing adjustment in response to ongoing volcanicity, marine and slope activity. The islands extend south from Iceland and include the Azores, St. Paul and St. Peter Rocks, Ascension, St. Helena, Tristan da Cunha, Gough Island and subantarctic Bouvet Island.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 6.5.1 Location map.

in ash and contributed to accretion of the beach. At Porto Pim, also on Faial, a substantial sandy spit has developed.

2.2. St. Paul and St. Peter (São Paulo e São Pedro) Rocks More than 800 km from South America, these 12 Brazilian islets (0° 56' N, 29° 22' W) are composed of mylonitic

­ eridotite and are thought to be unaltered mantle matep rial. Dated isotopically at 3,500–4,500 million years, they are among the oldest outcrops on the earth’s surface (Smith et al. 1974), the submarine mountain of which these rocks are the pinnacles extending 4,000 m into the ocean depths. The only source of freshwater on the St. Peter and St. Paul’s Rocks is from rain, and there is almost no vegetation. The lack of vascular vegetation is most likely due to constant sea-spray bombarding the rocks, and the terrestrial

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⊡⊡ Fig. 6.5.2 Stormy weather in the Azores on the north coast of Såo Miguel Island, near Capela. (Courtesy A. Stutterd.)

v­ egetation is dominated by fungus and algae. The rocks bear extensive deposits of white guano: calcium phosphate derived from seabird droppings. The highest point is only 23 m above sea level, and there is an emerged marine platform at 5 m. The coasts are cliffed and rocky, with wavescoured potholes.

Lithothamnion) growing in the narrow shallow water fringing the island. Erosion of adjacent rock shores and terrestrial sediment in-washed by infrequent rains probably account for the remainder of the beach sediment.

2.3.  Ascension Island

St. Helena (British) (15° 58' S, 5° 43' W) lies slightly east of the mid-Atlantic Ridge, and is an isolated volcanic peak rising to a total of 5,223 m from the sea floor to its highest point at Diana Peak. With an area of 122 sq km it is a composite extinct volcanic cone, deeply eroded with radiating valleys intersected by deep valleys. The coastline ­comprises high stepped and sometimes vertical cliffs cut by steep v-sided valleys. The highest cliffs are 670 m at Flagstaff Bay in the north, where the valleys are marked by 200 m high coastal waterfalls, and at Man and Horse Cliffs in the south– west where they reach 580 m. The vertical steps in the coastal cliffs often correspond to differences in the composition of the outcropping lava flows. At Great Stone Top, Turks Cap and Prosperous Bay on the east coast, thick flows of trachyte tform vertical faces which ­cascade debris onto stone chutes below, some of which extend into the sea to form fringing small boulder beaches. Emerged shore platforms stand at 4–6 m above sea level and the numerous offshore islets and skerries common on the south and west coast, such as at Egg Island and Cockburn Battery, are almost absent elsewhere.

Ascension Island (British) (7° 57' S, 14° 22' W), lying in mid-ocean about 1,300 km north of Saint Helena, is the peak of a dormant Holocene stratovolcano that rises to a total of 3,860 m from the sea floor to its highest point at Green Mountain. More than 100 vents are found within its 98 sq km area and although eruptions continued into Quaternary times, there are no known historic eruptions. The climate is arid and vegetation is sparse, with scattered guano deposits. Imposing sea cliffs line the more exposed eastern coast, where Boatswain Bird islet, a monolithic stack of trachyte 98 m high, has a natural arch. The west coast comprises low lava fields, whose ­serrated seaward edge is truncated by low cliffs fronted by narrow and uneven shore platforms (>Fig. 6.5.3). However, Ascension also has superb white sand beaches derived from shell and coral. Little seems to be known of the nature of the sediment supply to the Ascension beaches, but the carbonates can only come from shells and coralline algae (such as

2.4.  St. Helena

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⊡⊡ Fig. 6.5.3 The west coast of Ascension Island is characterised by low gradient lava flows, whose termini have been eroded into low cliffs fronted by uneven platforms.

v­ egetation occurs above 1,350 m. The only township, Edinburgh, is located on a low-lying platform called Settlement Plain on the northwest side of the main island. Emerged beaches, platforms and caves are found at 5 m above sea level on Tristan and at 12 m above sea level on the small adjacent islands, indicating differential tectonic uplift. The cliffs are lower along the Edinburgh coast and a narrow gravel beach has developed along this stretch of coast, fed by material eroded from the cliffs. In the months following the 1961 eruption at Big Point, rapid recession of the cliff was noted (10 m in 8 weeks), this contributing to accretion of the downdrift gravel beaches which formed a spit enclosing a small lagoon. In some places, small and inaccessible beaches are found, some of which are sandy. Four small islands are associated with Tristan: Inaccessible (10 sq km; peak, 548 m), Nightingale (2.6 sq km; peak, 396 m), Stoltenhoff (0.25 sq km; peak, 105 m), and Middle (0.26 sq km; peak, 65 m), all are bordered by high lava and tuff cliffs with small and narrow sand and gravel beaches.

2.6.  Gough Island

Only in favoured locations, where the deep gullies reach the coast do small beaches, some of sand occur, such as at Lemon Valley Bay. At Sandy Cove on the south coast, the sandy beach is backed by calcareous sand dunes.

2.5.  Tristan da Cunha Tristan da Cunha (British) (37° 15' S, 12° 30' W) is a 98 sq km circular volcanic island, some 350 km NW of Gough Island. The base of the central volcanic peak lies 3,700 m deep on the sea-bed and the summit rises to almost 2,100 m above sea level at Queen Mary’s Peak. The lower slopes are almost entirely lava and form high cliffs that surround the island. Queen Mary’s Peak has about 30 parasitic scoria cones encircling the main crater with bordering cliffs up to 900 m high. The island is bleak and barren, but zoned

Gough Island (British) (40° 30' S, 10° 00' W), is also known as Diego Alvarez and lies 2,500 km east of the Cape of Good Hope. It is the eroded summit of a deeply dissected Tertiary volcano. The island is mountainous, rising to 910  m above sea level, and the lower slopes are covered with a luxuriant evergreen scrub forest that thins out above the 300-m contour. The steep slopes are cut by deep gullies or gulches descending to hanging valleys and waterfalls, and the coastline is characterised by steep cliffs that have been undercut to produce a variety of narrow boulder beaches at the foot of the cliffs. There are also lava arches, numerous islets, stacks (such as Tristania Rock (165 m)) and skerries that mostly lie within 100 m of the coast. Landings are prohibited and the islands bouldery beaches are ideal breeding grounds for fur seals.

2.7.  Bouvet Island Bouvet Island (Norway) (60 sq km; peak, 935 m) (54° 00' S, 3° 00' E) stands at the southern end of the mid-Atlantic ridge, and has a rugged ice-mantled Quaternary volcanic terrain with five glaciers, one of which ends on the east coast in ice cliffs up to 122 m high (>Fig. 6.5.4). The icecapped dome rises to 935 m. There is a central ice-filled crater, and the surface is more than 90% ice-covered. The island is extremely remote: there is no land in any direction for

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⊡⊡ Fig. 6.5.4 Bouvet Island. (Courtesy R. Muench, Earth and Space Research.)

more than 1,500 km. The adjacent seas have pack ice in winter, but are disturbed by frequent gales. Basalt cliffs cut back by stormy seas attain 490 m on the more exposed west coast, where in 1955–1957 an eruption or uplift added a new lava shelf 25 m above sea level from which a gravelly beach has been derived. Thompson Island, charted in the nineteenth century to the northeast, is thought to have disappeared as the result of a volcanic explosion.

3.  Islands of the Western Atlantic 3.1. South Shetland, South Orkney, South Sandwich and South Georgia Islands The Scotia Arc is a tectonically disrupted link between the fold mountain belts of South American cordillera and the  Antarctic Peninsula, and includes Late Palaeozoic– Mesozoic sedimentary formations with granitic intrusions and associated volcanics. The combined area of islands of the Scotia Arc is about 10,000 sq km, with peaks ­exceeding 2,000 m. South Shetland (4,700 sq km) is barren and heavily glacierised, the South Sandwich Islands (British) (310 sq km) have active volcanoes, and South Georgia (British) (3,756 sq km) is glacierised and mountainous, rising to 2,934 m at Mount Paget. South Georgia, South Orkney (British) and South Shetland Islands (British) contain

metamorphic gneisses, schists, and marble, but the South Sandwich islands are composed entirely of Cenozoic volcanics, and a submarine eruption north of Zavodovski Island in 1962 generated pumice deposits, which are found on the narrow beaches and were distributed far afield on the Antarctic Circumpolar Current. The islands are all barren and rugged, with glaciers, ice caps and Alpine peaks. Their coastlines are generally cliffed and rocky, but many inlets are found and plentiful glacial and glacifluvial debris provides the sediment sources for beach development. Rapid cliff retreat rates of 1 cm/year have been calculated for the cliffs of Fildes Peninsula in the South Shetland Islands, the result of efficient frost shattering. These coasts are affected by sea ice for up to 8 months of the year and the relative order associated with wave processes is annually swept away and replaced by the non-selective ­bulldozing and grinding action of floating ice, rendering beach landforms disorganised. In the South Shetland Islands, the coast is either rock or ice-cliff although some glacier termini rest on a low rocky platform at sea level. However, significant beach development occurs on the peninsulas that protrude beyond the ice cover and peninsulas on King George Island and Livingston Island are adorned with extensive emerged beach deposits, some of which connect offshore islets and skerries to the shore platforms of the main island. Prominent shore platforms are found at several altitudes including at

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present sea level, where extensive sub-horizontal shore platforms in sheltered locations give way to steeper, ramplike platforms in more exposed locations. This is a function of the persistence of floating ice blocks that horizontally abrade the surface in sheltered spots and wave activity that produces ramps in exposed sites. The upper part of the platforms is often adorned with an undulating layer of boulders that have been smoothed and compacted into extensive boulder pavements by the action of floating ice (Hansom and Gordon 1998). Some islands remain volcanically active and Penguin, Bridgeman and Deception are relatively new and active volcanoes. The sheltered inner crater of Deception Island has been inundated and has sand and gravel beaches in sharp contrast to the steep rocky outer coast. The coastline of the South Orkney Islands is formi­ dable and three of the four largest islands, Coronation, Powell and Laurie are dominated by large ice caps that spill over steep rocky cliffs, which plunge below sea level.  A  few bouldery beaches are found, for example, on Laurie Island. The remaining island, Signy, is low and rolling with a rocky sloping shore only with a few bouldery beaches. The South Sandwich group consists of a 400 km-long volcanic island arc that rises from a deep ocean trench. The 11 main islands are all volcanoes and the coastline is steep and cliffed with infrequent small bouldery beaches. Reefs and skerries abound especially between the smaller islands

in the extreme south at Thule, Cook and Bellingshausen islands. South Georgia lies on the northern part of the Scotia Arc, with several peaks rising to over 2,700 m. Long ridgelike and moraine-draped peninsulas are separated by deep glacial troughs (>Fig.  6.5.5), with the south coast more heavily glacierised than the north with steep cliffs and few small cliff-foot gravelly beaches. Glacier cover is more restricted on the north coast and the substantial ice-free peninsulas are flanked by high and steep cliffs, often with narrow shore platforms at their base together with numerous offshore islets and skerries. Several emerged marine platforms and beaches are found that relate to the glacial history but a systematic study of distribution and altitude is absent. The most extensive beaches of the Antarctic region lie on the north coast of South Georgia. Numerous glacifluvial streams deliver substantial amounts of material to extensive sand and gravel beaches, the largest of which are found at Salisbury Plain, Fortuna Bay, Cumberland Bay and St. Andrews Bay (Hansom and Gordon 1998).

3.2.  Falkland Islands The Falkland Islands (British) (52°S, 58°W) are also called Islas Malvinas and have a total area of 11,718 sq km. The coastline of the Falkland Islands, like the islands themselves, is low, flat and reminiscent of the coastline of the ⊡⊡ Fig. 6.5.5 The north coast of South Georgia has morainic plains fronted by gravelly beaches, in sharp contrast with the steep and heavily glacierised south coast.

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Outer Hebrides of Scotland. Although hundreds of smaller islands exist, the main island group comprises West and East Falkland separated by the steep-sided, fiord-like Falkland Sound (San Carlos Strait). Mt. Adam on West Falkland reaches 698 m. The intricate and crenulate nature of the lengthy coastline is related to the submergence and faulting of the underlying Devonian, Carboniferous and Permian limestones and sandstones. The landscape is treeless with extensive peat cover, and prominent periglacial stone runs of Pleistocene age are found (Hansom et al. 2008). In spite of the majority of the coastline being rocky, steep and high cliffs are mainly absent except in the west at Beaver Island where 100–200 m cliffs are found. In general, coasts are subdued, although the frequent westerly gales produce rapid recession where the rocky coast is replaced by periglacial sediment. Streams, although abundant and rarely dry, do not deliver much sediment to the beaches and the larger beaches are fed by shelly sand derived from nearshore shallow water. Significant beach development occurs on the outer coast and in many places strong winds have resulted in sand dune development (>Fig. 6.5.6). Many small sandy spits and tombolos connect small islets to the mainland.

Atol das Rocas (3° 50' S, 33° 40' W), (7.1 km2) lies 250 km off the Brazilian coast, but contrary to its name, it is not a coral atoll but an oval algal bank capping a seamount and enclosing a 6 m deep small lagoon. Two low sandy islands, Farol (106,000 sq m2) and Cemeteria (53,200 sq  m2), rise above the intertidal platform of algae, and have a capping of dune calcarenite. The highest point is a sand dune in the south of larger Farol Cay, at 6 m. Both islets are overgrown with grasses, bushes and palm trees.

3.3.  Trindade and Martin Vaz Rocks

3.5.  Fernando de Noronha

Trindade (8 sq km; peak, c. 600 m) (20° 30' S, 29°W) and the nearby Martin Vaz rocks are a complex of successively built volcanoes, the last (Paredo) a Holocene cone 217 m

The Fernando de Noronha archipelago (Brazilian) (3° 50' S, 32° 25' W) is volcanic and, with a main island (16.9 sq km) and about 20 associated islets, it is the emergent part

⊡⊡ Fig. 6.5.6 Sandy beaches on the east coast of East Falkland. Shallow water nearshore promotes beach development and strong winds carry the sand landward to produce dunes.

high on the southeast coast. The margins show high receding cliffs and truncated ravines, and a wave-cut tunnel 130 m long has been excavated in a volcanic headland. The exposed northeast coast has a coral reef, and there are beaches and dunes of calcareous sand as well as older dune calcarenites. Principe, on the south coast, has a sandy beach in front of a degraded cliff, and there is widespread evidence of about 3.5 m of emergence (platforms and beach rocks).

3.4.  Atol das Rocas

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of a large cone that rises 4,000 m from the sea floor. The relief is gentle, but the coastline is rocky and cliffed, with headlands in phonolite and bays carved in tuff. Wave action is strongest on the southeast coast, where Lith­ othamnion reefs have developed, while on the rugged north coast are beaches of sand, gravel and boulders, and some tombolos. Dune calcarenites and marine limestones are found above present sea level on the east coast. The tide range, generally small on oceanic islands, here attains 3.2 m at springs.

3.6.  Bermuda Bermuda (54 sq km) lies about 1,000 km east of Cape Hatteras on the east of the United States. It is a group of 150 limestone islands resting on the southern margin of a 650 sq km platform that is presently submerged to depths of 20 m. Beneath a thin capping of calcarenites, lava occurs at a variety of depths and the platform of volcanic rocks supports the northernmost coral reefs in the North Atlantic. Four geomorphological-ecological provinces are found: the 20 m depth reef-front terrace; the main 4 m deep coral-algal reef; the 16 m deep lagoon; and a northeast trending chain islands near the southern edge of the platform (Vacher 1973). The exposed limestones are mainly aeolian calcarenite, with only 10% beach and shallow water calcarenite. The supply of reef-derived sediment to beaches and dunes continues today but the present ­supply appears to be more restricted than in former times, since only on the island’s south shore do extensive beaches  backed by dune exist. These extensive sandy beaches are pink because of large numbers of fresh Homotrema clasts derived from the reefs, 1 km from the shoreline. Much of the south coast is cliff with only limited sediment. The sediment of the lagoon now plays little part in beach supply. This may be the reason why the lagoonfacing shoreline is erosional with a line of cliffs and few pocket beaches.

3.7.  Sable Island Sable Island (Canada) (43° 57' N, 60°W), lying on the shallow Sable Island Bank on the continental shelf 150 km southeast of Nova Scotia, is a low and wind-swept chain of sandy islands famous for their sandy shoals (>Fig. 6.5.7) and shipwrecks. The chain is 40 km by 1 km and of a variable but arcuate configuration. The Sable Island Bank is thought to be re-worked glacial sediment sourced from the maximum ice positions of the Late

Wisconsin and later readvances (King 2001). The islands have spectacular but desolate sandy beaches backed by sand dunes up to 30 m high (Byrne and McCann 1995). The dunes are primary in-situ dunes together with secondary dunes that have migrated across the island, a process exacerbated by a reduction in vegetation cover produced by the grazing of introduced ponies. Erosion during storms has caused eastward migration of 14.5 km over the past two centuries and necessitated the relocation of the lighthouse, although other buildings have been buried. However, deposition also occurs, and over the years, the land area has remained the same at about 30 sq km.

4.  Islands of the Eastern Atlantic 4.1.  Jan Mayen Dominated in the north by the 2,277 m stratovolcano, Beerenberg, the 54 km long-island of Jan Mayen (373 sq km) (Norway) is steep and rugged, comprising 200–400 m high cliffs eroded everywhere into ash, lava and tephra and in the north (Nord-Jan) into the ice of glaciers that flow from Beerenberg into the sea. The volcano remains active and the coastline has advanced as a result of lava tongues from  eruptions in 1970 and 1985, which have extended into the ocean. Such additions have resulted in pre-­existing  cliffs becoming buried by newer lavas that then become eroded by the sea, such as at Kroksletta, near the northern cape where the recent lava is now cliffed to a height of 5–13 m. Sud-Jan is a hilly cratered ridge of scoria and trachyte domes whose coast is composed of low but  steep rocky cliffs and low gradient but prominent rocky platforms connected by emerged and modern gravelly beaches. Unusually for a small mid-oceanic volcanic island, Nord-Jan is connected to Sud-Jan by a substantial 11 km-long barrier beach that encloses a lagoon on the south-central coast at Lagunevollen (>Fig. 6.5.8), with a smaller barrier and lagoon at Stasjonsbukta on the north coast. Although the north coast is subjected to variable amounts of sea ice for an average of 4 months, the south coast is mainly free of ice. Elsewhere, the rocky coastline of Sud-Jan is not as high as Nord-Jan and small gravelly cliff-foot beaches fed by wave erosion of the adjacent lavas are found.

4.2.  Faeroes The Faeroes (Denmark) (1,398 sq km) (62°N, 7°W) are 16 steep grassy volcanic islands whose landscape is the

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⊡⊡ Fig. 6.5.7 Sable Island, Nova Scotia, is 40 km long and 1.5 km wide and is the exposed part of a sand shoal that extends for more than 160 km. (Courtesy National Aeronautics and Space Administration.)

⊡⊡ Fig. 6.5.8 South coast of Jan Mayen looking north towards the lower slopes of Beerenberg Volcano. (Courtesy Per-Einar Dahlen.)

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result of the erosion of past valley glaciers into a thick pile of horizontally-bedded Tertiary flood basalts. The horizontal nature of the lavas produces a stepped landscape whose interior is boggy. The islands are almost continuously exposed to storm waves, especially from the north and west and so the outer coast is erosional, rugged and dominated by high and plunging cliffs that reach 882 m at Slaettaratindur on Eysturoy. The rate of cliff recession is unknown, but is higher on the seaward facing sea cliffs than the immediately adjacent glaciated cliffs: Tindholmur island is effectively a mountain cut in half, with a vertical seaward cliff and a steep grassy slope on the landward side (>Fig.  6.5.9). Everywhere the coastline is indented with fiord-like inlets, deep clefts and geos and shore platforms are found intermittently. There are few beaches in the Faeroes and these are almost exclusively found at the sheltered fiord heads, where wave approach is unidirectional and small pocket beaches composed of black basalt sand, occasionally backed by low sand dunes are found, such as at Sandur and Saksun.

4.3.  Madeira Madeira (Portuguese) (32° 40' N, 17°W) consists of one large island and together with the smaller island of Porto Santo, the nearby Islas Desertas, and the uninhabited Selvagens to the south, they total 810 sq km and are the emergent summits of extinct (mainly Miocene) volcanoes

much dissected by deep ravines running down to the sea. Terraces mark stages in successive (Quaternary) uplift. The summit of Pico Rivo de Santana rises to 1,861 m. Madeiran slopes are steep, but well vegetated (annual rainfall, c. 600 mm, dry summer) and between the ravines they terminate in high coastal cliffs like Cabo Girao, west of Funchal, at 580 m among Europe’s highest (>Fig. 6.5.10). In the north, fragments of shore-platform are found as well as small islets and skerries, particularly near Faial. There are several well-developed stacks near Ponta do Sao Lourenco in the northeast. Small beaches of rounded grey basalt gravel are found at several places, particularly where river mouths occur. The only sandy beach is at Prainha, on the extreme east of the island where the sand is mainly basalt, but the south coast of nearby Porto Santo is dominated by a long white sand beach consisting of shelly material from the sea floor and aeolian sand blown in from the Sahara Desert. The Selvagens are composed of limestone capped by lava and ash and are cliffed to 100 m in the north, but with a gentler beach-fringed coast in the south. Emerged beaches and dune calcarenites are found on Pleistocene marine terraces at 3 and 7 m above sea level.

4.4.  Canary Islands The Canary Islands (Spanish), between 27° 30' and 29° 30' N and 13° 30'–18° 12' W, comprise seven main volcanic islands (total area, 7,500 sq km), six of which host volcanoes (Teneguia on La Palma, erupted in 1972). From east

⊡⊡ Fig. 6.5.9 The island of Tindholmur in the extreme west of Faeroe is what remains of an interfluve between two ice streams, the inner glacial trough retaining its shape while the outer one is an active marine cliff.

Other Atlantic Ocean Islands

6.5

⊡⊡ Fig. 6.5.10 Cabo Girao, Madeira.

to west they are Lanzarote, Fuerteventura, Gran Canaria, Tenerife, La Gomera, La Palma, and Hierro. Pico de Teide on Tenerife (3,718 m) is the highest peak of these Atlantic islands but they all share a central volcanic spine falling steeply to a mainly rocky and cliff coast cut by deeply incised ravines. The main morphological features of the coastlines of the Canary Islands are related to vulcanicity: layered cliffs of superposed basalt and trachyte lavas, semicircular shapes corresponding to calderas, and long and low craggy volcanic peninsulas that have allowed beach development to occur. The islands have a dry subtropical climate, influenced by southeasterly trade winds from the Sahara: there is a little rain in November–December. Under the hot and dry northeast trade winds the eastern islands of Lanzarote and Fuertoventura receive dust and sand from the Sahara Desert 90 km away, resulting in white sand beaches rather than the black basaltic sand beaches of the other islands. The island slopes show radiating ravines that are short, steep, and gravelly. The north and west coasts are bold, with cliffs up to 300 m high in massive lavas, surrounded by fallen boulders. In the NW of Tenerife at Acantilado de Los Gigantes, the steep coastal slope rises to 500 m. Locally fans of gravelly material swept down ravines have been uplifted and are now cliffed, as on the north coast of Gran Canaria, where there are also emerged shore platforms. There are no biological reefs, but the north coast has ­sectors of shelly beach and calcareous dune sands.

Some of the dunes have been blown inland to form ­desertlike landscapes of bare, mobile dunes on Lanzarote and Fuerteventura. On the southeast coast valley mouth inlets are partly blocked by pebbly beaches, and there are incipient deltas; beaches of black volcanic sand are found locally. On the south coast of Gran Canaria there are beaches of shelly sand, which at Fataga enclose a freshwater lagoon. The tourist beach at San Andreas in SE Tenerife has been artificially nourished with sand shipped from the Sahara Desert. Pared is a sandy tombolo linking the former island of Jandia to mainland Fuertoventura.

4.5.  Cape Verde (Cabo Verde) Islands The Cape Verde Islands (14° 30'–17° 30' N and 22° 40'–25° 30' W), lie about 500 km west of Dakar in Senegal and have a very dry climate because of the northeast trade winds and the desiccating Harmattan winds from the Sahara, which are misty with airborne sand and dust. Mindelo on Sao Vicente has a mean temperature of 22°C in February and 27°C in September, and with a mean annual rainfall of 240 mm (mainly August/September), there are no permanent streams. Barren dissected volcanoes partly overlie Mesozoic marine limestone. There are 15 islands (total area, 4,033 sq km, coastline length 965 km), in two arcuate groups, the southern Sotavento group and the northern Barlavento group. Some of the islands in the

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east are low and undulating, such as Boa Vista, but those in the west are much higher and steeper, and surrounded by craggy plunging cliffs, such as Fogo. Strong winds mobilise sand (volcanic ash) to scour rock outcrops into rugged shapes. Some bays contain narrow fringing algal communities with patchy corals (gardens rather than reefs). Stormy seas have hurled boulders up to 200 m inland and undermined cliffs to produce massive rock falls. Emerged features have been noted at six levels up to 100 m. Four main stratovolcanoes are found in the Cape Verde Islands, Fogo, Brava, Santo Antao and San Vicente but only Fogo is active, its asymmetrical volcanic cone rising 2,829 m above the sea at Pico de Cano. The volcano has erupted ten times between 1500 and 1995. On its steeper east coast several lava flows formed in successive eruptions descend to basal cliffs, below which are small beaches of boulders and gravel. Porto Grande at Mindelo is a deep-water harbour in a volcanic crater, the edge of which has been breached and flooded by the sea. A similar picture characterises the other islands of the group, each being surrounded by steep cliffs cut by deep gullies at the foot of which occur infrequent gravelly beaches.

4.6.  Bioko (Fernando Po) The group of islands in the Gulf of Guinea (3° 30' N, 8° 40' W) include the large Bioko (Equatorial Guinea) (2,017 sq km), a high volcanic island densely covered in tropical rainforest, rising to a peak of 2,850 m at Santa Isabel, with steep rocky coasts interspersed with sheltered sandy bays and mangrove swamps. The capital, Malabo, stands by a bay formed where the sea has invaded a volcanic crater.

4.7.  Principe The 110 sq km island of Principe (Principe and Sao Tome) (1° 37' N,  7° 27' E) has a low, embayed north coast and a higher cliffed south coast. Pico de Principe in the south rises to  1,100 m. The island consists of Tertiary ­volcanic deposits, emergent since the Miocene with shelly ­limestone deposits at 130 m in the Forca Valley. It is dissected

by streams that deliver muddy waters to mangrove-fringed inlets behind barrier beaches, notably at Burras and Lapa.

4.8.  São Tomé Close to the Equator (0° 25' N, 6° 35' E), São Tomé (Principe and Sao Tome) is a large (854 sq km; peak, 2,023 m), densely vegetated island, with a generally subdued volcanic relief having occasional cones and plugs (Cao Grande, 663 m). Beaches of white calcareous sand interrupt high cliffs cut in basalt and phonolite, with columnar jointing well exposed in the Baio do Iologologo. To the south, Ilheu das Rolas is a much smaller volcanic island off the southern coast, and beyond is outlying Annó Bon, last of the group.

References Byrne ML, McCann SB (1995) The dunescapes of Sable Island. Canadian Landform Examples 31. Can Geogr 39:363–368 Guilcher A, Battistini R (1982) Erosional and constructional shore platforms and ancient beaches in the Azores. (in French). Revista Portuguesa de Geografia 15:221–242 Hansom JD, Gordon JE (1998) Antarctic environments and resources: a geographical perspective. Longman, London, 402p Hansom JD (2005) Atlantic Ocean islands, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Kluwer Academic Publications, Dordrecht, the Netherlands, pp 88–95 Hansom JD, Evans DJA, Sanderson DCW, Bingham R, Bentley M (2008) Constraining the age and formation of stone runs in the Falkland Islands using optically stimulated luminescence. Geomorphology 94:117–130 King RL (2001) A glacial origin for Sable Island. Geological Survey of Canada, Current Research D19:18 Mitchell-Thome RC (1970) Geology of the South Atlantic Islands. Beitrage zur Regionalen Geologie der Erde, 10. Gerbrüder Borntraeger, Berlin Mitchell-Thome RC (1976) Geology of the Middle Atlantic Islands. Beitrage zur Regionalen Geologie der Erde, 12. Gerbrüder Borntraeger, Berlin Smith HG, Hardy P, Leith IM, Spaull VW, Twelves EL (1974) A biological survey of St. Paul’s Rocks in the equatorial Atlantic Ocean. Biol J Linn Soc 6:89–96 Vacher L (1973) Coastal dunes of younger Bermuda. In: Coates DR (ed)  Coastal geomorphology. George Allen & Unwin, London, pp 355–391

7.0 British Isles – Editorial Introduction

The coasts of the British Isles are considered severally in the following chapters: England and Wales (> England and Wales Introduction, > Cumbria, > Lancashire, > Isle of Man, > North Wales, > West Wales, > South Wales, > Severn Estuary, > Gloucestershire, Somerset and North Devon, > North Cornwall, > Isles of Scilly, > South Cornwall, > South Devon, > Dorset, > Hampshire, > Isle of Wight, > Sussex, > Kent, > Essex, > Suffolk, > Norfolk, > Lincolnshire, > Yorkshire and Cleveland, > Durham, Tyne and Wear and > Northumberland ), > Scotland, > Northern Ireland, The Republic of > Ireland and the > Channel Islands.

1.  Coastal Geology A distinction was made by Mackinder (1902) between Highland and Lowland Britain, separated by a line drawn from the Exe estuary in South Devon to the Tees estuary on the North Sea coast. Highland Britain consists mainly of mountains and moorlands, dominated by Pre-Cambrian and Palaeozoic geological formations, Lowland Britain of escarpments and vales on Mesozoic, Tertiary and Quaternary formations. The outcrop of Permo-Triassic rocks marks a transition between the two (Stamp 1946). The coasts of Highland Britain (including Ireland and the Channel Islands) are steep and cliffy on the outcrops of older and harder rock formations, while the coasts of Lowland Britain are cliffed, where they intersect the sandstone, limestone and Chalk ridges and low-lying on intervening clay lowlands (Steers 1953). The landscape bears the imprint of Pleistocene glaciations, which at their maximum extended as far south as the Thames valley, South Wales and southern Ireland. Glacial drift deposits resulting from glaciation are widespread within this region. Where they reach the coast they are generally cliffed. In Highland Britain, there are cliffs where pockets of glacial drift occupy valleys, as on parts of the Welsh coast, notably in the northern part of Cardigan Bay, the Lleyn Peninsula, the Isle of Man, Lancashire and Cumbria, and locally in western Scotland. In Lowland Britain, cliffs cut in glacial drift are extensive in East Anglia and Holderness, and are found locally further north on the North Sea coast.

South of the limits of Pleistocene glaciation, there are rubble drifts (Head deposits) produced by intensive frost shattering. These periglacial deposits mantle hillsides and coastal slopes, and are well seen in SW England, where they have been undercut to a varying extent (related to exposure to strong wave action), forming slope-over-wall coast profiles. Similar coastal slopes existed in south-east England, but have been generally destroyed by Holocene marine erosion, and persist only where they have been protected by Holocene deposition, as in the lee of Slapton Sands in South Devon and Chesil Beach in Dorset, and behind alluvial lowlands such as Romney Marsh. Slopeover-wall profiles are also found on parts of the coasts of Highland Britain, as in Cardigan Bay, where periglaciation followed glacial deposition.

2.  Coastal Environments With a prevalence of westerly winds, Highland Britain has higher precipitation (rain and snow) than Lowland Britain, which is comparatively dry. Tide ranges are macrotidal (4–6 m) and megatidal (>6  m) around much of the British coastline, reducing to mesotidal (2–4 m) in NW Wales and W Scotland, Orkney and Shetland and NE Scotland and E Norfolk, and microtidal (<2 m) in the central south (Portland Bill to Christchurch). The Irish coast is mesotidal in the SE, NW and NE, macrotidal in the south and west and microtidal between Coleraine and Cushendal in the north. Wave action is generally strong on the coasts of Highland Britain, particularly on sectors exposed to the Atlantic Ocean in S and SW England, S and W Ireland and W Scotland. Coasts bordering the Irish Sea and the North Sea are exposed to frequent storm waves, and the North Sea is subject to recurrent storm surges that cause coastal flooding and erosion on the east coasts of England and Scotland. In general, erosion is relatively slow on the coasts of Highland Britain and more rapid on the coasts of Lowland Britain, especially on east coast sectors where waves attack soft glacial drift. Sandy beaches are extensive on the Atlantic coasts of Britain, where they include varying proportions of

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calcareous biogenic sand from the sea floor, and are locally shelly. Many are backed by dunes. Shingle beaches are common on the south and east coasts of England, particularly, where they consist of flint gravel eroded from Chalk cliffs or derived from glacial and periglacial drift deposits. Mixed sand and shingle beaches are found along coasts where waves attack cliffs cut in glacial drift, notably in East Anglia and northward. In general, rivers have delivered little sand or gravel to beaches, an exception being in NE Scotland, where the River Spey has contributed gravel of glacifluvial origin to shingle beaches. The evolution of the coasts of Britain has been influenced by changing sea levels, culminating in the Holocene

marine trangsression, which submerged valley mouths to form rias and estuaries. Pleistocene glaciation has been followed by isostatic recovery, with uplift in the ­north-west (NW Scotland and Northern Ireland) and subsidence in the south-east.

References Mackinder HJ (1902) Britain and the British seas. Heinemann, London Stamp LD (1946) Britain’s structure and scenery. Collins New Naturalist, London Steers JA (1953) The sea coast. Collins New Naturalist, London

7.1  England and Wales

Eric Bird

1.  Introduction The Coast of England and Wales is presented with an Introduction followed by 24 illustrated sections comprising the counties of England and three divisions of Wales. In each section the coast is described in a sequence that runs from north to south (> Cumbria to > Cornwall), west to east (> Cornwall to > Kent) and south to north (> Essex to > Northumberland). In England and Wales most of the stages in the geological column are somewhere represented. A notable gap is the Miocene (5–23 million years ago), when Britain was a land area, subject to erosion rather than deposition, apart from some gravel deposits, while some stages, such as the Rhaetic and the Palaeocene, are better developed elsewhere, particularly in Europe. There has been tectonic uplift in the north and west of Britain and subsidence in the south and east. This transverse tilting results in the oldest (Pre-Cambrian) rocks outcropping in parts of Anglesey while the youngest are in East Anglia. The pattern has been complicated by three phases of tectonic activity, the Caledonian, producing structures that run SW-NE in Wales and Cumbria, the Hercynian, producing east-west structures in South Wales and South-West England, and the Alpine, also producing east-west structures across southern England, including the Weald and the Hampshire Basin. The result is that the older rock formations (PreCambrian and Palaeozoic) form the uplands of Wales and the south-west peninsula, and the overlying Mesozoic strata dip generally south-east, the more resistant formations outcropping in roughly parallel escarpments between vales cut in weaker formations in a sequence of diminishing age up to the Chalk. The Tertiary formations are largely confined to the synclinal London and Hampshire Basins in the transversely folded Chalk. The various structures have been truncated by erosion to form the present coastline, and thus expose many different geological formations in cliffs and rocky shores. The influence of Caledonian structures is shown by the elongated SW-NE rock outcrops in Anglesey and the Lleyn

Peninsula in North Wales. The Hercynian structures of the Mendip Hills produce the ridges that protrude into the Bristol Channel at Weston super Mare. The Alpine folding is responsible for the high, narrow Chalk ridge that runs through the Isle of Wight, protruding as cliffy headlands at the eastern and western extremities: Culver Cliff and The Needles. Cambrian rocks outcrop in the Isle of Man, the Lake District and Snowdonia, and Ordovician and Silurian rocks over much of Wales. Devonian and Carboniferous rocks are exposed in Devon and Cornwall, South Wales, and Northumbria. These various formations can be seen in the cliffs and rocky shores of the western coasts of England and Wales. The Permian and Triassic outcrops, which form a low zone between Upland Britain to the west and Lowland Britain to the east, outcrop on the Devon coast between Torquay and Seaton and on the north-east coast between the mouth of the River Tees and South Shields. The Jurassic formations are well exposed in the cliffs of Lyme Bay between Seaton and Portland Bill, in Weymouth Bay and Ringstead Bay, and locally to the east, particularly between Lulworth Cove and Swanage. They are also seen on the north-east coast between Filey Bay and Redcar. The Lower Cretaceous formations have been brought up above sea level in upfolded areas in Purbeck, the Isle of Wight and The Central Weald. In Dorset they outcrop on either side of Lulworth Cove, in Worbarrow Bay and in Swanage Bay, and in Sussex and Kent they extend from Eastbourne to Folkestone. Above them the Chalk, the most distinctive of the Cretaceous formations, outcrops in sectors of cliffed coast at Beer Head in Devon, between White Nothe and Worbarrow Bay in Dorset. The Chalk passes below sea level in the Hampshire Basin, but rises sharply to form a central ridge between The Needles and Culver Cliff in the Isle of Wight. The South Downs and North Downs are in-facing Chalk escarpments overlooking the Weald and reaching the Sussex coast between Brighton and Eastbourne and the Kent coast from west of Dover round to Deal. Chalk cliffs reappear in Thanet extending from Pegwell Bay past Ramsgate and Margate to Minnis

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Bay, west of Birchington. On the North Norfolk coast there is ice-rafted Chalk in the cliffs between East and West Runton, and Chalk outcrops on the shore platform at Sheringham. In Yorkshire the Chalk of the Wolds comes to the coast at Flamborough Head, with Chalk cliffs extending from Sewerby (NE of Bridlington) round to Bempton. The Tertiary formations (Palaeocene, Eocene and Oligocene) overlie the Chalk, and occur in the Hampshire Basin and the Isle of Wight, where they outcrop in cliffs along the northern coast. They are also seen on the north coast of the Isle of Thanet, and in East Anglia, where the cliffs near Walton-on-the-Naze are cut in Eocene formations, while Pliocene formations outcrop in cliffs north of Aldeburgh in Suffolk and on the Norfolk coast between Trimingham, Cromer, Sheringham and Weybourn. Pleis­­ tocene formations are dominated by glacial drift, which has been cliffed on several parts of the east coast of Britain, notably in East Anglia and Holderness, and locally on the west coast bordering the Irish Sea and on the northern peninsula of the Isle of Man. There are also Pleistocene periglacial drift deposits, notably the Head formations prominent on coastal slopes in south-west England and South Wales.

2.  The Coastline The combined coastline of England and Wales (>Fig. 7.1.1) is about 2,734 miles (4,400 km) in length. Its outlines are related to geological outcrops, the more resistant rocks forming headlands while weaker formations have been excavated as bays and inlets (Clayton and Shamoon 1998). Palaeozoic rocks dominate the intricately dissected western peninsulas of Cumbria, Wales and S.W. England, and a sequence of Mesozoic sandstones and limestone cuestas with intervening clay vales ending in alternating cliffs and lowlands on the southern and eastern coasts. Tertiary sedimentary formations outcrop only locally in Dorset, Hampshire and Thanet, but Pleistocene glacial deposits are extensive on the east coast, especially in Holderness and East Anglia. A general account was provided by Steers (1964), and a review of scientific sites by May and Hansom (2003). The cliffs and depositional features formed during the later stages of the Flandrian (Late Pleistocene to Holocene) marine transgression and its aftermath (during the past 5,000 years) of relative stillstand, but relict Pleistocene raised beaches and periglaciated slopes persist where erosion has been slow, mainly on the western coast. Remains of

forests submerged by the Flandrian marine transgression are seen at low tide in the Bristol Channel and ­elsewhere, and continuing subsidence in south-eastern England is indicated by the foundations of Roman London, now 2–3 m below high-tide level. Tide gauge records show that northern England has been rising and the south and south-east sinking at rates of a few millimetres per year. There is much debate on the significance of actual sea level oscillations, of tectonic uplift and subsidence, of upward and downward isostatic movements in response to deglaciation and rising sea levels, of erosion and sedimentation, of land reclamation and coastal engineering as factors influencing ongoing changes in the relative levels of land and sea around England and Wales. The rivers of England and Wales flow into estuaries formed by submergence of valleys that had been deepened during Pleistocene low sea level phases, and fluvial deposition has formed valley floors, but there are no protruding deltas. Embanking and reclamation of intertidal areas has been extensive, especially in Romney Marsh north of Dungeness, the Somerset Levels and the Fens south of The Wash. Marine sand has been washed into several of the drowned valley mouths (rias) of south-west England, producing broad intertidal sandflats, but those that open on sheltered parts of the coast, such as Carrick Roads in Cornwall and Milford Haven in Pembrokeshire, have escaped this accretion, and remain as deep inlets. Wave energy is strong from the southwest through the Atlantic approaches, diminishing into the English Channel and through St. George’s Channel to the Irish Sea. The east coast is sheltered from the prevailing westerly winds, but receives waves from northerly and easterly directions over the North Sea. Refracted patterns of dominant waves have shaped curved bays and beaches, as in Lyme Bay on the Dorset coast (Bird 1989). Longshore drifting of beach material is eastward along the south coast of England and southward down the east coast, except in north Norfolk where westward movement of shingle has contributed to the intermittent growth of compound recurved formations at Blakeney Point and Scolt Head Island. Beach drifting and cliff recession are most pronounced when onshore winds and low barometric pressure bring large waves in at unusually high levels. Major storm surges have occurred on the east coast several times in the past few centuries, the most recent being in 1953 and 1978. The duration and intensity of wave action on the shore is also influenced by tidal regimes. Mean spring tide ranges (>Fig. 7.1.1) are generally between 3 and 5 m, augmented in the Bristol Channel to more than 12 m at Avonmouth, with tidal bores transmitted up the Severn, Avon, and

England and Wales

⊡⊡ Fig. 7.1.1 The coastline of England and Wales. (Courtesy Geostudies.)

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Parrett rivers, and falling to less than 2 m on the central south coast, where double high tides occur. Strong currents are generated in the Irish Sea and through Menai Strait, and tidal currents have shaped complex shoal and channel topography off the east coast, where wave action is relatively weak. Tidal variations also influence the ecological zonations of plant and animal communities on rocky shores (Lewis 1964). Cliffs are found on a variety of geological formations where uplands have been truncated by marine erosion. Much of the coastline of England and Wales runs across the geological strike, so that many formations outcrop in rapid succession along the coast. The most characteristic are the vertical white cliffs of thick gently dipping Cretaceous Chalk on either side of the transected Weald, between Brighton and Eastbourne and between Dover and Deal; however, cliffs in strongly folded chalk occur on the Dorset coast and the Isle of Wight, while at Flamborough Head chalk cliffs are capped by slopes in glacial drift, a sequence that results in slope-over-wall profiles. Steep and high cliffs are also well developed on limestones at either end of the Jurassic outcrop that crosses England from Dorset to Yorkshire, with an outlier in Glamorgan; on the red Permian sediment of St. Bees Head and East Devonshire; and on the Tertiary sands of Bournemouth Bay and the Isle of Wight. Sloping cliffs (<40°) are found on the weaker sands and clays, and slumping occurs on soft Tertiary clays east of Bournemouth and on glacial drift deposits, as in north Norfolk and Holderness. In detail, cliff profiles and cliffy coastlines show the influence of structures such as joints, faults and bedding planes, as well as local variations in rock resistance, notably in the presence of coves, clefts, caves, stacks and natural arches. These features are very well illustrated on the limestone cliffs of South Pembrokeshire. In sectors where the coastline runs parallel to the geological strike, as on the Dorset coast east of Weymouth, a landward dip results in escarpment cliffs, as at Ballard Down, horizontal strata often form stepped profiles, as on Hounstout Cliff, while a seaward dip produces gentler slip-slopes, as at Portland Bill. Where the Chalk dips seaward over clay beds there have been massive landslides (also termed landslips), as be­­ tween Axmouth and Lyme Regis, at Ventnor on the Isle of Wight, and at Folkestone Warren. Coastal landslides were particularly common during the wet winter of 2000–2001. In places, cliff recession has truncated small valleys (chines in the Isle of Wight), and streams emerge in coastal waterfalls. The older and harder rocks of the SW peninsula show cliffs on sectors exposed to strong Atlantic wave action, such as the castellated outcrops of jointed granite at

Land’s End and the intricate patterns of headland and cove, related to faults, thrust planes, and jointing as local variations on the north Cornish coast. More sheltered sectors retain periglaciated slopes formed during the late Pleistocene low sea level and undercut following the Flandrian marine transgression to give slope-over-wall profiles which differ in origin from those related to lithological contrasts, as at Flamborough Head (above). Evencrested cliffs and slopes result from Pleistocene marine planation at higher levels, especially the 130 m level in SW England and Pembrokeshire. Elsewhere, as on the Gower coast in South Wales, cliffs cut back into hilly country increase in height until they intersect an interfluvial ridge, and thereafter diminish. The outlying Isles of Scilly are a largely submerged granite upland with bouldery shores and beaches of white quartzose sand. The north Devon coast, also formed of hard rock formations, faces a relatively sheltered sea area where wave action is impeded by large tide ranges; it is a fine example of a steep coast, with only limited basal cliffing. Across the Bristol Channel the more exposed south coast of Wales shows cliffs of Carboniferous Limestone on the Gower Peninsula and older formations in Pembrokeshire, but in Cardigan Bay exposure to wave attack diminishes northward and sloping coast profiles reappear, becoming prominent on the Lleyn Peninsula and around Anglesey and the Isle of Man. On the coasts of the Irish Sea where glacial drift deposits frequently border the uplands, coastal plains are extensive. Glacial drift fronts bluffs marking a Pleistocene coastline at Ramsey on the Isle of Man, and Walney Island is a glacial deposit. On the east coast the soft glacial drift cliffs of Holderness have been cut back up to 200 m in a century. In north-east England cliffs in Permian and Carboniferous rock include craggy outcrops of the Whin Sill dolerite intrusion which also dominates the outlying Farne Islands. Shore platforms are well developed on Chalk, sandstone and limestone outcrops in southern and eastern England and on rock outcrops of moderate resistance in the south-west, around Wales and locally on the shores of the Irish Sea. Typically they are abrasion platforms (intertidal sloping ramps) formed where waves move shingle to and fro across the foreshore, and they rise to a cliff-base around high tide level, but with considerable variation laterally (Wright 1970). Solution and bioerosion features occur locally on limestones that have escaped abrasion, but the subhorizontal shore platforms developed by weathering processes on microtidal oceanic coasts (as in Australia and New Zealand) are poorly developed in England and Wales: even on the central south coast, where the tide range is small, wave action armed with sand and

England and Wales

pebbles produces abrasion platforms or exposes structural ledges. Planation is less advanced on older and harder formations, which often present a rib-and-corridor topography reflecting lithological variations, while on softer sediment (e.g. Holderness) the crumbling cliffs are fronted by beach-mantled abrasion ramps that decline gradually into a smoothed sea floor. Depositional features include shingle beaches derived from the flint nodules that occur in the Chalk, both directly in the vicinity of eroding cliffs and indirectly by reworking of fluvial, marine, or glacial gravels in post-Cretaceous sediment. Shingle beach trains dominated by flint extend eastward from Lyme Bay along the Channel coast. They include the major barrier at Chesil Beach, spits built by longshore drifting on the Solent shore and deflecting the mouth of the Adur River near Brighton, and the cuspate foreland at Dungeness. Shingle is also present on parts of the east coast at Flamborough Head, on the spit that has regrown periodically at Spurn Head, on Scolt Head Island and Blakeney Point (where it is partly derived from gravelly moraines deposited at the limits of the last glaciation), at Lowestoft, and in the long fluctuating barrier spit of Orfordness. The flint shingle that forms barrier beaches enclosing freshwater lagoons at Porthleven in Cornwall and Slapton in Devon has been swept in from gravel deposits on the sea floor (probably laid down by rivers, including the ThamesRhine-Seine system, during low sea level stages) by wave action during the Flandrian marine transgression, as has Chesil Beach, which is backed by The Fleet lagoon and hinterland slopes that escaped marine cliffing. Beach gravels are also produced by the erosion of Jurassic limestones and Palaeozoic sandstones, volcanic rocks, and vein quartz, but away from the flint-dominated shingle beaches there is usually an admixture of sand derived from the weathering mantles of granitic and other quartzose formations. Some beaches are almost entirely sandy, those of Cornish coves reflecting the mineralogy of weathering rocks on the coast and hinterland stream catchments. Atlantic coast beaches include shelly sand washed in from the sea floor by swell, but on the south side of Barnstaple Bay the massive Westward Ho! cobble beach is being driven intermittently landward and lengthened northward by oceanic storm waves. The spit formations of Cardigan Bay are derived mainly from glacial drift deposits, relics of which persist offshore as partly submerged bouldery sarns, but in the Irish Sea where wave action is weaker, sea-floor shoals and intertidal sand flats remain extensive, particularly in Morecambe Bay, where a current-scoured topography is exposed at low tide. Interacting waves and tides produce variable shoals and

7.1

ebb-flood channels in the Goodwin Sands off Kent and in The Wash off Gibraltar Point (King 1972). Off the East Anglian coast, the changing pattern of shoals has influenced the location and movement of wave-built coastal formations such as Benacre Ness in Suffolk. Sands derived from glacial drift are also extensive along the Northumbrian coast, especially around Holy Island, and on the shores of the Irish Sea. Where the tide range is large, incident waves may shape intertidal sandy shores into ridge and runnel patterns, as on the south Lancashire coast. In general, the beaches of England and Wales consist of shingle and sand derived from the erosion of cliffs and rocky shores, or carried in from sea floor sources during (and to some extent since) the Flandrian marine transgression. Very little beach material is of fluvial origin, most rivers carrying only fine sediment, and sandy deposits in estuaries have generally been swept in from the sea. Sandy beaches at St. Ives Bay and Perranporth in Cornwall and Barnstaple Bay to the north are backed by derived dune formations that have drifted inland and are now largely held in place by grassy vegetation. Elsewhere, successive parallel dune ridges have been built behind sandy beaches, as at Newborough Warren in Anglesey, Studland Bay in Dorset (where they enclose a lake, Little Sea, on the site of a former marine embayment), Winterton Ness in East Anglia and on the south Lancashire coast (Gresswell 1953). Deposition of sand and shingle has narrowed and deflected the mouths of several estuaries, notably the Taw-Torridge and Exe in Devon, and the tidal entrances to Poole, Christchurch, and Pagham harbours on the Channel coast. The pattern of beach drifting in England and Wales has been much influenced by groyne and breakwater construction during the past two centuries. In several localities there is updrift beach accretion alongside breakwaters and downdrift starvation leading to  accelerated erosion: the coast at Newhaven, east of Brighton, is an example. Beach gravels are still being extracted from parts of the Dorset coast (West Bay, Seatown) and in Cornwall (Gunwalloe) despite the risk of accelerating erosion. On the other hand, several Cornish beaches have been augmented by the arrival of sand and gravel derived from mining and quarry waste on the coast, or carried down by rivers from mining areas inland. There have been some artificial beach nourishment projects (e.g., at Bournemouth, Seaford and Minehead), and near Workington on the Cumbrian coast mining waste and industrial slag dumped on the coast are being incorporated in nearby beach sediment. Muddy sediment extensive in east coast estuaries and around The Wash have come partly from hinterland clay

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England and Wales

⊡⊡ Fig. 7.1.2 Coastline changes in England and Wales 1880–1980. (Courtesy Geostudies.)

outcrops by way of suspended silt and clay in river loads, and partly from wave-winnowing of fine-grained material from coastal and sea-floor outcrops of clay, especially glacial boulder clay deposits, for eventual deposition in sheltered environments. Salt marshes dominated by halophytic Salicornia, Suaeda, and Puccinellia spp. have grown alongside estuaries and in areas sheltered by spits, as at Blakeney Point and barrier islands, as at Scolt Head, both

in Norfolk (Steers 1960). Salt marsh vegetation occupies the upper intertidal zone and builds a depositional terrace, which is intersected by branching and meandering tidal creeks and fronted by sandy or muddy areas exposed at  low tide. There is a contrast between the firm sandy marshlands of west coast estuaries and the softer muddy estuaries of the south and east coasts, where occasional storm waves have emplaced shelly cheniers. In 1870

Cumbria

hybridization took place in Southampton Water between local Spartina maritima and American Spartina alterniflora, yielding a vigorous new species, Spartina anglica. This spread rapidly and was introduced to Poole Harbour and other estuaries in England and Wales, and became established on board intertidal mudflats in Bridgwater Bay. It has accelerated the building of depositional terraces, but in some areas it is dying back and the marshlands are eroding. Changes in the outline of England and Wales (>Fig.  7.1.2) were reviewed by a Royal Commission in 1911, which concluded that over a period of about 35 years the gains (14,344 ha) largely by reclamation around estuaries substantially exceeded the losses by erosion (1,899 ha). Historical evidence shows that on the outer coastline only a few sectors (spits, cuspate forelands, accretion alongside breakwaters) have prograded, while on the Chalk (May 1971) and on the softer formations along the  east and south coast rapid cliff recession has consumed  much agricultural land, several settlements (e.g. the fourteenth-­century port of Dunwich, Suffolk) having been lost (Bird and May 1976). It is difficult to find beaches that are actively prograding, and the seaward margins of coastal dunes are commonly cliffed by erosion during recent decades. The response to coastal erosion has been extensive construction of sea walls and boulder ramparts, to­­gether with groynes intended to retain protective beaches, but sea walls and boulder ramparts have often led to beach depletion and shore lowering by wave reflection scour while (as has been noted) groynes impede longshore drifting and result in erosion down-drift. In recent years, coastal management has begun to take account of coastal sediment cells, within which longshore, onshore and offshore movements of sand and gravel occur, and can be assessed in terms of sediment budgets. On this basis, sand and gravel from offshore or hinterland sources can be deposited to augment beaches, and reference has been

7.2

made to such artificial beach nourishment on parts of the coast as an alternative method of coast protection. Predicted changes in climate and sea level are likely to lead to substantial changes on the coast of England and Wales (Cracknell 2003). Low-lying areas, including many that have been embanked and reclaimed during the past few centuries, are likely to be submerged by the sea, and erosion will develop and intensify on cliffy coasts, beaches and dunes, and marshlands. One response may be to extend and elaborate coastal defences, building up lowlying areas by infilling, but another may be managed retreat, by which coastal areas of little economic value will be abandoned to marine incursion.

References Bird ECF (1989) The beaches of Lyme Bay. Proc Dorset Nat Hist Archaeol Soc 111:91–97 Bird ECF, May VJ (1976) Shoreline changes in the British Isles during the past century. Bournemouth College of Technology, Bournemouth, England Clayton KM, Shamoon N (1998) A new approach to the relief of Great Britain, 1. The machine-readable database. Geomorphology 25:31–42 Cracknell BE (2003) Outrageous waves: global warming and coastal change in Britain. Phillimore, Chichester, West Sussex Gresswell RD (1953) Sandy shores in South Lancashire. University of Liverpool, Liverpool King CAM (1972) Beaches and Coasts. St. Martin’s, London Lewis JR (1964) The ecology of rocky shores. English Universities Press, London May VJ (1971) The retreat of chalk cliffs. Geogr J 137:203–206 May VJ, Hansom JD (2003) Coastal geomorphology of Great Britain. Geological Conservation Review Series 28. Joint Nature Conservation Committee, Peterborough, ON, pp 737 Steers JA (ed) (1960) Scolt Head Island. Cambridge University Press, Cambridge Steers JA (1964) The coastline of England and Wales. Cambridge Uni­ versity Press, Cambridge Wright LW (1970) Some characteristics of the shore platforms of the English Channel. Zeitschrift für Geomorphologie 11:36–46

7.2  Cumbria 1.  Introduction The mountains of the Lake District have developed on Lower Palaeozoic slates and volcanic rocks, and have been shaped by Pleistocene glaciation into peaks and ridg­ ­es, separated by the deep troughs which contain the lakes. The Lower Palaeozoic formations have had little influ­­ ence  on the Cumbrian coastline. On the northern side, they pass beneath Carboniferous Limestone and the Coal

Measures, which reach the coast between Maryport and Whitehaven, and on the southern side, they pass beneath Carboniferous Limestone, which forms the hills of the Furness Peninsula, Grange-over-Sands and the Arnside district. There is a fringe of Triassic sandstones and marls which reaches the north coast near Maryport, and fringes Solway Firth is prominent on St Bees Head and underpins the coast down to Barrow-in-Furness and the Isle of Walney, with an outlier on the Flookburgh Peninsula.

353

Cumbria

hybridization took place in Southampton Water between local Spartina maritima and American Spartina alterniflora, yielding a vigorous new species, Spartina anglica. This spread rapidly and was introduced to Poole Harbour and other estuaries in England and Wales, and became established on board intertidal mudflats in Bridgwater Bay. It has accelerated the building of depositional terraces, but in some areas it is dying back and the marshlands are eroding. Changes in the outline of England and Wales (>Fig.  7.1.2) were reviewed by a Royal Commission in 1911, which concluded that over a period of about 35 years the gains (14,344 ha) largely by reclamation around estuaries substantially exceeded the losses by erosion (1,899 ha). Historical evidence shows that on the outer coastline only a few sectors (spits, cuspate forelands, accretion alongside breakwaters) have prograded, while on the Chalk (May 1971) and on the softer formations along the  east and south coast rapid cliff recession has consumed  much agricultural land, several settlements (e.g. the fourteenth-­century port of Dunwich, Suffolk) having been lost (Bird and May 1976). It is difficult to find beaches that are actively prograding, and the seaward margins of coastal dunes are commonly cliffed by erosion during recent decades. The response to coastal erosion has been extensive construction of sea walls and boulder ramparts, to­­gether with groynes intended to retain protective beaches, but sea walls and boulder ramparts have often led to beach depletion and shore lowering by wave reflection scour while (as has been noted) groynes impede longshore drifting and result in erosion down-drift. In recent years, coastal management has begun to take account of coastal sediment cells, within which longshore, onshore and offshore movements of sand and gravel occur, and can be assessed in terms of sediment budgets. On this basis, sand and gravel from offshore or hinterland sources can be deposited to augment beaches, and reference has been

7.2

made to such artificial beach nourishment on parts of the coast as an alternative method of coast protection. Predicted changes in climate and sea level are likely to lead to substantial changes on the coast of England and Wales (Cracknell 2003). Low-lying areas, including many that have been embanked and reclaimed during the past few centuries, are likely to be submerged by the sea, and erosion will develop and intensify on cliffy coasts, beaches and dunes, and marshlands. One response may be to extend and elaborate coastal defences, building up lowlying areas by infilling, but another may be managed retreat, by which coastal areas of little economic value will be abandoned to marine incursion.

References Bird ECF (1989) The beaches of Lyme Bay. Proc Dorset Nat Hist Archaeol Soc 111:91–97 Bird ECF, May VJ (1976) Shoreline changes in the British Isles during the past century. Bournemouth College of Technology, Bournemouth, England Clayton KM, Shamoon N (1998) A new approach to the relief of Great Britain, 1. The machine-readable database. Geomorphology 25:31–42 Cracknell BE (2003) Outrageous waves: global warming and coastal change in Britain. Phillimore, Chichester, West Sussex Gresswell RD (1953) Sandy shores in South Lancashire. University of Liverpool, Liverpool King CAM (1972) Beaches and Coasts. St. Martin’s, London Lewis JR (1964) The ecology of rocky shores. English Universities Press, London May VJ (1971) The retreat of chalk cliffs. Geogr J 137:203–206 May VJ, Hansom JD (2003) Coastal geomorphology of Great Britain. Geological Conservation Review Series 28. Joint Nature Conservation Committee, Peterborough, ON, pp 737 Steers JA (ed) (1960) Scolt Head Island. Cambridge University Press, Cambridge Steers JA (1964) The coastline of England and Wales. Cambridge Uni­ versity Press, Cambridge Wright LW (1970) Some characteristics of the shore platforms of the English Channel. Zeitschrift für Geomorphologie 11:36–46

7.2  Cumbria 1.  Introduction The mountains of the Lake District have developed on Lower Palaeozoic slates and volcanic rocks, and have been shaped by Pleistocene glaciation into peaks and ridg­ ­es, separated by the deep troughs which contain the lakes. The Lower Palaeozoic formations have had little influ­­ ence  on the Cumbrian coastline. On the northern side, they pass beneath Carboniferous Limestone and the Coal

Measures, which reach the coast between Maryport and Whitehaven, and on the southern side, they pass beneath Carboniferous Limestone, which forms the hills of the Furness Peninsula, Grange-over-Sands and the Arnside district. There is a fringe of Triassic sandstones and marls which reaches the north coast near Maryport, and fringes Solway Firth is prominent on St Bees Head and underpins the coast down to Barrow-in-Furness and the Isle of Walney, with an outlier on the Flookburgh Peninsula.

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Cumbria

Glacial drift deposits are extensive on the lowlands around the Lake District, and prominent in low cliffs along parts of the coastline, as on the Isle of Walney (Huddart et al. 1977). Mean spring tide range at Glasson on Solway Firth is 5.2 m, increasing to 8.3 m at Silloth, then declining to 7.7 m at Maryport, 7.5 m at Workington and 7.3 m at Whitehaven. Along the coast south past St Bees the range remains around 7 m, with 7.5 m at Tarn Point and 7.6 m on the Duddon Bar, while the port of Barrow in Furness has 8.2 m.

2.  The Coastline The Cumbrian coast begins at the head of Solway Firth, where the rivers Sark, Esk and Eden flow to a coast with wide salt marshes declining seaward to mudflats exposed at low tide. The Solway salt marshes, like others on the west coast of Britain, are on firm sand or muddy sand, and form lawn-like swards when grazed. Solway Firth widens ­rapidly west­ward, and salt marshes and tidal mudflats are  extensive. There are successive salt marsh terraces (>Fig.  7.2.1), rising landward and related to isostatic uplift, complicated by the effects of migrating river channels and a recent slow rise in relative sea level. The seaward margins of these salt marshes are often cliffed and eroding, despite the emergence. There are scattered scars (boulders washed out of the glacial drift) along the shore. Moricambe Bay has

extensive sandy flats backed by high salt marsh and peat bogs. Numerous small parallel tidal creeks drain the salt marsh, which has cliffed or steep seaward margins exposing layered mud. There are a few Spartina clones on the mudflats. Grune Point is a dune-capped spit of sand and shingle that constricts the entrance to Moricambe Bay. It has been supplied by longshore drifting NE from Silloth, indicated by angled beach segments between groynes. SW from Silloth the sandy intertidal zone widens to an area of ridge and runnel exposed at low tide. The coast is bordered by dunes underlain by glacial drift, from which the shingle and sand beaches, nearshore sand bars and boulder scars have been derived. A 7.6-m emerged beach south of Dubmill Point, has Holocene sand, clay and peat deposits, thought to have formed in a dune-fringed lagoon during the Late Qua­ ternary (Flandrian) marine transgression, and been subsequently raised by postglacial isostatic rebound. In general, emerged beaches decline in level southward along the Cumbrian coast. Allonby Bay has low cliffs cut in boulder clay with occasional outcrops of Triassic St Bees Sandstone. At Maryport, a transverse fault brings up Carboniferous rocks, Millstone Grit and Coal Measures. Near Workington, there are segments of cliff cut into basic slag cliff dumped from a former steelworks at Moss Bay, which have yielded grey sand and gravel and dark brown ironstone to adjacent beaches (>Fig. 7.2.2). South of Workington, the coast was modified by coal mining. The Coal Measures dip beneath sandstones at ⊡⊡ Fig. 7.2.1 A grassy terrace, microcliff and salt marsh at Cardurnock, Solway Firth. (Courtesy Geostudies.)

Cumbria

7.2

⊡⊡ Fig. 7.2.2 Slag heaps at Moss Bay, Workington. (Courtesy Geostudies.)

Whitehaven, and to the south Permian and Triassic sandstones form the cliffed St Bees peninsula, rising to 98 m on North Head (Eastwood et al. 1931). Cliffs in stratified red  sandstone decline to Flenswick Bay at the incised mouth of a small valley. There is a grey pebble beach and the cliffs  have basal ledges of seaward-­dipping Triassic sandstone that show differential erosion along joint planes (>Fig.  7.2.3). The cliffs are locally ­overhanging, and an abrasion notch is prominent along their base. The cliffs then rise towards South Head (>Fig. 7.2.4), where the St Bees Sandstone dips landward and forms an upper cliff, then decline towards the valley of Rottington Beck, which contains glacial drift, and to the south cliffs and bluffs in boulder clay extend for several kilometres behind a grey shingle beach. Three rivers, the Irt, Mite and  Esk, converge into the estuary at Ravenglass, the Irt ­de­­flected southward by the dune-capped spit at Drigg Point. At low tide, the estuary is bordered by gravels  and  sand with mud below mid-tide level. Eskmeals Point to the south, is also dune-covered, and Selker Point is a low eroding cliffed promontory of glacial drift (>Fig. 7.2.5). The River Duddon opens into a large funnel-shaped estuary where at low tide the Duddon Sands are exposed up to 3 km wide and traversed at low tide by the narrow winding river channel and some smaller tidal creeks. On the east coast of the Duddon estuary, salt marshes are backed by steep slopes in Carboniferous Limestone. The marshes pass southward behind a sand dune fringe at Sandscale Haws and Lowsy Point is a dune-covered spit

bordering the shallow northern entrance to the channel behind the Isle of Walney. The Isle of Walney has the form of a barrier island, but it consists of glacial drift. The western part of the island is a series of low hills of glacial drift linked by barriers of sand and gravel, and the northern end is a sandy spit capped by dunes, North End Haws. The Isle of Walney has been frequently overwashed by the sea and coastal erosion has been severe. Parts of the coast have been protected against erosion by sea walls and concrete blocks. The southern part is narrower, with low cliffs that fade as the coast swings towards South End Haws. To the east of the island, an extensive area of intertidal sand and ­mudflats is split by the Walney Channel, which gives ships high tide access to the industrial port of Barrowin-Furness. Morecambe Bay extends across to Heysham in ­Lancashire. The coast runs northward past Piel Island and Roa Island with segments of cliff cut in glacial drift (including truncated drumlins, as at Wadhead Hill), ­fringing salty  mar­­shes and bouldery scars. Beyond Ulverston is ­Greeno­­dd, an old port below the confluence of the Rivers Crake and Leven, which drain from Coniston Water and Lake Windermere respectively, and open into  a  wide estuary with channels and shoals bordered by  salt marsh. Notices warn of quicksands and rapid incoming tides. The hilly country to the east of the estuary is on Carboniferous Limestone, but the undulating  country to the south, around Flookburgh, includes drumlins.

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Cumbria

⊡⊡ Fig. 7.2.3 Flenswick Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.2.4 Shingle beach and cliffs in dipping St Bees Sandstone south of Flenswick Bay, looking towards South Head. (Courtesy Geostudies.)

Humphrey Point stands at the southern end of a prominent Carboniferous Limestone ridge, Humphrey Head, rising to a summit of 53 m. The limestone dips eastward, so that the western side is a high escarpment cliff (>Fig. 7.2.6), while the eastern side is a gentler dip-slope, declining to a  low basal ramp with low wave energy because of the wide salt marsh. Humphrey Point projects

as a rocky spur south-ward into the mudflats of Morecambe Bay (>Fig. 7.2.7). The coastline turns northward, bordered by salt marshes, past Grange-over-Sands. Where there are views across Morecambe Bay, with vast sand flats exposed at low tide. It has tidal channels that change rapidly, deepening, shallowing and migrating. There are sandflats firm enough

Cumbria

⊡⊡ Fig. 7.2.5 Cliff recession at Selker Point has cut into a house. (Courtesy Geostudies.)

⊡⊡ Fig. 7.2.6 Escarpment cliff of Carboniferous Limestone on the western side of Humphrey Head. (Courtesy Geostudies.)

7.2

357

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7.3

Lancashire, Merseyside and Cheshire

⊡⊡ Fig. 7.2.7 Spur of Carboniferous Limestone at Humphrey Point, projecting into the tidal flats of Morecambe Bay. (Courtesy Geostudies.)

to drive a tractor across, but there are also areas of soft mud, particularly where fresh water flows out from rivers. East of the Kent estuary, a limestone ridge steep wood­ ­ed slopes descending to Arnside Point. The Cumbrian coast ends where the Lancashire boundary crosses the Silverdale marshes.

References Eastwood T, Dixon EEL, Hollingworth SE, Smith D (1931) The geology of the Whitehaven and Workington district. Memoirs of the Geological Survey, London Huddart D, Tooley MJ, Carter PA (1977) The coasts of northwest England. In: Kidson C, Tooley MJ (eds) The Quaternary history of the Irish sea. Seel House Press, Liverpool, pp 119–154

7.3  Lancashire, Merseyside and Cheshire 1.  Introdution The Lancashire, Merseyside and Cheshire coast is generally low-lying and dominated by glacial drift deposits, with only local outcrops of the underlying solid formations. It is notable for its wide sandy estuaries, Morecambe Bay, the Lune, the Ribble, the Mersey and the Dee, at the mouths of valleys, which were incised during Pleistocene low sea level phases, submerged by the Late Quaternary (Flandrian) marine transgression, and then partly filled with sediment brought down by the rivers or swept in from the sea floor. Shore and near shore sediment range from shingle and sand to silt and clay, derived from the glacial drift by wave action as the tides (which are substantial) rise and fall. Salt marshes are extensive in the upper intertidal zone, and there has been much embanking and reclamation of these

to secure for farmland or develop ports and industrial areas. Mean spring tide range is generally large, 8.4 m at Morecambe, Heysham and Fleetwood, 7.5 m at St. Anne’s, 8.3 m at Liverpool and 7.8 m at Hilbre Island. In the Dee estuary, the mean spring tide range at Connah’s Quay is 4.6 m.

2.  Lancashire The Lancashire coast begins near the northern end of Silverdale Marsh, which is backed by steep wooded bluffs of Carboniferous Limestone marking a former coastline that extends round to Jenny Brown’s Point. At Silverdale the salt marsh fringe is eroding, and a gravelly beach is forming in front of the eroded cliff (>Fig. 7.3.1).

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Lancashire, Merseyside and Cheshire

⊡⊡ Fig. 7.2.7 Spur of Carboniferous Limestone at Humphrey Point, projecting into the tidal flats of Morecambe Bay. (Courtesy Geostudies.)

to drive a tractor across, but there are also areas of soft mud, particularly where fresh water flows out from rivers. East of the Kent estuary, a limestone ridge steep wood­ ­ed slopes descending to Arnside Point. The Cumbrian coast ends where the Lancashire boundary crosses the Silverdale marshes.

References Eastwood T, Dixon EEL, Hollingworth SE, Smith D (1931) The geology of the Whitehaven and Workington district. Memoirs of the Geological Survey, London Huddart D, Tooley MJ, Carter PA (1977) The coasts of northwest England. In: Kidson C, Tooley MJ (eds) The Quaternary history of the Irish sea. Seel House Press, Liverpool, pp 119–154

7.3  Lancashire, Merseyside and Cheshire 1.  Introdution The Lancashire, Merseyside and Cheshire coast is generally low-lying and dominated by glacial drift deposits, with only local outcrops of the underlying solid formations. It is notable for its wide sandy estuaries, Morecambe Bay, the Lune, the Ribble, the Mersey and the Dee, at the mouths of valleys, which were incised during Pleistocene low sea level phases, submerged by the Late Quaternary (Flandrian) marine transgression, and then partly filled with sediment brought down by the rivers or swept in from the sea floor. Shore and near shore sediment range from shingle and sand to silt and clay, derived from the glacial drift by wave action as the tides (which are substantial) rise and fall. Salt marshes are extensive in the upper intertidal zone, and there has been much embanking and reclamation of these

to secure for farmland or develop ports and industrial areas. Mean spring tide range is generally large, 8.4 m at Morecambe, Heysham and Fleetwood, 7.5 m at St. Anne’s, 8.3 m at Liverpool and 7.8 m at Hilbre Island. In the Dee estuary, the mean spring tide range at Connah’s Quay is 4.6 m.

2.  Lancashire The Lancashire coast begins near the northern end of Silverdale Marsh, which is backed by steep wooded bluffs of Carboniferous Limestone marking a former coastline that extends round to Jenny Brown’s Point. At Silverdale the salt marsh fringe is eroding, and a gravelly beach is forming in front of the eroded cliff (>Fig. 7.3.1).

Lancashire, Merseyside and Cheshire

7.3

⊡⊡ Fig. 7.3.1 The eroding salt marsh at Silverdsle. (Courtesy Geostudies.)

The salt marshes extend along the NE shore of Morecambe Bay, across the Kent estuary and round to Scalestones Point. The mean spring tide range here is over 8 m and the tide rises and falls rapidly across muddy sands that are up to 10 km wide at the lowest tides. The salt marshes have been partially drained and embanked, and are grazed by cattle. Nevertheless, the salt marsh fringe has been eroding. South of Scalestones Point, the eroded coastline has been heavily armoured with large limestone block ramparts, with groynes intended to retain a beach. The beach remains sparse and gravelly, with extensive muddy sand exposed at low tide. At Morecambe, attempts have been made to improve it by dumping sand and shingle in front of the promenade. South of Morecambe, there are low cliffs and a rocky shore in Millstone Grit at Heysham. To the south, the coast is low and marshy with many tidal creeks, and towards Sunderland Point, there is a shingle beach. It becomes a spit on the northern side of the Lune estuary, which is bordered by salt marshes (>Fig. 7.3.2). The River Cocker opens through the marshes to a little gulf at Bank End, and at low tide, the wide Cockerham Sands are exposed. A long embankment begins on the south bank of the River Cocker, enclosing reclaimed land. It is armoured with large limestone boulders along a 3 mile stretch westward from Fluke Hall, where the salt marsh fringe fades out. A concrete sea wall then extends

along the Esplanade at Knott End on Sea, which at low tide is fronted by rough grassy salt marsh. This grades into a wide sandy area known as Preesall Sands and Bernard Wharf, besides the mouth of the River Wyre. West of the River Wyre, the coastline is artificial, with a  high embankment, concrete sea walls and numerous groynes extending past Fleetwood to Cleveleys and Black­ pool. It borders a peninsula, which is of glacial drift underlain by Keuper Marl. The shingle fades out towards Blackpool, where the famous Blackpool Sands are largely submerged at high tide. To the south of Blackpool, the coast swings eastward past Lytham St. Annes, where it passes into salt marshes on the northern shore of the Ribble estuary.

3.  Merseyside The southern shore of the Ribble estuary has salt marshes behind wide intertidal Marshside Sands. The coast be­­ comes sandy towards Southport, where the sands are calcareous, incorporating the residues of shelly marine organisms as well as inorganic sand derived from glacial drift deposits. The coastline has prograded during the past century as the result of sand accretion, forming an outer barrier in front of an elongated lagoon. To the south, Birkdale Sands and Ainsdale Sands are backed by grassy dunes on which famous golf courses have

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Lancashire, Merseyside and Cheshire

been developed (Hansom et al. 1993). The dunes are of calcareous sand winnowed from the wide intertidal sand flats. The wide shore has ridge and runnel topography (Gresswell 1953) with ridges and swales (amplitude up to 1.2 m) parallel to the coastline. There has been ­progradation between Southport and Formby Point, where the coast turns SE, but beyond Formby Point the seaward margin of the coastal dunes has been cliffed by erosion. This erosion has extended northward.

Much of the coastline at Crosby, Bootle and the lower Mersey estuary is artificial, with port, industrial and urban development, and the same is true of the east coast of the Wirral Peninsula. Northward the Mersey cuts through the intertidal sandflats of Liverpool Bay as it opens to the Irish Sea, and the north coast of the Wirral Peninsula is fronted by sandy beaches and wide intertidal sand and mudflats (>Fig. 7.3.3). Red Rocks (Hilbre Point) are a Triassic sandstone outcrop, marking a sharp change to the SW facing ⊡⊡ Fig. 7.3.2 Salt marsh bordering the River Lune at Glasson, looking NW to the power station at Heysham. (Courtesy Geostudies.)

⊡⊡ Fig. 7.3.3 The shore at Hoylake, Wirral Peninsula. (Courtesy Geostudies.)

The Isle of Man

7.4

coastline of the Wirral Peninsula. The shore is sandy, backed by dunes and very extensive intertidal sand and muddy sand areas are exposed at low tide besides the River Dee.

i­ ntertidal zone is predominantly muddy in the south, passing northward to mudflats.

4.  Cheshire

Gresswell RK (1953) Sandy shores of South Lancashire. University of Liverpool, Liverpool, UK Hansom JD, Comber DPM, Fahy FM (1993) Ribble estuary. Coastal Pro­ cesses and Conservation Report, University of Glasgow, Glasgow, UK Marker ME (1967) The Dee estuary: its progressive silting and salt marsh development. Trans Inst Brit Geogr 41:65–71

There are extensive salt marshes on the short Cheshire coast, towards the mouth of the River Dee (Marker 1967) and deposition, and within the wide Dee estuary the

References

7.4  The Isle of Man The Isle of Man trends NE–SW and has an interior upland with valleys radiating from Snaefell (620 m), a transverse vale between Douglas and Peel, and a southern upland rising to South Barrule (477 m). In the south the hill country subsides to a low limestone plateau around Castletown, and in the north is the broad lowland of Ayre, composed of Pleistocene glacial drift deposits. The island is dominated by strongly folded Upper Cambrian Manx (Barrule) Slates, with a Caledonian (NE–SW) trend, and includes small areas of Carboniferous Limestone in the south, some intrusions of basalt and diorite, and Permian sandstones near Peel. Permian rocks occupy subsided basins in the Irish Sea, the Isle of Man being an uplifted horst within the Irish Sea basin. In the Pleistocene the Isle of Man was at times covered by an ice sheet moving south across the Irish Sea basin. The glacial drift in the north of the Isle of Man was part of an extensive area of drift deposits across what is now the floor of the Irish Sea. When the ice finally melted there was an episode of glacial readvance, marked by a push moraine in the Bride Hills. Submerged glacial drift, worked over by wave action during and after the Late Quaternary marine transgression, has yielded a vast quantity of gravel, sand and finer sediment to the shores of the Irish Sea. Late in Pleistocene times the Isle of Man was strongly periglaciated, and the most obvious legacy of this is the slope-over-wall profile that occurs around much of the coast. The slope was formed by periglacial degradation of a Pleistocene cliffed coast, and mantled with solifluction (Head) deposits, while the basal cliff has been cut in Holo­cene times, since the Late Quaternary marine transgression brought the sea to about its present level. Apart from the northern peninsula, which is beachfringed with cliffs cut into glacial drift deposits, much of the Isle of Man coast is rocky, with only limited beaches.

Erosion has been prevalent on much of the coastline, ­especially on the soft glacial drift (Rouse 1990). Features of the coastline are described by Evans (1987). At Douglas in the SE of the island the esplanade is fronted at low tide by a wide beach of sand with some grey sandstone and quartz pebbles above high tide level, widening northward in front of the curved sea wall. To the north a steep slope-over-wall coast extends round to Onchan Head. Several streams flow down to the east coast from the interior uplands, and have incised valleys that open on to the steep coast, such as the deep Groudle Glen opening to Port Groudle. Between Clay Head and Laxey steep grassy coastal slopes on Head deposits descend to shore outcrops of grey Cambrian Manx Slates (>Fig. 7.4.1). The steep coast continues north to Ramsey, where an escarpment of Cambrian slates runs inland across to Ballaugh and Kirkmichael. This was an early Pleistocene coast, cliffed before the glaciers lay down the deposits to the north, and subject to intense periglaciation when the ice edge readvanced to form the Bride push-moraine. North of Ramsey the coast has cliffs cut in glacial drift (>Fig. 7.4.2) extending to the Point of Ayre at the northern end of the Isle of Man. This stands on a heathy plain dotted with gravel pits, which incorporates a raised beach. A broad and high gravel beach of shingle and sand wraps round the Point, and has multiple berms. On the NW coast sand and gravel beaches (>Fig. 7.4.3) backed by dunes with heath vegetation extend south towards Jurby, where they give place to cliffs cut in glacial drift (>Fig. 7.4.4). Near Kirkmichael the cliffs are cut in glacifluvial sands (>Fig. 7.4.5), fronted by a sand and gravel beach with some glacial erratic boulders on the shore. To the south the cliffs of glacial drift give place to cliffs in hard rock formations as the north-facing escarpment reaches the coast.

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The Isle of Man

7.4

coastline of the Wirral Peninsula. The shore is sandy, backed by dunes and very extensive intertidal sand and muddy sand areas are exposed at low tide besides the River Dee.

i­ ntertidal zone is predominantly muddy in the south, passing northward to mudflats.

4.  Cheshire

Gresswell RK (1953) Sandy shores of South Lancashire. University of Liverpool, Liverpool, UK Hansom JD, Comber DPM, Fahy FM (1993) Ribble estuary. Coastal Pro­ cesses and Conservation Report, University of Glasgow, Glasgow, UK Marker ME (1967) The Dee estuary: its progressive silting and salt marsh development. Trans Inst Brit Geogr 41:65–71

There are extensive salt marshes on the short Cheshire coast, towards the mouth of the River Dee (Marker 1967) and deposition, and within the wide Dee estuary the

References

7.4  The Isle of Man The Isle of Man trends NE–SW and has an interior upland with valleys radiating from Snaefell (620 m), a transverse vale between Douglas and Peel, and a southern upland rising to South Barrule (477 m). In the south the hill country subsides to a low limestone plateau around Castletown, and in the north is the broad lowland of Ayre, composed of Pleistocene glacial drift deposits. The island is dominated by strongly folded Upper Cambrian Manx (Barrule) Slates, with a Caledonian (NE–SW) trend, and includes small areas of Carboniferous Limestone in the south, some intrusions of basalt and diorite, and Permian sandstones near Peel. Permian rocks occupy subsided basins in the Irish Sea, the Isle of Man being an uplifted horst within the Irish Sea basin. In the Pleistocene the Isle of Man was at times covered by an ice sheet moving south across the Irish Sea basin. The glacial drift in the north of the Isle of Man was part of an extensive area of drift deposits across what is now the floor of the Irish Sea. When the ice finally melted there was an episode of glacial readvance, marked by a push moraine in the Bride Hills. Submerged glacial drift, worked over by wave action during and after the Late Quaternary marine transgression, has yielded a vast quantity of gravel, sand and finer sediment to the shores of the Irish Sea. Late in Pleistocene times the Isle of Man was strongly periglaciated, and the most obvious legacy of this is the slope-over-wall profile that occurs around much of the coast. The slope was formed by periglacial degradation of a Pleistocene cliffed coast, and mantled with solifluction (Head) deposits, while the basal cliff has been cut in Holo­cene times, since the Late Quaternary marine transgression brought the sea to about its present level. Apart from the northern peninsula, which is beachfringed with cliffs cut into glacial drift deposits, much of the Isle of Man coast is rocky, with only limited beaches.

Erosion has been prevalent on much of the coastline, ­especially on the soft glacial drift (Rouse 1990). Features of the coastline are described by Evans (1987). At Douglas in the SE of the island the esplanade is fronted at low tide by a wide beach of sand with some grey sandstone and quartz pebbles above high tide level, widening northward in front of the curved sea wall. To the north a steep slope-over-wall coast extends round to Onchan Head. Several streams flow down to the east coast from the interior uplands, and have incised valleys that open on to the steep coast, such as the deep Groudle Glen opening to Port Groudle. Between Clay Head and Laxey steep grassy coastal slopes on Head deposits descend to shore outcrops of grey Cambrian Manx Slates (>Fig. 7.4.1). The steep coast continues north to Ramsey, where an escarpment of Cambrian slates runs inland across to Ballaugh and Kirkmichael. This was an early Pleistocene coast, cliffed before the glaciers lay down the deposits to the north, and subject to intense periglaciation when the ice edge readvanced to form the Bride push-moraine. North of Ramsey the coast has cliffs cut in glacial drift (>Fig. 7.4.2) extending to the Point of Ayre at the northern end of the Isle of Man. This stands on a heathy plain dotted with gravel pits, which incorporates a raised beach. A broad and high gravel beach of shingle and sand wraps round the Point, and has multiple berms. On the NW coast sand and gravel beaches (>Fig. 7.4.3) backed by dunes with heath vegetation extend south towards Jurby, where they give place to cliffs cut in glacial drift (>Fig. 7.4.4). Near Kirkmichael the cliffs are cut in glacifluvial sands (>Fig. 7.4.5), fronted by a sand and gravel beach with some glacial erratic boulders on the shore. To the south the cliffs of glacial drift give place to cliffs in hard rock formations as the north-facing escarpment reaches the coast.

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7.4

The Isle of Man

⊡⊡ Fig. 7.4.1 Grassy slopes on Head deposits north of Laxey on the east coast of the Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.2 Cliffs cut in glacial drift north of Ramsey. (Courtesy Geostudies.)

The Isle of Man

⊡⊡ Fig. 7.4.3 Cliffed dunes and zones of sand and shingle on the beach at Lhen, Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.4 Cliffs cut in glacial drift at Jurby Head, Isle of Man. (Courtesy Geostudies.)

7.4

363

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7.4

The Isle of Man

⊡⊡ Fig. 7.4.5 Cliffs in fluvioglacial sands near Kirkmichael, Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.6 Fleshwick Bay. (Courtesy Geostudies.)

The west coast north and south of Peel has steep slopes with basal outcrops of Cambrian rock and bold cliffs on headlands. There are some small bays with beaches of sand and gravel (>Fig. 7.4.6). Port Erin is a wide valley-mouth bay between hogsback cliffs, with quartz-veined phyllites on rocky shores and a sandy beach backed by a sea wall. The steep slopeover-wall coast continues round the southern promontory to Spanish Head. The Calf of Man (>Fig. 7.4.7) is an

­ utlying island separated by Calf Sound, where strong o tidal currents swirl on either side of the rocky iset of Kitterland. Numerous basaltic dykes intersect cliff and shore outcrops, some weathered out as clefts. In the SE of the Isle of Man Carboniferous Limes­­ tone  outcrops on cliffed peninsulas, as at Scarlett Point. Castletown Bay is wide and south-facing, with mainly limestone shores. To the north Santon Head is an outcrop of Carboniferous grits, and the Cambrian Manx Slates come to the steep coast. The coastal slopes are interrupted

The Isle of Man

7.4

⊡⊡ Fig. 7.4.7 View across Calf Sound and Kitterland to The Calf of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.8 Steep coast south of Douglas, Isle of Man. (Courtesy Geostudies.)

at Port Soderick, where a cove stands at the mouth of Glen Soderick, a deeply incised valley. The steep coast on strongly folded Cambrian sandstones and shales continues from Little Ness to Douglas Head (>Fig. 7.4.8), the steep slopes mantled by earthy Head  deposits. There are coves, clefts (zawns) and deep inlets in shales and buttresses on the harder sandstones. Landslides have disrupted the Marine Drive towards Douglas Head, where vertical and steeply dipping slates show breakaways along bedding planes. Landslide debris

has supplied grey shingle to pocket beaches in coves, and there is a natural arch in the cove below the largest slide.

References Evans A (1987) Isle of Man coastal path. Cicerone, Milnthorpe, Cumbria, UK Rouse C (1990) The Isle of Man’s unstable coast. In: Robinson V, McCarroll D (eds) The Isle of Man: celebrating a sense of place. Liverpool University Press, Liverpool, UK

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7.5

North Wales and Anglesey

7.5  North Wales and Anglesey 1.  Introduction The geological formations of North and West Wales show the Caledonian trend (NE–SW), well displayed in the grain of Anglesey and the projection of the Lleyn Peninsula, but complicated by the Harlech (Cambrian) Dome. The geological strike curves westward behind Cardigan Bay and meets the Armorican trend (E–W) in Pembrokeshire and extending through South Wales. The interior includes several planation surfaces, shaped by erosion in relation to higher sea levels in Tertiary and Quaternary times (Brown 1960). There is an early Pleistocene coastline at about 180 m, and a series of marine terraces at lower ­levels, notably between 45 and 75 m, prominent in Pembrokeshire. Glacial drift deposits are extensive, and in coastal regions where they have been cliffed by marine erosion they have yielded sand and gravel to beaches and tidal sand flats and finer sediment to salt marshes and tidal mudflats. Beaches have also received sand and gravel swept in from the sea floor by wave action, but fluvial sediment yields have generally been too fine for retention on beaches. Much of the coast is cliffed, with interruptions at the mouths of valleys that were submerged by the Late Quaternary marine transgression and have been to varying degrees infilled with sediment brought down by rivers and swept in from the sea.

2.  North Wales and Anglesey The north coast of Wales has been much influenced by  the  deposition of Pleistocene glacial drift deposits, which largely conceal the solid rock formations that outcrop in the hinterland. Carboniferous Limestone comes to the coast west of Abergele, and is succeeded by Lower Palaeozoic volcanic rocks west of Conway. Anglesey has parallel outcrops of Pre-Cambrian, Ordovician and Car­ boniferous formations, trending NE–SW, as in Menai Strait which separates it from the mainland. Mean spring tide range on the north coast of Wales is generally between 6 and 7 m (Llandudno 6.9 m, Trwyn Du 6.9 m, Menai Bridge 6.6 m and Amlwch 6.4 m), but diminishes round Anglesey (Wylfa Head 5.7 m, Holyhead 4.9 m, Porth Rhuffydd 4.3 m, Llanddwyn Island 3.9 m). The Point of Air in NE Wales is a foreland of grassy dunes over shingle. To the west a wide alluvial coastal plain is backed by steep hills at the northern end of the Clwydian Range. Wide sand flats are exposed at low tide. Between Prestatyn and Rhyl erosion has long been a

­ roblem, and much of the coastline is armoured by sea p walls and boulder ramparts. Longshore drifting is predominantly eastward along the coast, and a spit that has grown east from Kinmel Bay deflecting the mouth of the River Clwyd. Towards Abergele the sandy shore is backed by a narrow coastal dune fringe. The coastal plain narrows westward in front of a steep slope cut in hard Silurian and Lower Carboniferous rocks, which comes to the coast at Colwyn Bay. The coast then curves out to a promontory of  Carboniferous Limestone at Rhos-onSea. To the west the land rises to Little Ormes Head, also Carboniferous Limestone, with a 141 m summit and much bare rock. Llandudno has a promenade with a high wide beach of grey shingle declining westward to sand as the coast swings out to limestone cliffs of Great Ormes Head. This is a massive, cliff-edged block of Carboniferous Limestone, rising to a summit of 207 m and linked to mainland by an isthmus of dune sand over glacial drift. It is an elongated synclinal structure with an axis SW–NE. The northern coast is an escarpment on dipping limestone (>Fig. 7.5.1). To the south is Conwy Bay, which at low tide has a  wide expanse of sandflats and mudflats with bars and  tidal channels, and sandy beaches on either side of  the  Conwy estuary. West of Conwy sharply rising slopes on Ordovi­cian shales and grits run out to a steep promontory at Penmaen-bach Point. At Penmaenmawr (>Fig.  7.5.2) a steep slope of rhyolite rises behind a coastal fringe of glacial drift, and has been extensively quarried. There has been slope stabilisation above the coast road and railway here, with contour fencing to intercept falling rocks. Much of the coastline is artificial, with defensive walls. The intertidal zone widens rapidly to the extensive Lavan Sands, an area of sand and mud (reworked glacial drift). Seaweed-covered boulders, Llys Helig, extend a mile offshore, uncovered at low tide, the remains of a glacial moraine from which the finer material was swept away by the sea during the later stages of the Late Quaternary marine transgression. Menai Strait is 20 km long, and was formed by the submergence of valleys that diverged from a divide that is now submerged. It is a drowned lowland following the NE–SW geological strike, and is now sluiced by strong tidal currents. On the northern side of the strait a steep wooded coast runs eastward past the foreland of Gallows Point and along

North Wales and Anglesey

7.5

⊡⊡ Fig. 7.5.1 Eastward-dipping Carboniferous Limestone at St Tudwal’s Buttresses on Great Ormes Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.5.2 The steep rhyolite slope west of Penmaenmawr. (Courtesy Geostudies.)

to Beaumaris. Towards Beaumaris the beach becomes sandy, and drifting sand has accreted alongside the breakwater. A succession of bays and minor headlands with cliffs cut into drumlins and fluvioglacial deposits then extends north to Penmon, and off Black Point is Puffin Island, a steep-sided grassy island of Carboniferous Limestone. West from Black Point a steep coast in Carboniferous Limestone extends round to Red Wharf Bay, a 4 km wide bay with extensive sand flats exposed at low tide. It stands at the north-eastern end of the lowland corridor that runs

across the island from Malltraeth Bay. Bluffs and cliffs run round to Benllech. The coast to Moelfre is cliffed and rocky, with a sandy cove at Traeth Bychan. Moelfre stands on a cliffed promontory and cliffs then pass behind long sandy beaches, Traeth Lligwy and Traeth-yr-Ora. The north coast of Anglesey has slope-over-wall cliffs in hard Pre-Cambrian rock (>Fig. 7.5.3), dissected into coves and clefts, headlands and stacks, in relation to variations in resistance of component formations.

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North Wales and Anglesey

⊡⊡ Fig. 7.5.3 Cliffs in Pre-Cambrian gneiss at St. Patrick’s church, Llanbadrig, Anglesey. (Courtesy Geostudies.)

⊡⊡ Fig. 7.5.4 The White Lady, a stack on the east coast of Cemmaes Bay, Anglesey. (Courtesy Geostudies.)

Cemmaes Bay has cliffs cut into Ordovician shales, an  exhumed sub-drift landscape. Two glacial drift-filled ­valleys enter the bay. On the southern coast is White Lady  Beach, a bay with a tall white stack in the middle (>Fig. 7.5.4) and a beach of well-rounded pebbles. At Cemmaes Head the cliffed coast swings south past a series of small coves, declining behind Church Bay, a sandy cove. Sandy beaches alternate with rocky and shore platforms. Holyhead Bay has wide intertidal sand flats on either

side of a channel that leads out from a lagoon. On its western coast a drumlin has been truncated by a marine cliff. Holy Island is separated from Anglesey by an irregular strait running south from Holyhead Bay to a narrow outlet to Cymyran Bay. Inland slopes rise to Holyhead Mountain (220 m). North Stack is a tall cliffy islet of grey strongly-folded quartzite, and rugged cliffs extend behind to South Stack, also of quartzite, linked to the mainland by a suspension bridge.

West Wales

Bold cliffs continue along the south coast of Holy Island to Cymyran Bay, which has a 3 km long sandy beach, wide at low tide, in front of low cliffs and grassy dunes, extending to the rocky reefs at Rhosneigr. Head­ lands and bays alternate to the SE at the ends of ridges and valleys down to Malltraeth Bay, which has 5 km of  sand between the low spring tide line and a high grassy embankment that keeps the sea out of Malltraeth Marsh. Llanddwyn Island, a long rocky peninsula of PreCambrian volcanic rocks, rises to 12 m and is notable for  exposures of pillow lavas in cliff and foreshore outcrops. It  projects southward from a sandy beach and a coast fringed by hummocky dunes. To the east the dunes form Newborough Warren, where three successive ridges of dune sand have migrated inland, with associated

7.6

v­ egetation zones (Ranwell 1958, 1972). On the southern shore is a sand and shingle beach, and eastward drifting has formed a spit with several recurves at Abermenai Point, bordering  the southern entrance to Menai Strait (Robinson 1980).

References Brown EH (1960) The relief and drainage of Wales. University of Wales Press, Cardiff, UK Ranwell DS (1958) Movement of vegetated sand dunes at Newborough Warren, Anglesey. J Ecol 46:83–100 Ranwell DS (1972) Ecology of salt marshes and sand dunes. Chapman & Hall, London Robinson AHW (1980) The sandy coast of south-west Anglesey. Trans. Anglesey Antiq. Field Club for 1980:37–66

7.6  West Wales The NE–SW Caledonian trend is evident in the PreCambrian, Silurian and Ordovician rocks of the Lleyn Peninsula but to the east the Cambrian formations come to the surface in the Harlech Dome, the coastline between Criccieth and Barmouth skirting its western edge. To the south, between Tywyn and Fishguard, the Caledonian trend brings Lower Silurian and Upper Ordovician formations to the coast of Cardigan Bay. West of Fishguard the coast to St David’s Head has Lower Ordovician formations, with numerous igneous intrusions forming headlands between valleys and bays cut in the softer Ordovician rocks. The coasts of Cardigan Bay are dominated by bevelled cliffs, typically with straight slopes descending to a steep or vertical rock wall. There is often a basal terrace of boulder clay, the remains of an extensive boulder clay apron that has been cut back by the sea to form the slope-overwall profiles. Irish Sea ice extended along the coast in late Pleistocene times, impounding proglacial lakes in valleys with the result that overflow channels were cut locally. Mean spring tide ranges are generally about 4 m. Porth Dinnle and Bardsey Island have 3.7 m, Pwllelhi 4.3 m, Porthmadog and Barmouth 4.4 m, Aberdovey and Aberystwyth 4.3 m, New Quay and Port Cardigan 4.1 m and Fishguard 3.9 m. A shingle spit borders the low-lying Morfa Dinnle, south of the entrance to Menai Strait, and Dinas Dinnle is a coastal hill where an Iron Age camp has been truncated by a cliff cut in glacial drift deposits. South from Dinas Dinlle the cliffs are cut in boulder clay at Llandwrog

(>Fig. 7.6.1) and rise past Pontllyfni behind a bouldery shore with some sand and gravel beaches. South of Trefor the coast steepens as hinterland slumping slopes rise to Yr Eifl, and there are coastal quarries in the igneous rocks. Quarry waste has spilled down the very steep slope to the rocky shore (>Fig. 7.6.2). Beyond the rocky headland at Penrhyn Glas an apron terrace of glacial drift descends to cliffs with basal rock outcrops at Porth Pistyll. The Trwyn Porth promontory is of Pre-Cambrian pillow lavas. At Porth Nefin glacial drift occupies a deep hollow excavated by the sea, and a wide sandy beach fronts cliffs and bluffs incised by narrow steep-sided chines (>Fig. 7.6.3). The cliffs and bluffs continue SW past Rhos-y-Llan, punctuated by small coves with sandy beaches and rock pools exposed at low tide. At Porth Colmon (>Fig. 7.6.4) a Pleistocene 6 m pebbly raised beach is well exposed, overlain by glacial drift deposits (Whittow 1960). At Porth Oer cliffs cut in glacial drift interrupt the PreCambrian rocky coast, and there is a long beach with the Whistling Sands, which squeak when walked upon, backing bluffs in glacial drift and a rocky foreshore with pools exposed at low tide. SW of Porth Oer the cliffs are cut into the slopes descending from Mynydd Anelog, and then the lower slopes of Mynydd Mawr. Bardsey Island is of Pre-Cambrian rocks, and intervening Bardsey Sound is a tidal strait scoured to 40 m by strong tidal currents. At the southern end of the Lleyn Peninsula the cliffs turn northward into Aberdaron Bay along a steep and cliffed coast. This bay has been cut out

369

West Wales

Bold cliffs continue along the south coast of Holy Island to Cymyran Bay, which has a 3 km long sandy beach, wide at low tide, in front of low cliffs and grassy dunes, extending to the rocky reefs at Rhosneigr. Head­ lands and bays alternate to the SE at the ends of ridges and valleys down to Malltraeth Bay, which has 5 km of  sand between the low spring tide line and a high grassy embankment that keeps the sea out of Malltraeth Marsh. Llanddwyn Island, a long rocky peninsula of PreCambrian volcanic rocks, rises to 12 m and is notable for  exposures of pillow lavas in cliff and foreshore outcrops. It  projects southward from a sandy beach and a coast fringed by hummocky dunes. To the east the dunes form Newborough Warren, where three successive ridges of dune sand have migrated inland, with associated

7.6

v­ egetation zones (Ranwell 1958, 1972). On the southern shore is a sand and shingle beach, and eastward drifting has formed a spit with several recurves at Abermenai Point, bordering  the southern entrance to Menai Strait (Robinson 1980).

References Brown EH (1960) The relief and drainage of Wales. University of Wales Press, Cardiff, UK Ranwell DS (1958) Movement of vegetated sand dunes at Newborough Warren, Anglesey. J Ecol 46:83–100 Ranwell DS (1972) Ecology of salt marshes and sand dunes. Chapman & Hall, London Robinson AHW (1980) The sandy coast of south-west Anglesey. Trans. Anglesey Antiq. Field Club for 1980:37–66

7.6  West Wales The NE–SW Caledonian trend is evident in the PreCambrian, Silurian and Ordovician rocks of the Lleyn Peninsula but to the east the Cambrian formations come to the surface in the Harlech Dome, the coastline between Criccieth and Barmouth skirting its western edge. To the south, between Tywyn and Fishguard, the Caledonian trend brings Lower Silurian and Upper Ordovician formations to the coast of Cardigan Bay. West of Fishguard the coast to St David’s Head has Lower Ordovician formations, with numerous igneous intrusions forming headlands between valleys and bays cut in the softer Ordovician rocks. The coasts of Cardigan Bay are dominated by bevelled cliffs, typically with straight slopes descending to a steep or vertical rock wall. There is often a basal terrace of boulder clay, the remains of an extensive boulder clay apron that has been cut back by the sea to form the slope-overwall profiles. Irish Sea ice extended along the coast in late Pleistocene times, impounding proglacial lakes in valleys with the result that overflow channels were cut locally. Mean spring tide ranges are generally about 4 m. Porth Dinnle and Bardsey Island have 3.7 m, Pwllelhi 4.3 m, Porthmadog and Barmouth 4.4 m, Aberdovey and Aberystwyth 4.3 m, New Quay and Port Cardigan 4.1 m and Fishguard 3.9 m. A shingle spit borders the low-lying Morfa Dinnle, south of the entrance to Menai Strait, and Dinas Dinnle is a coastal hill where an Iron Age camp has been truncated by a cliff cut in glacial drift deposits. South from Dinas Dinlle the cliffs are cut in boulder clay at Llandwrog

(>Fig. 7.6.1) and rise past Pontllyfni behind a bouldery shore with some sand and gravel beaches. South of Trefor the coast steepens as hinterland slumping slopes rise to Yr Eifl, and there are coastal quarries in the igneous rocks. Quarry waste has spilled down the very steep slope to the rocky shore (>Fig. 7.6.2). Beyond the rocky headland at Penrhyn Glas an apron terrace of glacial drift descends to cliffs with basal rock outcrops at Porth Pistyll. The Trwyn Porth promontory is of Pre-Cambrian pillow lavas. At Porth Nefin glacial drift occupies a deep hollow excavated by the sea, and a wide sandy beach fronts cliffs and bluffs incised by narrow steep-sided chines (>Fig. 7.6.3). The cliffs and bluffs continue SW past Rhos-y-Llan, punctuated by small coves with sandy beaches and rock pools exposed at low tide. At Porth Colmon (>Fig. 7.6.4) a Pleistocene 6 m pebbly raised beach is well exposed, overlain by glacial drift deposits (Whittow 1960). At Porth Oer cliffs cut in glacial drift interrupt the PreCambrian rocky coast, and there is a long beach with the Whistling Sands, which squeak when walked upon, backing bluffs in glacial drift and a rocky foreshore with pools exposed at low tide. SW of Porth Oer the cliffs are cut into the slopes descending from Mynydd Anelog, and then the lower slopes of Mynydd Mawr. Bardsey Island is of Pre-Cambrian rocks, and intervening Bardsey Sound is a tidal strait scoured to 40 m by strong tidal currents. At the southern end of the Lleyn Peninsula the cliffs turn northward into Aberdaron Bay along a steep and cliffed coast. This bay has been cut out

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⊡⊡ Fig. 7.6.1 Boulder clay cliffs at Llandwrog, looking south towards the Yr Eifl mountains. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.2 Quarry waste spilling down coastal slope, Yr Eifl. (Courtesy Geostudies.)

between cliffy promontories of Pre-Cambrian rock to the west and Ordovician to the east, but at the head of the bay the south-facing cliffs are cut in soft boulder clay. On the eastern side of Aberdaron Bay the cliffs run out to a headland at Trwyn-y-Penrhyn, and there is a 6 m raised beach around offshore islands. A cliffed promontory runs past Porth Ysgo, where the coastal slope has been disturbed by former mining, and round beside Porth Neigwl, a wide bay with receding cliffs cut in soft boulder clay along a gently curved coastline, shaped by

refracted south-westerly waves. East from Abersoch is a  succession of asymmetrical embayments with long beaches of sand and gravel between rocky headlands. Extensive intertidal sands at the head of Tremadog Bay, passing into the Glaslyn estuary, are the outcome of northward drifting of beach material from Cardigan Bay and eastward drifting from Lleyn, together with shoreward drifting from the sea floor. The long sandy estuaries of the Glaslyn and the Dwyryd have been much modified by reclamation.

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⊡⊡ Fig. 7.6.3 Porth Nefin. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.4 Emerged beach at Porth Colmon. (Courtesy Geostudies.)

Morfa Harlech is a broad triangular foreland south of the Glaslyn estuary, consisting of an undulating lowland of glacial drift and alluvium, with hollows occupied by salt marshes and freshwater swamps. The western shore is backed by shingle and sand beaches backed by dunes. The cuspate foreland at Morfa Dyffryn is similar to Morfa Harlech in that it consists of glacial drift with a fringe of beaches and dunes, particularly on the SW flank. Both forelands have grown northward as the result of longshore drifting and accretion, but their SW coasts are eroding.

South of Llanaber cliffs and shingle beaches extend to Barmouth, which stands on a small sandy foreland at the mouth of the Mawddach estuary. This estuary occupies a valley cut into Cambrian beds on the southern side of the Harlech Dome, and has been formed by Late Quaternary marine submergence. Ro Wen is a shingle and sand spit that runs north from Fairbourne and almost blocks the mouth of the estuary. South of Fairbourne a steep coast with basal cliffs and a boulder beach is interrupted by a delta-like lobe of glacial drift at

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the mouth of Afon Gwrill. The coastal plain extends southwards past Tywyn. An alluvial coastal plain continues southward, bordered by dunes and a beach of shingle and sand which curves eastward to Aberdovey. The Dyfi estuary is cut in strongly-folded slates, grits and sandstones which strike NNE–SSW. It is bordered by salt marshes with grazed turf, cliffed edges seaward and undercut creek banks (Burd 1989). On the southern side the Borth sand and shingle spit has grown about 3 km northward from Upper Borth, supplied with sand and shingle drifting from the south and shaped by westerly waves. At Upper Borth (>Fig. 7.6.5) a storm beach of grey platy cobbles is backed by a sea wall, and fronted by a wide sandy surf beach with remnants of a submerged forest exposed at low tide, with peat deposits are over estuarine clay. The peat formed about 6,000 years ago, and is exposed because of eastward migration of the spit. The beach ends as cliffs rise behind a shore platform. From Upper Borth to Aberystwyth and New Quay the Silurian Aberystwyth Grits form cliffs and shore platforms, locally mantled by boulder clay and interrupted by valleys floored with glacial drift. Where the dip is landward the cliffs slope seaward, where it is horizontal they are more or less vertical, and where it is seaward there is landslide topography. At the mouth of the Wallog valley there is a mantle of boulder clay, and a gravelly cause­way, Sarn Cynfelyn, well exposed at low tide, runs out into Cardigan Bay.

The cliffs come to an end at Aberystwyth, where a promenade runs behind a beach of shingle and a sandy foreshore exposed at low tide. To the south the Afon Rheidol comes in from the east to meet the Afon Ystwyth on the seaward side of high Pendinas Hill. The Ystw­­ yth flows in from the south across an alluvial plain behind a shingle barrier beach, then the coast rises to cliffs at Peny-bwlch, truncating a coastal ridge. The shore is littered with boulders, apart from a sector of abrasion platform cut across the steeply dipping mudstones (>Fig. 7.6.6). The cliffs to the south are steep and up to 30 m high. They are interrupted by valleys, and in places fronted by a boulder clay apron. A deep, narrow valley opens to the coast at Aberarth, where a 4 m cliff stands behind a bouldery cobble beach and the river mouth has been diverted northward by a small spit. Between Aberarth and Aberaeron there are cliffs cut in boulder clay. At Aberaeron beach shingle drifts northward, and by the early nineteenth century it had deflected the river mouth of the Afon Aeron about 650 m northward, but in 1807 it was diverted back to the harbour. To the south the shingle beach narrows in front of rising cliffs. New Quay stands in a bay with slumping cliffs in soft grey shale, and at New Quay Head the cliffed coastline turns WSW, rising along the flank of a coastal ridge. Another deep narrow valley descends to Cwm Tydi, a cove with a sea wall behind a beach of sand and platy blue– grey cobble gravel, and to the SW is an intricate coastline with coves and small headlands, including the narrow

⊡⊡ Fig. 7.6.5 The upper beach of shingle behind a sandy foreshore at Upper Borth. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.6.6 Shore platform on seaward dipping mudstones at Pen-yBwlch. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.7 Coastal slope west of Cwm Tydi. (Courtesy Geostudies.)

irregular Ynys Lochtyn peninsula (>Fig. 7.6.7) with a 30 m high coastal slope descending to cliffs and clefts. The steep coast continues westward behind Traeth Penbryn, a long beach of yellow sand with some gra­­ vel,  backed by low dunes and cliffs cut in boulderclay. Aberporth is in a double cove where two valleys ­converge, then the coast rises to the high headland above Cribach Bay. The coastal plateau is trenched by ­deeply-incised V-shaped valleys that may have been meltwater channels cut by overflow from an impounded glacial lake.

To the south Traeth-y-Mwnt (>Fig. 7.6.8) is a cove bordered by cliffs cut in steeply-dipping dark grey sedimentary rocks and backed by a 20 m high cliff cut in ­glacial drift and capped by grassy dunes. There is a beach of ­yellow–brown sand, wide at low tide. South from Mwnt the cliffs are rugged, with coves, caves and natural arches. They have slope-over-wall profiles, which also extend round Cardigan Island offshore. A funnel-shaped estuary at the mouth of the River Teifi opens between a broad sandy lobate foreland on the

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⊡⊡ Fig. 7.6.8 Traeth-y-Mwnt. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.9 Dinas Head. (Courtesy Geostudies.)

s­ outhern shore, the beach backed by grassy dunes, and a dune-capped spit at Gwbert on the northern shore. At low tide the waves break around the outer edge of Poppit Sands, a wide intertidal sand flat with bands of paler yellow sand marking multiple strandlines. Salt marshes have been dissected by tidal creeks and cliffed at the outer edge by wave erosion. Cemmaes Head is backed by overflow channel from a glacial lake that occupied the Teifi valley when ice blocked its mouth. Slope-over-wall cliffs continue south–west to Newport Bay, where Newport Sands are extensive at low

tide, the Afon Nyfer winding across them to the sea between sand spits. Inland the ground rises towards Presceli Mountain. The coast swings out to the high promontory of Dinas Head, rising 142 m above sea level (>Fig. 7.6.9), bordered on its landward side by a valley cut by meltwater overflow from a glacial lake. West from Dinas Head cliffs in dark Ordovician shales and sandstones are broken by deep and narrow valleys. Towards Fishguard the cliffs are on paler igneous rocks, on an indented coast, and the large breakwater at

South Wales

Fishguard Harbour lies under a cliff on the western side of the bay. Strumble Head is a bold heathy promontory with cliffs exposing pillow lavas, rhyolites, agglomerates and volcanic ash. To the south are coves cut in Ordovician shales between steep bluffs, cliffs and rocky shores on volcanic and intruded rocks. There is an extensive capping of boulder clay, and flint gravel is found on beaches. Abercastle is behind a long narrow drowned valley mouth. The sequence of headlands in dolerite or rhyolite and bays and inlets in softer Ordovician sedimentary rocks continues SW to St. David’s Head. As most of the

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rivers on the St. Davids Peninsula rise close to the north coast and flow south there are few sandy coves on the cliffy northern shore.

References Burd F (1989) The saltmarsh survey of Great Britain. Research and Survey in Nature Conservation No. 17, Peterborough Whittow JB (1960) Some comments on the raised beach platform of south-west Carnarvonshire and on an unrecorded raised beach at Porth Neigwl. Proc Geologis Assoc 71:31–39

7.7  South Wales Along the south coast of Wales the geological formations trend generally west to east, and the coastline intersects outcrops of rock formations that are often steeply dipp­­ ing.  The Pre-Cambrian and Lower Palaeozoic rocks of St. David’s Peninsula are succeeded southward in St. Bride’s Bay by Devonian (Old Red Sandstone) and Carboniferous formations, including Carboniferous Limestone, Millstone Grit and Coal Measures. These extend W–E across south Pembrokeshire, Carmarthen Bay and the Gower Peninsula to Swansea Bay and Porthcawl. There are broad plateaux ending in even-crested cliffs, except where incised valleys open to the coast. In Glamorgan the Carboniferous rocks subside beneath Trias and Lias formations. Keuper Marl forms a coastal fringe north–east from Cardiff, but is largely concealed by Pleistocene glacial drift and Holocene alluvium in the Wentlooge and Caldicot Levels. Mean spring tide ranges increase W–E along the coast. On Skomer Island the range is 5.9 m and at Milford Haven 6.3 m. Stackpole Quay has 7.0 m and Tenby 7.6 m. On Gower the range at Whitford Lighthouse is 7.8 m and in Swansea Bay it increases from 8.6 m at Swansea to 8.7 m at Port Talbot and 8.9 m at Porthcaw l8.9. The increase ­continues as the Bristol Channel narrows into the Severn estuary: Barry has 10.7 m, Cardiff 11.1 m, Newport 11.9 m, Sudbrook 12.4 m and Beachley, at the Severn Bridge, 12.4 m. In Pleistocene times the Devensian Glaciation is thought to have extended to South Wales, the limit (c 20,000 bp) of Irish Sea ice extending from the southern side of St. Brides Bay inland over north Pembrokeshire to meet Welsh Ice that extended back out across Carmarthen Bay and the northern part of the Gower Peninsula, across Swansea Bay to Porthcawl, then inland behind South Glamorgan to a lobe between Cardiff and Newport.

On the Pembrokeshire coast (John 1979) St. David’s Head is a promontory of gabbro. Cliffs cut in Cambrian rocks fringe the coast into Whitesands Bay, which has a typical west-facing Atlantic surf beach, wide at low tide and backed by dunes. Clefts and caves have been etched out along faults in the Cambrian rocks along the shore, and there is an emerged shore platform with a raised beach of sand and shingle. The St. David’s Peninsula is ­dominated by a broad plateau 45–75 m above sea level, generally ­considered to be a marine planation surface with residual higher hills that were once islands when the sea stood at a higher level. The Bitches is a tide race inshore from Ramsey Island, with strong currents as the tide rises and falls. Ramsey Island is a detached slab of the coastal plateau bordered by  cliffs in Ordovician shales between headlands of quartz-porphyry. The mainland coast swings eastward with slope-over-wall cliffs in Pre-Cambrian volcanic rocks, with intruded granites and dolerites forming headlands and islands. East of Pen-y-Cyrfrwy a coastal fringe of Cambrian rocks outcrops in the cliffs, and these continue past Caerfai Bay to the Solva inlet, and on to the mouth of the Cwm-mawr valley, where a major fault marks the beginning of the strongly folded Carboniferous formations (Coal Measures) which dominate the east coast of St. Bride’s Bay. At Newgale the sandy beach, wide at low tide, is backed by a straight ridge of storm-piled grey and brown cobbles. The cliffs are cut in Coal Measure Sandstones and Millstone Grit, overlain by glacial drift. Haroldston Bridge is a natural arch under a dipping ramp of Coal Measures sandstone, and blocks of cliff are falling away at Haraldston Chins, leaving deep chasms. Below the cliffs to the south is the Sleek Stone (>Fig. 7.7.1), a monoclinal protrusion of

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Fishguard Harbour lies under a cliff on the western side of the bay. Strumble Head is a bold heathy promontory with cliffs exposing pillow lavas, rhyolites, agglomerates and volcanic ash. To the south are coves cut in Ordovician shales between steep bluffs, cliffs and rocky shores on volcanic and intruded rocks. There is an extensive capping of boulder clay, and flint gravel is found on beaches. Abercastle is behind a long narrow drowned valley mouth. The sequence of headlands in dolerite or rhyolite and bays and inlets in softer Ordovician sedimentary rocks continues SW to St. David’s Head. As most of the

7.7

rivers on the St. Davids Peninsula rise close to the north coast and flow south there are few sandy coves on the cliffy northern shore.

References Burd F (1989) The saltmarsh survey of Great Britain. Research and Survey in Nature Conservation No. 17, Peterborough Whittow JB (1960) Some comments on the raised beach platform of south-west Carnarvonshire and on an unrecorded raised beach at Porth Neigwl. Proc Geologis Assoc 71:31–39

7.7  South Wales Along the south coast of Wales the geological formations trend generally west to east, and the coastline intersects outcrops of rock formations that are often steeply dipp­­ ing.  The Pre-Cambrian and Lower Palaeozoic rocks of St. David’s Peninsula are succeeded southward in St. Bride’s Bay by Devonian (Old Red Sandstone) and Carboniferous formations, including Carboniferous Limestone, Millstone Grit and Coal Measures. These extend W–E across south Pembrokeshire, Carmarthen Bay and the Gower Peninsula to Swansea Bay and Porthcawl. There are broad plateaux ending in even-crested cliffs, except where incised valleys open to the coast. In Glamorgan the Carboniferous rocks subside beneath Trias and Lias formations. Keuper Marl forms a coastal fringe north–east from Cardiff, but is largely concealed by Pleistocene glacial drift and Holocene alluvium in the Wentlooge and Caldicot Levels. Mean spring tide ranges increase W–E along the coast. On Skomer Island the range is 5.9 m and at Milford Haven 6.3 m. Stackpole Quay has 7.0 m and Tenby 7.6 m. On Gower the range at Whitford Lighthouse is 7.8 m and in Swansea Bay it increases from 8.6 m at Swansea to 8.7 m at Port Talbot and 8.9 m at Porthcaw l8.9. The increase ­continues as the Bristol Channel narrows into the Severn estuary: Barry has 10.7 m, Cardiff 11.1 m, Newport 11.9 m, Sudbrook 12.4 m and Beachley, at the Severn Bridge, 12.4 m. In Pleistocene times the Devensian Glaciation is thought to have extended to South Wales, the limit (c 20,000 bp) of Irish Sea ice extending from the southern side of St. Brides Bay inland over north Pembrokeshire to meet Welsh Ice that extended back out across Carmarthen Bay and the northern part of the Gower Peninsula, across Swansea Bay to Porthcawl, then inland behind South Glamorgan to a lobe between Cardiff and Newport.

On the Pembrokeshire coast (John 1979) St. David’s Head is a promontory of gabbro. Cliffs cut in Cambrian rocks fringe the coast into Whitesands Bay, which has a typical west-facing Atlantic surf beach, wide at low tide and backed by dunes. Clefts and caves have been etched out along faults in the Cambrian rocks along the shore, and there is an emerged shore platform with a raised beach of sand and shingle. The St. David’s Peninsula is ­dominated by a broad plateau 45–75 m above sea level, generally ­considered to be a marine planation surface with residual higher hills that were once islands when the sea stood at a higher level. The Bitches is a tide race inshore from Ramsey Island, with strong currents as the tide rises and falls. Ramsey Island is a detached slab of the coastal plateau bordered by  cliffs in Ordovician shales between headlands of quartz-porphyry. The mainland coast swings eastward with slope-over-wall cliffs in Pre-Cambrian volcanic rocks, with intruded granites and dolerites forming headlands and islands. East of Pen-y-Cyrfrwy a coastal fringe of Cambrian rocks outcrops in the cliffs, and these continue past Caerfai Bay to the Solva inlet, and on to the mouth of the Cwm-mawr valley, where a major fault marks the beginning of the strongly folded Carboniferous formations (Coal Measures) which dominate the east coast of St. Bride’s Bay. At Newgale the sandy beach, wide at low tide, is backed by a straight ridge of storm-piled grey and brown cobbles. The cliffs are cut in Coal Measure Sandstones and Millstone Grit, overlain by glacial drift. Haroldston Bridge is a natural arch under a dipping ramp of Coal Measures sandstone, and blocks of cliff are falling away at Haraldston Chins, leaving deep chasms. Below the cliffs to the south is the Sleek Stone (>Fig. 7.7.1), a monoclinal protrusion of

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⊡⊡ Fig. 7.7.1 The Sleek Stone, St. Bride’s Bay. (Courtesy Geostudies.)

Coal Measures sandstone, steeper and eroded on its northern flank, the upper surface of a sandstone layer from which overlying softer rocks have been removed. Broad Haven has been cut into strongly-folded Coal Measures sandstones, with smooth bedding planes exposed in the cliffs and Little Haven has been cut into soft shales between sandstone headlands. At Falling Cliff, where subsidence is due partly to the collapse of old coal mineshafts, the Coal Measure sandstones come to an end against dolerite. The coastline swings westward and the cliffs pass into scrubby bluffs on Coal Measures with bolder headlands on Pre-Cambrian volcanic rocks. There are several steepsided coves, as at Foxes Holes, where the Pre-Cambrian rocks give place to Old Red Sandstone, with inlets along small faults as at Musselwick Sands. Beyond Martins Haven is a broad promontory of igneous rocks, Deer Park, with Wooltack Point a headland at the NW corner, and The Anvil to the SE. The western cliffs look across a strait where the tidal currents are strong to Midland Isle, then another strait to Skomer Island. This island is mainly composed of volcanic rocks exposed in cliffs, and Grassholm Island, out to the west, is a cliffy island of basalt. Cliffs on the south coast of Deer Park are strongly influenced by the structure of outcrops of Silurian slate (>Fig. 7.7.2), and the coast runs out to a small headland behind Gateholm Island. This flat-topped island on steeply-dipping Old Red Sandstone is attached to the coast by a rocky reef. Offshore to the SW is Skokholm, a flattopped cliff-edged island of Old Red Sandstone rising 33 m above sea level.

The cliffs run on past Raggle Rocks, and The Three Chimneys are vertically projecting sandstone beds separated by erosion of intervening soft mudstones. Along Marloes Bay the cliffs are cut in colourful strongly folded Silurian rock formations, with bedding planes exposed in the cliff face where the dip is seaward Marloes Beach (>Fig. 7.7.3) is wide and sandy at low tide, when jagged reefs and stacks of Silurian rock protrude from the sand. ­Slope-over wall profiles extend south to St. Ann’s Head and round to Dale Fort. To the east the coast runs along the northern side of Milford Haven, with cliffs in Old Red Sandstone and Silurian rocks. Sandy Haven is a wide inlet at low tide with green weed and brown wrack on gravel and mud. Red Marls underlie the south-facing slopes west and east of the town of Milford Haven. Milford Haven is a large branching ria formed by Holocene marine submergence of the Cleddau River and its tributaries, incised into the Pembrokeshire plateau. In the Cresswell River broad sloping tidal mudflats fringe the low tide channel, and there are grassy salt marshes (>Fig. 7.7.4). On the south coast of Milford Sound coastal slopes descend to low cliffs and rocky shores. Angle Bay is sheltered from strong wave action, and has low tide rocky outcrops and gravel with much seaweed, descending to mudflats. The shore is fringed by salt marsh, dissected by tidal creeks. At the western end of the Angle valley is West Angle Bay, a synclinal cove in Carboniferous Limestone. Slopeover-wall cliffs run out to a headland at East Blockhouse, cut first in Carboniferous Limestone then Old Red

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⊡⊡ Fig. 7.7.2 Cliff in dipping Silurian slates at Horse Neck, near Gateholm, with Skomer Island in the distance. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.3 Marloes Sands. (Courtesy Geostudies.)

Sandstone rising on the southern side of the Angle ­syncline. The cliffs continue to Freshwater West, where Silurian rocks emerge from beneath the Old Red Sandstone as the shore platform disappears beneath the wide Broomhill Sands. Dunes and bluffs facing WSW back a wide sandy shore receiving Atlantic surf and dunes overlie a cobble beach. To the south Old Red Sandstone dips beneath Carboniferous rocks on Limney Head, where the coastline turns ESE.

A very fine stretch of Carboniferous Limestone coast extends ESE from Linney Head, and is well known for its geological structures and the cliff scenery formed by dissection along joints and bedding planes (John 1973). It includes a prominent natural arch at the Green Bridge of Wales (>Fig. 7.7.5) and the outlying Elegug stacks (>Fig. 7.7.6). The cliffs continue eastward to St. Govan’s Head (>Fig. 7.7.7), a bold promontory that shows high vertical

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⊡⊡ Fig. 7.7.4 Salt marsh and tidal mudflats in Cresswell River, Milford Haven. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.5 The Green Bridge of Wales. (Courtesy Geostudies.)

cliffs dissected along rectangular joints, so that blocks of ­limestone have fallen to the shore. Long Matthew Point marks the end of the south-facing cliffed coast. The coastline turns northward and New Quay is a sanded inlet at low; at high tide it is much like a Mediterranean calanque. Behind Broad Haven two valleys incised into the limestone plateau converge east of Bosherston. They were formerly tidal inlets, but in the eighteenth century they were enclosed by a dam as

Bosherston Lakes. The dunes of Stackpole Warren have drifted across from Broad Haven. The cliffs from Stackpole Point north to Trewent Point show the effects of struc­­ ture  and stratigraphy, with sloping ramps on dipping ­bedding-planes, hollows and valleys on shales, and clefts cut out along joints and bedding planes. To the north is Freshwater East Bay, with a wide sandy beach backed by dunes emplaced by easterly winds. It occupies an anticlinal flexure, with Silurian rocks

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⊡⊡ Fig. 7.7.6 Elegug Rocks. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.7 Cliffs of Carboniferous Limestone extending east to St. Govan’s Head. (Courtesy Geostudies.)

­ ipping steeply away on either side of a central outcrop of d Ordovician rocks. The slope-over-wall coast to the east is on steeply-dipping Old Red Sandstone, which in places stand almost vertical, and strike parallel to the coast. Swanlake Bay is a fault-bounded re-entrant with a broad shore platform cut across steeply-dipping Old Red Sandstone strata, also seen at Manorbier (>Fig. 7.7.8). In Skrinkle Haven are vertical strata from the upper Old Red Sandstone to Carboniferous Limestone, which outcrops along the steep coast eastward to Giltar Point

and on outlying Calder Island. Limestone cliffs run northward past Tenby, with Millstone Grit and Coal Measure sandstones towards Monkstone Point. Towards Saundersfoot the coast consists of wooded bluffs with basal cliffs of varying height in folded and faulted Coal Measures, and the beach includes ballast dumped by ships that came in to collect coal. East of Amroth the underlying Millstone Grit outcrops in cliffs that run out to Telpyn Point, and the Carboniferous Limestone forms cliffs along to Dolwen Point. At Pendine

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⊡⊡ Fig. 7.7.8 Shore platform cut across steeply dipping Old Red Sandstone in Manorbier Bay. (Courtesy Geostudies.)

these cliffs have several high caves have been cut out along vertical joints in the limestone (>Fig. 7.7.9). Pendine Sands, up to 800 m wide at low tide, border a dune-capped foreland that is backed by bluffs that extend to Ginst Point (Savigear 1952). This shelters a salt marsh beside the sandy Taf estuary, which meets the Tywi and Gwendraeth estuaries in an area of inwashed intertidal sand in the NE of Carmarthen Bay. South of the Gwendraeth estuary these sands are backed by a wide sandy foreland that includes dunes at Tywyn Burrows and Pembrey Burrows. To the east the coast is low and marshy to Burry Port, beside the wide Loughor estuary north of the Gower Peninsula. On its southern coast is the wide Llanrhidian salt marsh (>Fig. 7.7.10) with winding creeks and scattered pools and salt pans, some of which formed as the result of sub-surface tunnelling. The Gower Peninsula consists of folded and faulted Carboniferous Limestone with synclinal furrows occupied by Millstone Grit shales forming bays and anticlinal ridges forming headlands where Old Red Sandstone outcrops. On the west coast, south of the spit at Whitford Burrows, Burry Holms is an island of limestone seaward of a sandy intertidal isthmus. Rhossili Bay has a beach 5 km long, the northern part backed by dunes, the southern part by a gravel beach, low receding cliffs cut into an apron of glacial drift and a gently sloping terrace that rises to steep slopes of Old Red Sandstone on Rhossili Down (>Fig. 7.7.11). Worms Head is a long narrow promontory with a rugged island of strongly folded Carboniferous Limestone at the western end of an intertidal boulder-strewn shore platform.

The SW coast of the Gower Peninsula has steep escarpment cliffs in landward-dipping Carboniferous Limestone rising to a broad plateau. There are there are many minor bays and indentations related to faults or major joints, and segments of 4–5 m raised beach. Paviland Cave is one of several Palaeolithic occupation sites with animal bones and prehistoric implements. A high limestone promontory runs out to flattopped Port Eynon Point and Port Eynon occupies a broad synclinal bay cut in shales along a syncline of Millstone Grit. Oxwich Point is a 75 m high promontory with cliffs on the southern side passing into wooded bluffs beside Oxwich Bay to the north. This has been cut in soft Millstone Grit shales, and has a 2 km wide gentle intertidal sandy shore. On the northern side Great Tor, a barnacled rock pinnacle, protrudes at the western end of Three Cliffs Bay (>Fig. 7.7.12), which is bordered by limestone cliffs and bluffs and has a wide beach backed by grassy dunes. Pwll-du Head is a high (225 m) limestone promontory. To the east Pwll-ddu Bay has been cut out along fault lines, and contains shingle beach ridges. A succession of cliffy headlands and small bays extends to Mumbles Head. Swansea Bay has been cut in soft Upper Carboniferous rocks at the submerged mouths of several ­confluent river valleys. East of the Tawe estuary Kilvey Hill (171 m) rises steeply behind a narrow coastal plain, and sandy beaches backed by dunes extend to Porthcawl. At Porthcawl beach depletion has been countered by making a tarmac shore with embedded cobbles.

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7.7

⊡⊡ Fig. 7.7.9 Caves in Carboniferous Limestone in the cliff at Pendine. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.10 Llanrhidian Marsh. (Courtesy Geostudies.)

Newton Point is a headland projecting southward, and beyond it a curving bay with a broad sandy shore is backed by the grassy dunes of Newton Burrows and Merthyr Mawr. The beach ends at the mouth of Ogmore River, and at Ogmore-by-sea the coast consists of a steep grassy escarpment in thinly-bedded Lias with an apron of  periglacial Head, descending to cliffs cut in the layered  limestone. To the east Carboniferous Limestone strata form sloping stepped ledges, and at Southern­­ down an almost vertical cliff in gently dipping Lias rests

c­ onformably upon Carboniferous Limestone which runs across the shore. At  Nash Point (>Fig. 7.7.13) the cliffs are cut in layered Lias limestone and shales are fronted at low tide by shore platforms up to 500 m wide, with swirling, undulating limestone ledges separated by sinuous scarplets, dissected along joints and bedding planes l (Trenhaile 1972). The limestone cliffs continue SE past Summerhouse Point to Aberthaw and Barry Island (Williams et al. 1993). l Sully Island is a grassy and scrubby island of Triassic rocks

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South Wales

⊡⊡ Fig. 7.7.11 Rhossili Bay, Gower Peninsula. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.12 Three Cliffs Bay, Gower Peninsula. (Courtesy Geostudies.)

with bordering low cliffs and limestone ledges with beach gravels. The low Lias limestone cliffs continue to Lavernock Point, where the coast swings sharply northward to Penarth, crossing an outcrop of Keuper Marl in the core of the Penarth anticline. Off Lavernock Point are Flat Holm and Steep Holm, islands of Carboniferous Limestone in the wide mouth of the Severn estuary.

To the east is the wide mouth of the Taff estuary opening into Cardiff Bay, and the extensive Cardiff docks and waterfront. Beyond the Rhymney estuary is a wide coastal lowland, Peterstone Wentlooge, a former marsh that was drained and reclaimed behind a sea wall. The broad meandering estuary of the River Usk interrupts the coastal plain at Newport, but it resumes to the

South Wales

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⊡⊡ Fig. 7.7.13 Shore platform cut across dipping Lower Lias limestone near Nash Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.14 Cliff in Trias at Black Rock, with fronting salt marsh. (Courtesy Geostudies.)

east as the Caldicot Levels, another area of reclaimed marsh. Much of the coastline consists of a high sea wall, grassy on the inner slope and armoured with boulders on the seaward side, where wide mudflats are exposed at low tide. There are minor cliffs at Sudbrook and Black Rock (>Fig. 7.7.14). At Chepstow the River Wye meanders through a gorge between cliffs cut in massive Carboniferous Limestone before crossing low ground and marshland to open into the Severn estuary west of Beachley Point.

References John BS (1979) The geology of Pembrokeshire. Abercastle Savigear RAG (1952) Some observations on slope development in South Wales. Trans Inst Brit Geogr 31:23–42 Trenhaile AS (1972) The shore platforms of the Vale of Glamorgan. Trans Inst Brit Geogr 56:127–144 Williams AT, Davies P, Bomboe P (1993) Geometrical simulation studies of coastal cliff failures in Liassic strata, South Wales. Earth Surf Process Landforms 18:703–720

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The Severn Estuary

7.7.1  The Severn Estuary The River Severn flows through the wide Vale of Gloucester and downstream from Gloucester widens into the Severn estuary, opening into the Bristol Channel. The South Wales coast east from Penarth shows some estuarine features, as does the Avon coast northeast from Clevedon, but here the shores of the Severn estuary are those upstream from the first Severn Bridge, which crosses from Beachley to Aust. The river is tidal to well above Gloucester, but above Minsterworth the margins are essentially river banks. The Severn estuary is megatidal (mean spring tide range >6 m) and funnel-shaped, narrowing upstream. Mean spring tide range decreases from 12.4 m at Beach­­ ley to 8.8 m at Sharpness Dock and about 2 m at Gloucester. Ebb and flow currents can attain 20 kph, and at spring tides the Severn Bore travels rapidly upstream (>Fig. 7.7.1.1). Wave action is limited by tidal rise and fall, but can be strong with onshore winds on bordering shores at high tide. The present morphology of the Severn estuary results from sediment accumulation in deeper channels that were  cut during Pleistocene phases of low sea level and ­submerged during the Late Quaternary (Flandrian) marine transgression. At low tide, the river banks are fringed by a muddy intertidal slope that widens as the tide range increases downstream. The muddy deposits have been derived from the broad clay vales drained by the Severn and its tributaries, but there has been much

­ ovement of muddy sediment from the Bristol Channel, m brought back upstream by rising tides. Mudflats are shaped by tidal ebb and flow, as well as by wave action as the tides rise and fall (Pethick 1996). The stratigraphy of bordering salt marsh terraces shows alternations of almost horizontal layers of mud interdigitating with backshore peat horizons some of which outcrop as ledges on the seaward slope (Allen 1992).

1.  The North Shore Above Beachley the estuarine shore is fringed by salt  marshes, backed at Sedbury by minor cliffs and steep  wooded bluffs, cut in Rhaetic and Lower Lias (>Fig. 7.7.1.2). The bordering alluvium is interrupted by segments of low cliff at Lydney and Purton, and cliffs cut in thinly bedded grey Triassic limestone and pink shales over red-brown marls towards Awre. Further upstream the meandering river intersects Triassic rocks on the north bank near Newnham.

2.  The South Shore The Severn widens rapidly below Longney, and at Epney the bank at high tide is protected by a stone wall. The river meanders ESE past Hock Cliff, where layers of ⊡⊡ Fig. 7.7.1.1 The Severn Bore passing Stonebench. (Courtesy Geostudies.)

The Severn Estuary

limestone and shale outcrop in the cliff and on the shore (>Fig. 7.7.1.3). At Slimbridge managed retreat is being implemented on land that was formerly embanked and reclaimed. A new clay bank has been built 40 m back from the pre­ vious one, leaving a new intertidal strip to be managed

⊡⊡ Fig. 7.7.1.2 Salt marsh in front of Sedbury Cliff on the Severn shore. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.1.3 Cracking shales over a ledge of limestone at Hock Cliff on the Severn estuary. (Courtesy Geostudies.)

7. 7.    1

as an estuarine wetland. At Purton the grassy salt marsh ends seaward in a low cliff (>Fig. 7.7.1.4), below which are sloping mudflats with mounds of mud bearing Spartina. Similar features extend downstream, but the shore has  been modified by a large tidal reservoir near Shepperdine.

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Gloucestershire, Somerset and North Devon

⊡⊡ Fig. 7.7.1.4 Erosion of the Severn-side salt marsh terrace at Tites Point, near Purton. (Courtesy Geostudies.)

References Allen JRL (1992) Tidally induced salt marshes in the Severn estuary, southwest Britain. In: Allen JRL, Pye K (eds) Saltmarshes:

morphodynamics, conservation and engineering significance. Cambridge University Press, Cambridge, New York, pp 123–147 Pethick JS (1996) The geomorphology of mudflats. In: Nordstrom KE, Roman CT (eds) Estuarine shores. Wiley, New York, pp 185–211

7.8  Gloucestershire, Somerset and North Devon The coast of Gloucestershire, Somerset and North Devon begins where the Severn estuary runs beside the Vale of Belvoir, a lowland of Keuper Marl. The Jurassic formations of the Cotswolds do not extend to the coast, but there are minor outcrops of Trias at Aust. The Carboniferous Limestone of the Bristol district reaches the seaboard between Portishead and Clevedon (with some Upper Devonian), and again in the westward promontories north and south of Weston-super-mare. The Somerset Levels form a lowland behind Bridgwater Bay, and to the west the Lower Lias outcrops in low cliffs between Lilstock and Blue Anchor Point. Keuper Marl occupies the Minehead lowland and the valley opening to Porlock, and the Devonian rocks on the northern limb of the mid-Devon syncline form the steep coast at Selworthy Beacon and along the  wnorthern slopes of Exmoor between Porlock and Ilfracombe. The Devonian rocks strike ESE–WNW and the coast cuts successively across Lower and Middle Devonian formations as far as Morte Point and Upper Devonian ­formations south to Saunton. Strongly folded Carbonif­ erous Culm Measures (dark grey and greenish shales and

s­ andstones with grit bands and some limestones) then dominate the coast from the ­Taw-Torridge estuary to Hartland Point and south to the border with Cornwall. Mean spring tide range is large in the Severn estuary, attaining more than 12 m off Avonmouth. As the estuary widens into the Bristol Channel the tide range diminishes from 11.9 m at Clevedon to 11.3 m at Weston-super-mare, 10.5 m at Watchet, 9.9 m at Minehead, 9.5 m in Porlock Bay and 8.5 m at Ilfracombe, and as the Bristol Channel opens to the Atlantic coast they diminish further to 7.4 m at Clovelly 7.2 m on Lundy Island, and just under 7 m at  Hartland Quay. These megatidal conditions result in strong currents in the Severn estuary and wide intertidal zones, notably in Bridgwater Bay where marshes and mudflats are extensive.

1.  Gloucestershire On the Gloucestershire coast of the Severn estuary at Aust there is a cliff cut in Triassic sediment: the Red Marls

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⊡⊡ Fig. 7.7.1.4 Erosion of the Severn-side salt marsh terrace at Tites Point, near Purton. (Courtesy Geostudies.)

References Allen JRL (1992) Tidally induced salt marshes in the Severn estuary, southwest Britain. In: Allen JRL, Pye K (eds) Saltmarshes:

morphodynamics, conservation and engineering significance. Cambridge University Press, Cambridge, New York, pp 123–147 Pethick JS (1996) The geomorphology of mudflats. In: Nordstrom KE, Roman CT (eds) Estuarine shores. Wiley, New York, pp 185–211

7.8  Gloucestershire, Somerset and North Devon The coast of Gloucestershire, Somerset and North Devon begins where the Severn estuary runs beside the Vale of Belvoir, a lowland of Keuper Marl. The Jurassic formations of the Cotswolds do not extend to the coast, but there are minor outcrops of Trias at Aust. The Carboniferous Limestone of the Bristol district reaches the seaboard between Portishead and Clevedon (with some Upper Devonian), and again in the westward promontories north and south of Weston-super-mare. The Somerset Levels form a lowland behind Bridgwater Bay, and to the west the Lower Lias outcrops in low cliffs between Lilstock and Blue Anchor Point. Keuper Marl occupies the Minehead lowland and the valley opening to Porlock, and the Devonian rocks on the northern limb of the mid-Devon syncline form the steep coast at Selworthy Beacon and along the  wnorthern slopes of Exmoor between Porlock and Ilfracombe. The Devonian rocks strike ESE–WNW and the coast cuts successively across Lower and Middle Devonian formations as far as Morte Point and Upper Devonian ­formations south to Saunton. Strongly folded Carbonif­ erous Culm Measures (dark grey and greenish shales and

s­ andstones with grit bands and some limestones) then dominate the coast from the ­Taw-Torridge estuary to Hartland Point and south to the border with Cornwall. Mean spring tide range is large in the Severn estuary, attaining more than 12 m off Avonmouth. As the estuary widens into the Bristol Channel the tide range diminishes from 11.9 m at Clevedon to 11.3 m at Weston-super-mare, 10.5 m at Watchet, 9.9 m at Minehead, 9.5 m in Porlock Bay and 8.5 m at Ilfracombe, and as the Bristol Channel opens to the Atlantic coast they diminish further to 7.4 m at Clovelly 7.2 m on Lundy Island, and just under 7 m at  Hartland Quay. These megatidal conditions result in strong currents in the Severn estuary and wide intertidal zones, notably in Bridgwater Bay where marshes and mudflats are extensive.

1.  Gloucestershire On the Gloucestershire coast of the Severn estuary at Aust there is a cliff cut in Triassic sediment: the Red Marls

Gloucestershire, Somerset and North Devon

­ verlain by Tea Green Marls and a capping of dark grey o shales (>Fig. 7.8.1). To the SW the coast is fringed by salt marshes that have been partly reclaimed behind sea walls. At Avonmouth at low tide, wide, muddy, sandy and gravelly flats are exposed. Low cliffs of Carboniferous Limestone run out to Battery Point and Denny Island is a limestone outcrop offshore, while in Kilkenny Bay salt marshes front the esplanade. Steep coastal slopes in Carboniferous Limestone and Old Red Sandstone descend to scrubby bluffs and low basal cliffs behind rocky and gravelly shores, with mud exposed at low tide. Limestone cliffs continue to the west until River Kenn flows out by Wain Hill. The coast is then low-lying and marshy, with high tide embankments interrupted by the mouth of the River Yeo; then St Thomas’s Head marks the beginning of Middle Hope, a narrow coastal ridge of Carboniferous Limestone mantled by Pleistocene Head deposits. The ridge declines along a promontory to Sand Point, and the steep slopes and low basal cliffs run back along the northern shore of Sand Bay (>Fig. 7.8.2). Sand Bay is a 3 km wide west-facing bay between promontories of Carboniferous Limestone, backed by a sandy beach and low dunes. At low tide up to 2 km of sand and mud is exposed, with numerous sub-parallel creeks running seaward. Near the northern end is a Spartina marsh (>Fig. 7.8.3). Weston-super-mare stands behind Weston Bay, which is similar to Sand Bay, with a wide area of runnelled, silty

⊡⊡ Fig. 7.8.1 Aust Cliff. (Courtesy Geostudies.)

7.8

sand passing seaward below low tide mudflats. It is ­bordered southward by Brean Down, a ridge of Carboniferous Limestone running out to Howe Rock. The River Axe opens into the southern part of the bay, delivering muddy sediment to the shore.

2.  Somerset The long, narrow steep-sided ridge of Carboniferous Limestone running out to Brean Down rises to 97 m, and declines westward to Howe Rock. Steep Holm and  Flat Holm are limestone islands offshore. The Carboniferous Limestone dips steeply north, so that a dip-slope faces Weston Bay and the southern slopes are an escarpment with basal cliffs. To the south, intertidal sand and mudflats are backed by a wide sandy beach and a series of parallel dune ridges. At Burnham on Sea the sandy beach gives way to salt marshes at the mouth of the Brue estuary, which opens to a muddy shore. The larger Parrett estuary comes in from the south beside Stert Point and the elongated Stert Island, a shingle patch which rises above smoothly-sloping mudbanks offshore. There has been mud accretion in Bridgwater Bay and rapid spread of marsh west of the mouth of the Parrett since Spartina was introduced in about 1928 (Ranwell 1964). To the west, the marshes fade out and shingle beaches and beach ridges are arranged in a series of elongated spits

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⊡⊡ Fig. 7.8.2 Limestone promontory north of Sand Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.3 Spartina in Sand Bay. (Courtesy Geostudies.)

with recurves indicating intermittent eastward growth. A  backshore embankment becomes a concrete wall at Stolford, and a gravelly beach that extends westward to Hinkley Point consists of limestone and sandstone pebbles that have drifted from disintegrating Liassic rock outcrops to the west. Hinkley Point power station stands on a low plateau of Liassic limestone and to the west the coastline consists of cliffs fronted by broad shore platforms cut across undulating layers of grey Lias limestone and blue shale, which

­ utcrop in swirling patterns on the wide foreshore o exposed as the tide falls (>Fig. 7.8.4). To the west the vertical cliffs in stratified Lias rise along the seaward side of a coastal ridge to Quantocks Head (over 50 m), past the Blue Ben headland to St Audrie’s Bay. This has a wide foreshore of red sand and sandy mud, backed by disintegrated boulders, a sector of sandy beach at mean high tide, and an upper beach of grey and pink gravel along the base of red cliffs. West of St Audrie’s Bay, the Lias cliffs continue to Watchet. To the west Lias

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7.8

⊡⊡ Fig. 7.8.4 Cliffs and shore platform west from Kilve. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.5 Cliff near Blue Anchor Point cuts across a fault that juxtaposes Lower Lias limestone and shale on the left and Keuper Marl on the right. (Courtesy Geostudies.)

a­ lternates with Keuper Marl in cliff outcrops along Warren Bay to Blue Anchor Point (>Fig. 7.8.5). West from Blue Anchor is a coastal plain fringed by a shingle beach, extending out to a lobate foreland, and then fronting Minehead, where it has been artificially augmented. Beyond Minehead harbour a steep forested slope rises past Culvercliff to Greensleigh Point, and continues along the northern slopes of Selworthy Beacon in steeply dipping, much faulted hard Devonian sandstone.

At Hurlstone Point the steep slope runs in on the eastern side of the Porlock lowland. Porlock Bay has a long beach of cobbles and pebbles (Jennings et al. 1998), backed by a depositional plain that includes areas of reclaimed wetland. The beach was frequently breached by storms, but the breaches generally healed within a week. A breach formed in 1998 has persisted, and the area flooded by high tides has become a salt marsh. Former drains have become wider and deeper tidal creeks, and waterfalls form during the ebb (>Fig. 7.8.6).

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⊡⊡ Fig. 7.8.6 Waterfall over a receding salt marsh cliff, Porlock Bay. (Courtesy Geostudies.)

The beach curves round to the little harbour at Porlock Weir, in the lee of a spit of pebbles and cobbles, the Porlock Weir Ridge, which diverges from the coast to the west and is backed by a grassy marsh. The spit grew as the result of  eastward drifting of beach material derived from cliff recession and slumping and shore erosion of Devonian sandstones to the west, but the supply of pebbles and cobbles appears to have diminished during recent decades, and the spit is unstable. At its western end it passes beneath the steep coast of Culbone (>Fig. 7.8.7) to the Devonshire border at Glenthorne.

3. The North Devon Coast and Lundy Island West from the deep valley at Glenthorne, the bold coast of North Devon curves gradually out to Foreland Point (Arber 1911). Cut in Devonian formations, it is a high, steep hogsback coast with convex slopes curving down to small basal cliffs and a rocky boulder-strewn shore. The steep coast is interrupted at Lynmouth (>Fig. 7.8.8) where the gorges of the East and West Lyn, cut through Devonian slates and grits, converge, and a river descends over ­boulders to a deltaic fan of rock debris exposed at low tide. After wet weather, and when snow melts on the moors, the Lyn discharges large quantities of water, and there was a devastating flood in 1952.

Forested slopes rise along the coast west of Lynmouth, and a high craggy coastal ridge of hard shales and grits is crowned with tors, and has buttresses and screes on the steep coastal slope. The ridge is backed by The Valley of Rocks, a dry valley running parallel to the coast and opening through a gap to Lee Bay. At Heddon’s Mouth, a deep V-shaped valley has been cut across the strike of the Lower Devonian rocks, and the steep coastal slopes continue westward to Blackstone Point, where the slopes rise abruptly to Great Hangman (318 m) and Little Hangman (218 m). Combe Martin is a long, straight, incised valley, and at its mouth is a beach of pebbly sand between planed-off outcrops of steeply dipping sandstone. To the west a coastal ridge on southward dipping slates is divided by inset sandy coves at Sandy Bay and Small Mouth, and backed by a long inlet, Watermouth, which drains out at low tide to expose sand and gravel. The coast then rises and steepens to Widmouth Head and Rillage Point. The deep Chambercombe valley opens to Hele Bay, and the steep coast then runs out to Beacon Point above a rocky shore. On its western side a series of coves and rocky ribs extends to Larkstone Beach and the bay that is on its western side has a sandy beach in Ilfracombe Harbour, sheltered by a breakwater. Ilfracombe occupies a steep-sided east-west valley behind a coastal ridge interrupted by sandy Wildersmouth Bay. To the west, a steep coast runs along Seven Hills,

Gloucestershire, Somerset and North Devon

7.8

⊡⊡ Fig. 7.8.7 The western end of the Porlock Weir Ridge, showing the steep coast towards Culbone. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.8 The steep coast at Lynmouth, showing a lowtide delta of boulders and gravel carried down by the River Lyn during episodes of flooding. (Courtesy Geostudies.)

interrupted by a deep valley descending to Lee Bay. Morte Point runs out sharply westward with scrubby knolls that become craggy and projecting slate tors like teeth on a dragon’s back. The coastal slopes are mantled with periglacial rubble (Head) and descend to cliffs of ­collapsing grey slate that overlook a white water race as tides meet waves. Lundy Island is about 30 km west of Morte Point. It is almost flat-topped, rising 100–120 m above sea level, and

consists of granite, with a small outcrop of Devonian rock at the SE end. Its coasts show a slope-over-wall profile, the slope mantled by Pleistocene periglacial Head deposits above granite cliffs that are generally higher on the more exposed western side. Morte Bay has been cut out in relatively soft shales, and contains Barricane Beach, famous for shell accumulations in clefts between rocky ribs. Woolacombe Sand has surf breaking on fine, firm, wide, calcareous sands

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exposed at low tide, with megaripples and pools emerging as the tide ebbs. At the southern end is the broad Baggy Point peninsula with steep rock faces, and the craggy cliffs run back into sandy Croyde Bay, at the mouth of a valley cut into southward-dipping Upper Devonian beds. At Saunton Head-mantled slopes descend to an emerged shore platform with outcrops of bedded calcareous sandstone, a Pleistocene beach and dune sequence that incorporates glacial erratics (Campbell and Gilbert 1998). Saunton Sands are up to 500 m wide at low tide, backed by dunes (Braunton Burrows) up to 30 m high. At their southern end a recurved spit - Crow Point - borders the wide TawTorridge estuary where at low tide the river channel flows to the sea between wide sands, with waves breaking over Bideford Bar (Kidson 1963). At Fremington Quay on the southern side of the Taw estuary the river at low tide meanders close to the shore. To the east a beach fringes low cliffs cut in earthy gravel containing rounded and angular pebbles, mainly of culm but also some quartzite, basalt, Old Red Sandstone, Scottish erratics (including dolerite, granite, quartzite, andesite, tuff, gneiss and flint) as well as Irish Sea foraminifera. These erratics indicate a Pleistocene glacial drift deposit. Appledore is a small port on a promontory at the mouth of the Torridge River, where mean spring tide range is 7.4 m. The coast is walled, but there is an intertidal area of angular culm gravel and sand shoals.

Dissected salt marsh exists in a bay to the west, behind a broad foreland, Northam Burrows. On the western shore, extending north from Westward Ho!, is the Pebble or Popple Ridge, a barrier spit 25 m wide and up to 3.5 m above high spring tide level, consisting of rounded boulders (>Fig. 7.8.9). From Westward Ho! cliffs cut in the Culm Measures and truncate a series of east-west strike ridges and valleys. Steep scrubby coastal slopes are fronted by beaches of grey and pink gravel drifting east towards the Popple Ridge at Westward Ho! At Portledge, cliffs are cut in an outlier of Triassic Keuper Marl with sandstone and breccia layers (>Fig. 7.8.10). At Peppercombe the steep coast descends to a beach of coarse grey cobbles and boulders. East of the mouth of Peppercombe valley the steep forested coastal slopes have hollows and benches aligned with the geological strike, formed by slumping. Buck’s Mills is in a deeply incised forested valley which interrupts the steep coastal  escarpment. Clovelly stands on an anticlinal pro­montory, and the coast westward to Hartland Point is steep and wooded, incised by deep valleys, as at Mouth Mill. At Hartland Point the coast turns southward to face the Atlantic. High vertical cliffs rise to a plateau 60–80 m above sea level that truncates the strongly folded Culm Measures exposed in the cliffs and on the shore. There are smooth steeply inclined bedding planes in the cliffs and truncated strata and rocky ribs on the shore. ⊡⊡ Fig. 7.8.9 The Popple Ridge at Northam Burrows. (Courtesy Geostudies.)

Gloucestershire, Somerset and North Devon

Several short, steep streams descend from a drainage divide a few kilometres inland to end in coastal waterfalls (>Fig. 7.8.11). Caves, ribs and pillars have been cut out in sharp anticlines north of Hartland Quay, where the cliffs rise 100 m and the cliff base is fringed by intertidal shore platforms, where fan-folded strata have been dissected.

⊡⊡ Fig. 7.8.10 Cliffs of Keuper Marl at Portledge. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.11 Waterfall at Speke’s Mill Mouth. (Courtesy Geostudies.)

7.8

Between Speke’s Mill Mouth and Welcombe the cliffs are more than 120 m high. Several buttress reefs of thick hard shale run out from the coast, forming an irregular intertidal shore; Knaps Longpeak (>Fig. 7.8.12) is a salient of hard shale. At Welcombe Mouth a deep valley with ­bordering terraces of gravelly Head opens between cliffs to a rocky shore. A steep bluff extends to the mouth of the

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The North Coast of Cornwall

⊡⊡ Fig. 7.8.12 Knaps Longpeak, on Carboniferous (Culm sand­­stone. (Courtesy Geostudies.)

steep-sided Marsland valley, which marks the boundary between Devon and Cornwall.

References Arber EAN (1911) The Coast Scenery of North Devon. Dent, London Campbell S, Gilbert A (1998) The Croyde-Saunton coast. Quaternary of South-West England. Geol Conserv Rev Ser 14:214–224

Jennings SC, Orford JD, Canti M, Devoy RJN, Straker V (1998) The role of relative sea-level rise and changing sediment supply on Holocene gravel barrier development: the example of Porlock, Somerset. Holocene 8:165–181 Kidson C (1963) The growth of sand and shingle spits across estuaries. Zeitschrift für Geomorphologie 7:1–22 Ranwell DS (1964) Spartina salt marshes in southern England: Bridgwater Bay. J Ecol 52:79–94

7.9.1  The North Coast of Cornwall 1.  Introduction The north coast of Cornwall is bold, and receives strong swell which has been transmitted from stormy areas in the Atlantic Ocean, as well as occasional large waves generated by gales in the Western Approaches (Bird 1998). Mean spring tide ranges are from 5.9 m at St. Ives and 6.2  m at Perranporth to 6.6 m in the Camel estuary at Padstow. The coast is relatively straight, with high cliffs and rocky headlands, and short steep valleys cut by streams, some of which end in coastal waterfalls. There are  slope-over-wall coasts with steep, vegetated slopes descending to basal rocky cliffs, particularly on the flanks of headlands, and in several places the coastal slope has been removed by erosion and there are vertical cliffs, as in Watergate Bay north of Newquay and on the coast north

of Bude. The highest point on the coast is 213 m at the crest of the steep slope at High Cliff, north of Tintagel, but the highest rocky cliffs are between Navax Point and St. Agnes Head, from Newquay to Trevose Head, on Pentire Point and north of Bude.

2.  The Coastline South of the Devonshire border at Marsland Mouth the coast rises to steep cliffs cut in strongly folded and much faulted Carboniferous Culm Measures (Bude Formation). The Bude Formation consists of sandstones and shales in outcrop in sharp anticlines and synclines exposed and dissected in the cliffs and shore platforms up to 200 m wide at low spring tides. A dissected plateau comes to the coast,

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The North Coast of Cornwall

⊡⊡ Fig. 7.8.12 Knaps Longpeak, on Carboniferous (Culm sand­­stone. (Courtesy Geostudies.)

steep-sided Marsland valley, which marks the boundary between Devon and Cornwall.

References Arber EAN (1911) The Coast Scenery of North Devon. Dent, London Campbell S, Gilbert A (1998) The Croyde-Saunton coast. Quaternary of South-West England. Geol Conserv Rev Ser 14:214–224

Jennings SC, Orford JD, Canti M, Devoy RJN, Straker V (1998) The role of relative sea-level rise and changing sediment supply on Holocene gravel barrier development: the example of Porlock, Somerset. Holocene 8:165–181 Kidson C (1963) The growth of sand and shingle spits across estuaries. Zeitschrift für Geomorphologie 7:1–22 Ranwell DS (1964) Spartina salt marshes in southern England: Bridgwater Bay. J Ecol 52:79–94

7.9.1  The North Coast of Cornwall 1.  Introduction The north coast of Cornwall is bold, and receives strong swell which has been transmitted from stormy areas in the Atlantic Ocean, as well as occasional large waves generated by gales in the Western Approaches (Bird 1998). Mean spring tide ranges are from 5.9 m at St. Ives and 6.2  m at Perranporth to 6.6 m in the Camel estuary at Padstow. The coast is relatively straight, with high cliffs and rocky headlands, and short steep valleys cut by streams, some of which end in coastal waterfalls. There are  slope-over-wall coasts with steep, vegetated slopes descending to basal rocky cliffs, particularly on the flanks of headlands, and in several places the coastal slope has been removed by erosion and there are vertical cliffs, as in Watergate Bay north of Newquay and on the coast north

of Bude. The highest point on the coast is 213 m at the crest of the steep slope at High Cliff, north of Tintagel, but the highest rocky cliffs are between Navax Point and St. Agnes Head, from Newquay to Trevose Head, on Pentire Point and north of Bude.

2.  The Coastline South of the Devonshire border at Marsland Mouth the coast rises to steep cliffs cut in strongly folded and much faulted Carboniferous Culm Measures (Bude Formation). The Bude Formation consists of sandstones and shales in outcrop in sharp anticlines and synclines exposed and dissected in the cliffs and shore platforms up to 200 m wide at low spring tides. A dissected plateau comes to the coast,

The North Coast of Cornwall

ending in steep slopes down to rugged cliff and shore scenery, with high headlands and deep valleys (cleaves). The incised Yeolmouth valley has a waterfall pouring  down into a shingle cove, and near Morwenstow, another deep V-shaped valley truncated by cliff recession ends in a waterfall above St. Morwenna’s Well. There is a recent major landslide south of Higher Sharpnose Point (>Fig. 7.9.1.1). To the south a deeply incised valley opens to Stanbury Mouth, which has a cliff waterfall and a beach of sandy and grey gravels between rocky ribs. South from Warren Point the proportion of rocky ribs and shores diminishes as sandy beaches lengthen towards Bude. At Earthquake a major landslide has occurred down a sandstone bedding plane, producing a basal fan of talus, and large sandstone ribs run out into sea. Bude is built behind a sandy bay at the mouth of the River Neet valley. The beach sands are calcareous, including shelly debris washed in from the sea floor and quartzose sand derived from the erosion of weathering sandstones exposed in the nearby cliffs and shore platforms. Large quantities of shellysand were extracted from this beach in the eighteenth and nineteenth centuries, resulting in the reduction of Summerleaze Beach. Whale Back (folded sandstones) and Saddle Rock are en echelon pericline folds, which look like inverted boats when seen from the top of the cliff. To the south, marine erosion has cut steep to vertical cliffs across the sharply folded sandstones, mudstones and shales, fronted by a gently sloping intertidal shore platform cut across these formations, ⊡⊡ Fig. 7.9.1.1 The steep Atlantic coast south of Higher Sharpnose Point. (Courtesy Geostudies.)

7. 9.    1

truncating the anticlines and synclines that run at right angles to the coastline. Variations in the resistance of the outcropping strata are expressed as ribs and clefts in the cliffs and ridges and furrows across the shore platform (>Fig. 7.9.1.2). The promontory of Lower Longbeak, cut in massive mudstones, marks the southern end of the high cliffs. Widemouth Bay is backed by low cliffs cut in Head Deposits. The shore platforms show structural features developed across strongly-folded sandstone and shale formations in which the axes of folding run out westward across the shore. The sandstones and shales of the Bude For­mation give place to dark shales and thin grey sandstones of the Crackington Formation at Wanson Mouth. This formation has some outcrops mainly of sandstone, forming bold cliffs and resistant outcrops on the shore, and others (generally narrow) mainly of shale, excavated to form coves, inlets and caves. It dominates the coast south to Boscastle. Much of the coast is steep, with slope-overwall profiles cut into fault-bounded outcrops of various Upper Devonian and Lower Carboniferous formations. At Crackington Haven two deep valleys meet behind a sandy cove, backed by grey shingle and low cliffs in Head, between cliffs and steep slopes. As the coast runs out to Cambeak the cliffs are cut across strongly folded sandstones and shales. Cambeak (98 m) is a sloping sandstone headland with tumbled grassy slopes in Head and cliffs that show swirling folds and recumbent folding in the Crackington Formation. South of Fire Beacon Point there are landslides on steep coastal slopes to Beeny Cliff

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⊡⊡ Fig. 7.9.1.2 Shore platform cut across steeply dipping sandstones and shales of the Bude Formation at Lower Longbeak, south of Bude. (Courtesy Geostudies.)

(>Fig. 7.9.1.3), where convex coastal slopes intersect northward-dipping sandstones and descend to rocky cliffs broken by large caves. South towards Boscastle high ground recedes from the coast and the Trevena Platform (90–130 m) dominates the hinterland, ending in bevelled slope-over-wall profiles. Boscastle Harbour is the drowned mouth of the incised Valency valley, bordered by steep craggy slopes in the strongly folded slates and sandstones of the Crackington Formation. The coast between Boscastle and Bossiney Haven has many inlets at the mouths of deeply-incised valleys where streams flow steeply down to the coast. South of Trevalga Head a deep cove has been cut in rocks that dip steeply northeastward. Cliffs in green Tredorn Slates are cut by numerous low-angle faults and some steeper, almost vertical fractures. The Tintagel Volcanic Formation is prominent in the high craggy Trevalga Cliffs and Rocky Valley is a long and deep valley cut in Delabole Slates that descends to the coast. There is a sandy beach at low tide in Bossiney Haven. In the Tintagel district cliff outlines have been influenced by dissection along rectilinear joint patterns in Devonian and Carboniferous slates and volcanic rocks (Wilson 1952). Upper Devonian slates cap Tintagel Island, thrust up over Carboniferous Tintagel Volcanics. South of Tintagel cliffs, Devonian and Carboniferous rocks extend along Port Isaac Bay, past Port Quin, a curving drowned valley-mouth inlet, and out to the bold promontory between Rumps Point and Pentire Point. This borders the wide Camel estuary, which has extensive

inwashed sands. On its northern shore is the Trebetherick raised beach (>Fig. 7.9.1.4), on an emerged shore platform rising to about 3 m above high tide level, cut across the dipping slates (Arkell 1943). This dates from a phase when the sea stood higher than it does now, and shows features that formed during an ensuing phase of cold (periglacial) climate and much lower sea level. The south coast of the Camel estuary shows a succession of promontories on hard igneous rocks, notably greenstones, lava and tuff, and intervening bays cut out in the softer slates. Between Stepper Point and Trevose Head is a broad bay backed by low cliffs and coastal slopes of subdued outline, cut in weathered Middle to Upper Devonian slates. South of Porthmissen Bridge the cliff crest curves round the head of a deep bouldery inlet, the southern side of which is Marble Cliff, over 30 m high, cut in about 140 thin bands of alternating gently-dipping grey limestone, dark shale and striped grey slate. This is the lower part of an overfold, so that the strata are inverted, the youngest beds being near the bottom. Towards Trevose Head the steep coastal slope is cut in massive dark greenstone. Trevose Head and Dinas Head are windswept headlands, and Constantine Bay is a gentlycurving west-facing bay cut in shales, and occupied by a wide surf beach of shelly sand, backed by dunes held in place by marram grass. To the south the cliffs have been much dissected by marine erosion, isolating numerous stacks and small islands. At Bedruthan Steps (>Fig. 7.9.1.5) the cliffs in Devonian slates are almost vertical, and there are several high stacks including the pinnacle of Queen Bess Rock.

The North Coast of Cornwall

7. 9.    1

⊡⊡ Fig. 7.9.1.3 The slope-over-wall profile of Beeny Cliff, north of Boscastle, with High Cliff in the ­background. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.4 The emerged beach at Trebetherick Point, on the eastern shores of the Camel estuary, resting on a bench cut into Upper Devonian slates and overlain by laminated dune calcarenite. Brae Hill in the distance is an outcrop of greenstone. (Courtesy Geostudies.)

On Trenance Point there are sandstones and shales with hard quartzite forming protrusions on the bold cliffs. Mawgan Porth is a sanded inlet at the mouth of the Vale of Mawgan. High cliffs border Beacon Cove, and at Strasse Cliff the Watergate Bay surf beach begins. The coastline of Watergate Bay is cut across an E–W  anticline, where the Meadfoot Beds overlie Dart­­ mouth Beds rising through the cliffs. It has a sandy beach 2.5 miles long in front of vertical cliffs at the edge of the St. Mawgan plateau. Ocean swell washes the base

of  the cliffs at high tide, and rip currents are powerful (>Fig. 7.9.1.6). At Whipsiderry the coast turns westward, and the Meadfoot Beds continue along the north coast of Porth Island, where the southward dip produces an escarpment cliff. The great cave known as the Banqueting Hall was formerly about 60 m long, up to 20 m wide and nearly 20 m high, with two entrances from the sea and a third through the roof. Unfortunately it became unsafe and the local council felt obliged to destroy it with explosives

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⊡⊡ Fig. 7.9.1.5 The cliffs, stacks and low-tide sandy beach near Bedruthan Steps, looking towards Park Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.6 An air view of Watergate Bay, with strong Atlantic swell at high tide. The receding cliffs retain no remnants of the ­periglaciated slope seen on less exposed parts of the Cornish coast. (Courtesy Geostudies.)

in 1987. Porth Island runs out to Trevelgue Head (>Fig. 7.9.1.7), and is bordered by the long inlet at Porth, where sand is exposed at low tide. The Newquay coast has cliffs fronted by flat brown sands exposed at low tide. Towan Head is a rounded promontory of hard calcareous slates and thin limestones running out to the northwest. A low cliff exposes Holocene dune sand over gravelly Head up to 3 m of thinly bedded Pleistocene calcareous sandrock (a beach or dune calcarenite), with an emerged beach resting upon grey Mead­foot Slates.

Fistral Bay is cut into rapidly weathering shales, and has a sandy beach washed by surf from Atlantic swell, backed by grassy dunes. Pentire Point East and West are headlands in hard slates and grits, divided by The Gannel, another sanded inlet. Porth Joke, to the west, is similar. South of Kelsey Head is Holywell Bay, several caves below calving cliffs capped by Head and cliff-top dunes, and a wide beach backed by grassy dunes. Cliffed headlands run out to Penhale Point and Ligger Point, then Perran Sands, backed by high dune topography that extends just over 2 miles south to Perranporth.

The North Coast of Cornwall

7. 9.    1

⊡⊡ Fig. 7.9.1.7 An air view of Trevelgue Head and Porth at high tide, with the cove at Lusty Glaze on the right. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.8 The stack at Tobban Horse, a stack south of Porthtowan, with the remains of the coastal slope on its seaward side. Erosion during the past 6,000 years has excavated a cove here, cutting the cliffs back by up to 200 m. (Courtesy Geostudies.)

Droskyn Point was much modified by mining and slate quarrying in the eighteenth and nineteenth centuries, and a rugged slope-over-wall coast runs out to Cligga Head and on to St. Agnes Head and Navax Point. The proportion of slope to wall varies according to exposure to Atlantic waves. South of Porthtowan is Tobban Horse, a large stack with a grassy Head-mantled slope on its seaward side, from which the extent of cliff recession can be estimated (>Fig. 7.9.1.8).

At Portreath is a valley-mouth beach of calcareous sand mixed with sand and gravel from mining waste. To the SW, even-crested cliffs about 80 m high truncate the coastal plateau (Trevena Platform) and a slope-over wall coast runs out to Naxax Point. Godrevy Point marks the beginning of St. Ives Bay, and the cliffs cut in slates at Godrevy include an emerged Pleistocene beach and stratified dune calcarenite beneath Head deposits, as well as a fronting shore platform (>Fig. 7.9.1.9).

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⊡⊡ Fig. 7.9.1.9 Cliffs cut in Head deposits back the broad shore platform at Godrevy, on the north coast of St. Ives Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.10 Granite tor above Gamper, Land’s End, looking along the slope-over-wall coast to Pedn-men-du, with the stack known as the Irish Lady (arrowed). (Courtesy Geostudies.)

Red River comes down to the sea from the mines of the Camborne district. It was formerly stained by oxidised waste from the mining area, but this pollution has been halted. The beach sands of St. Ives Bay contain tin and other heavy minerals supplied by streams draining from mineral-rich hinterlands and eroded from coastal outcrops. They are backed by dunes that have been drifting inland since mediaeval times, burying farmland and early settlements.

The River Hayle flows through a broad estuary to reach the sea through an outlet between the coastal dunes behind Porth Kidney Sands. To the west cliffs and bluffs extend to St. Ives. The north coast of the Land’s End (Penwith) Peninsula has sectors of granite coast between extensive fringes of metamorphic rock and greenstone intrusions. It is a high, steep and cliffy coast facing stormy seas, and its outlines have been strongly influenced by patterns of jointing in the rocks. Zennor Head is a promontory of

The South Coast of Cornwall

greenstone (dolerite) with craggy slopes at the edge of a plateau about 90 m above sea level. Typical granite cliffs, dissected along vertical and almost horizontal joint planes, run from Gurnards Head to Portheras Cove, where there is a junction between the granite and Mylor Slates. The cliffs are incised by deep zawns, steep-sided inlets cut out along joints or fault planes, some of which have been excavated or enlarged by mining along tin and copper lodes in  the coastal rock outcrops. Beyond Cape Cornwall, in Priest’s Cove the metamorphic slates give place to granite, which dominates the coast from here to Lands End. At Porth Nanven there is an emerged beach of granite boulders, overlain by gravelly Head deposits. Whitesand Bay has an upper beach of coarse shelly brown sand behind a wide almost flat area of finer softer mica-rich sand exposed at low tide; to the south, Sennen, Cove has a beach of cobbles and boulders.

7. 9.       2            

The coast curves south, past the cliffed headland of Pedn-men-du and the Irish Lady, a stack with a conical hat. Grassy slopes on Head deposits have granite tors (>Fig. 7.9.1.10) and are trenched by large cliff cauldrons, possibly collapsed caves, as the coast runs out to the castellated granite cliffs of Dr. Syntax’s Head, the promontory that declines to Land’s End.

References Arkell WJ (1943) The Pleistocene rocks at Trebetherick Point, Cornwall: their interpretation and correlation. Proc Geol Assoc 54:141–170 Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Wilson G (1952) The influence of rock structures on coast-line and cliff development around Tintagel. North Cornwall. Proc Geol Assoc 63:20–49

7.9.2  The South Coast of Cornwall 1.  Introduction The south coast of Cornwall is more subdued and indented than the north coast, and receives waves generated by winds across the English Channel as well as south-westerly ocean swell from the Atlantic (Bird 1998). Wave action is generally light to moderate, with large storm waves arriving only occasionally. Mean spring tide range averages 4.7–4.8 m. Slope-over-wall profiles have steep vegetated slopes descending to rocky cliffs that are usually less than 10 m high (Arber 1949). The proportion of rocky cliff to coastal slope is less than on much of the north coast, but on the western side of the Lizard Peninsula, exposed to Atlan­­ tic  swell and storm waves, vertical cliffs rise locally to 50–70 m. Head­lands and promontories correspond with higher ground or harder outcrops, the intervening bays and coves occupying lower ground or softer outcrops. The coastline runs in along the sides of several large branching inlets known as rias, formed where the sea has invaded the lower parts of river valleys, notably the Tamar, the Helston and Fowey Rivers, and Carrick Roads. From time to time large waves break upon the southern shores of Cornwall. Some are generated by storms, but there are also records of sudden bursts of wave activity in calm weather, particularly in Mounts Bay and along the south coast to Whitsand Bay. Such waves were observed along the south coast on 1 November 1755, 4 h and 15 min after the catastrophic Lisbon earthquake generated a tsunami transmitted northward in the West Atlantic: large

granite boulders were thrown up well above high tide level at Lamorna Cove on the Land’s End Peninsula. Mean spring tide ranges show little variation: at Newlyn 4.8 m, Porthleven, Lizard Point, Coverack, Helford River mouth, Falmouth and Mevagissey 4.7 m, Fowey and Looe 4.8 m and Whitsand Bay 4.5 m.

2.  The Coastline The columnar granite cliffs of Land’s End are dissected along horizontal and vertical joint planes, and interrupted by grassy chutes of Head deposits. Many granite blocks have fallen to the shore. There are valley-mouth coves with little sandy beaches. In places there is a narrow cliff-base ramp, cut out along joint planes inclined gently seaward. The sea is green over a sandy floor, but darker over weedy areas amid inshore boulders. In Pendower Cove there are caves cut out along joint planes (>Fig. 7.9.2.1). The coast to Gwennap Head is bold, ­facing SW into an often stormy sea; even in calm weather Atlantic swell breaks as a heavy surf over reefs and shore rocks. On Gwennap Head rectilinear jointing is ­prominent, and the coast turns eastward. The southern coast of the Land’s End peninsula is almost entirely medium to coarse grained granite, with associated igneous rocks, such as the greenstones at Tater-du. Porthcurno Cove, at the mouth of a steep-sided valley, has a sandy beach backed by dunes, and to the east the coast curves out to the promontory of

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The South Coast of Cornwall

greenstone (dolerite) with craggy slopes at the edge of a plateau about 90 m above sea level. Typical granite cliffs, dissected along vertical and almost horizontal joint planes, run from Gurnards Head to Portheras Cove, where there is a junction between the granite and Mylor Slates. The cliffs are incised by deep zawns, steep-sided inlets cut out along joints or fault planes, some of which have been excavated or enlarged by mining along tin and copper lodes in  the coastal rock outcrops. Beyond Cape Cornwall, in Priest’s Cove the metamorphic slates give place to granite, which dominates the coast from here to Lands End. At Porth Nanven there is an emerged beach of granite boulders, overlain by gravelly Head deposits. Whitesand Bay has an upper beach of coarse shelly brown sand behind a wide almost flat area of finer softer mica-rich sand exposed at low tide; to the south, Sennen, Cove has a beach of cobbles and boulders.

7. 9.       2            

The coast curves south, past the cliffed headland of Pedn-men-du and the Irish Lady, a stack with a conical hat. Grassy slopes on Head deposits have granite tors (>Fig. 7.9.1.10) and are trenched by large cliff cauldrons, possibly collapsed caves, as the coast runs out to the castellated granite cliffs of Dr. Syntax’s Head, the promontory that declines to Land’s End.

References Arkell WJ (1943) The Pleistocene rocks at Trebetherick Point, Cornwall: their interpretation and correlation. Proc Geol Assoc 54:141–170 Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Wilson G (1952) The influence of rock structures on coast-line and cliff development around Tintagel. North Cornwall. Proc Geol Assoc 63:20–49

7.9.2  The South Coast of Cornwall 1.  Introduction The south coast of Cornwall is more subdued and indented than the north coast, and receives waves generated by winds across the English Channel as well as south-westerly ocean swell from the Atlantic (Bird 1998). Wave action is generally light to moderate, with large storm waves arriving only occasionally. Mean spring tide range averages 4.7–4.8 m. Slope-over-wall profiles have steep vegetated slopes descending to rocky cliffs that are usually less than 10 m high (Arber 1949). The proportion of rocky cliff to coastal slope is less than on much of the north coast, but on the western side of the Lizard Peninsula, exposed to Atlan­­ tic  swell and storm waves, vertical cliffs rise locally to 50–70 m. Head­lands and promontories correspond with higher ground or harder outcrops, the intervening bays and coves occupying lower ground or softer outcrops. The coastline runs in along the sides of several large branching inlets known as rias, formed where the sea has invaded the lower parts of river valleys, notably the Tamar, the Helston and Fowey Rivers, and Carrick Roads. From time to time large waves break upon the southern shores of Cornwall. Some are generated by storms, but there are also records of sudden bursts of wave activity in calm weather, particularly in Mounts Bay and along the south coast to Whitsand Bay. Such waves were observed along the south coast on 1 November 1755, 4 h and 15 min after the catastrophic Lisbon earthquake generated a tsunami transmitted northward in the West Atlantic: large

granite boulders were thrown up well above high tide level at Lamorna Cove on the Land’s End Peninsula. Mean spring tide ranges show little variation: at Newlyn 4.8 m, Porthleven, Lizard Point, Coverack, Helford River mouth, Falmouth and Mevagissey 4.7 m, Fowey and Looe 4.8 m and Whitsand Bay 4.5 m.

2.  The Coastline The columnar granite cliffs of Land’s End are dissected along horizontal and vertical joint planes, and interrupted by grassy chutes of Head deposits. Many granite blocks have fallen to the shore. There are valley-mouth coves with little sandy beaches. In places there is a narrow cliff-base ramp, cut out along joint planes inclined gently seaward. The sea is green over a sandy floor, but darker over weedy areas amid inshore boulders. In Pendower Cove there are caves cut out along joint planes (>Fig. 7.9.2.1). The coast to Gwennap Head is bold, ­facing SW into an often stormy sea; even in calm weather Atlantic swell breaks as a heavy surf over reefs and shore rocks. On Gwennap Head rectilinear jointing is ­prominent, and the coast turns eastward. The southern coast of the Land’s End peninsula is almost entirely medium to coarse grained granite, with associated igneous rocks, such as the greenstones at Tater-du. Porthcurno Cove, at the mouth of a steep-sided valley, has a sandy beach backed by dunes, and to the east the coast curves out to the promontory of

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⊡⊡ Fig. 7.9.2.1 Caves on a granite coast in Pendower Cove, near Land’s End. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.2 Logan Rock (L), perched on a cliff-top tor near Treen, on the south coast of the Land’s End Peninsula. The coarsegrained granite is divided by three intersecting joints, two almost vertical planes and one roughly horizontal and curving. (Courtesy Geostudies.)

Horrace, on which is an Iron Age fort and Logan Rock, a 65 tonne rocking boulder (> Fig. 7.9.2.2). The granite cliffs run eastward to Lamorna Cove, where the valley-side slopes are strewn with granite boulders. At Mousehole breakwaters enclose a former cove as a harbour, dry and sandy at low tide. Just south is the ­junction between the granite and the Mylor Slates, here metamorphosed into hard green splintery rock. Low cliffs in these slates extend northward past Penlee Point, an outcrop of greenstone. On South Pier at Newlyn is the tidal observatory, where the

rise and fall of the tide is recorded. There is evidence that mean sea level has been rising here in recent decades. From Newlyn the coast curves NE and is largely artificial, round to Penzance. Mounts Bay has brown muddy sand with outlying rocky reefs and a submerged forest. Behind the sandy beach towards Marazion, grassy dunes are backed by reedswamp and scrub (Marazion Marsh) in a former valley-mouth lagoon. The Causeway (> Fig. 7.9.2.3) that leads out at low tide to St. Michael’s Mount is an ebb road, passable for an hour or two before and after low

The South Coast of Cornwall

tide: the duration of emergence has diminished as the result of sea level rise (as indicated on the Newlyn tide gauge). St. Michael’s Mount is a high island of granite with an outcrop of Mylor Slates on its eastern shore. Granite ledges, sloping gently SE along joint planes, rise along the  southern shore, and granite boulders on the NE shore  were emplaced by storm surges or tsunamis. Low vertical cliffs in Head deposits and reefs and headlands of greenstone extend east from Marazion to Cudden Point. Prah Sands occupy an embayment backed by cliffs in Head ⊡⊡ Fig. 7.9.2.3 The causeway (ebb road) to St. Michael’s Mount crosses a rocky, sandy and bouldery isthmus exposed at low tide. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.4 A fault exposed in the cliffs behind Rinsey Cove, with a cave excavated along the fault plane. (Courtesy Geostudies.)

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and Mylor Slates and in Rinsey Cove the cliff exposes the contact between the Mylor Slates and the Tregonning granite (> Fig. 7.9.2.4). On the shore platform fronting cliffs west of Porthleven is Giant’s Rock (> Fig. 7.9.2.5), a large rounded block of smooth pink microcline gneiss, a rock type unknown in Britain is seen and probably brought here in an iceberg that floated on to the shore and melted. Porthleven Harbour occupies a valley-mouth inlet and to the east the beach consists of fine flinty shingle and

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coarse sand which extends SE as Loe Bar, a barrier ­enclosing Loe Pool in the drowned valley mouths of the River Cober and Carminowe Creek. The beach continues SE to Gunwalloe, where a Late Pleistocene emerged beach can be seen. To the south cliffs extend past headlands and coves to Church Cove, where a headland is capped by a stabilised dune. There is a contrast between the rocks on either side of the sandy beach in Polurrian Cove: Devonian Portscatho slates on the north and hornblende schists to the south.

The schists run round to Mullion Cove, where there are dykes and lenses of serpentine, which dominates the high cliffs south past Kynance Cove towards Lizard Head. Steep and rugged cliffs in mica schists extend along the coast to Lizard Point. The cliffs decline to Housel Bay from Lizard Point where The Lion’s Den (> Fig. 7.9.2.6) is an almost circular hole nearly 50 m in diameter at the top of a shaft over 70  m deep, formed when the roof of a cave collapsed in 1847. ⊡⊡ Fig. 7.9.2.5 Giant’s Rock, a large boulder of gneiss thought to have arrived on the shore west of Porthleven as an ice-borne erratic in Pleistocene times. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.6 Slope-over-wall cliffs on the western side of Housel Bay, Lizard Peninsula, showing cliff cauldrons and the Lion’s Den (L), a round hole formed by the collapse of a cave roof. (Courtesy Geostudies.)

The South Coast of Cornwall

Steep rugged cliffs with crags and tors run out on Pen Olver, where the coast turns northward. Steep grassy slopes descend to low cliffs in hornblende schist along this more sheltered coast to Cadgwith and Kennack Sands, and then higher cliffs run eastward to Black Head. At Coverack the shore is irregular, with serpentine outcrops cut by dykes of coarse crystalline gabbro, black epidiorite and other dark intrusive rocks. To the north on Lowland Point a wide terrace is cut by low cliffs in Head deposits. Godrevy Cove, Porthoustock and Porthallow are bays on a cliffed coast cut in gabbro, with beaches ­augmented by gravelly quarry waste. Beyond Nare Point the cliffs recede and decline beside drowned valleys at Gillan Harbour and the Helford River. The rounded peninsula of Rosemullion Head is cut into the dark contorted slates and sandstones of the Portscatho Beds, and to the north the cliffs rise to a succession of rocky promontories and small sandy coves, with ridges of slate running out eastward across the shore. Maen­porth has a sandy beach behind a wide valley-mouth cove, and near Newporth Head there are prominent cliff-base ledges bearing remnants of a gravelly raised beach (> Fig. 7.9.2.7). The coast at Falmouth consisted of receding cliffs of Head deposits stabilised by construction of a sea wall. This is fronted by a gravelly beach and a shore platform that is being cut into a similar Late Pleistocene shore platform, which stands 2–3 m higher. Pendennis Point is fringed by vegetated slopes that descend to cliffs of rubbly Head above steep rocky shores. This marks the entrance to

⊡⊡ Fig. 7.9.2.7 Sunny Cove, near Falmouth, showing remnants of the Late Pleistocene emerged shore platform (P) between segments of modern shore platform backed by cliffs. (Courtesy Geostudies.)

7. 9.       2            

Carrick Roads, a large branching drowned valley system. The River Fall, which opens into an upper branch of Carrick Roads, built a delta of china clay brought down from quarries on Hensbarrow Down (Ranwell 1974), but when this clay pollution was halted the delta began to erode (> Fig. 7.9.2.8). St. Mawes stands behind a south-facing bay on a peninsula north of the Porthcuel River. The shores are rocky, cut into outcrops of the soft slates of the Portscatho Formation. At St. Anthony Head, on the eastern side of the entrance to Carrick Roads, the coast swings round to the NE. The cliffs and shore rocks are interrupted by Towan Beach, mainly sandy with some gravel, in front of low cliffs of Head and a section of coastal dunes. North of Portscatho the cliffs follow smooth cleavage planes and the coastline runs along the strike of the shore outcrops. The shore becomes sandy as the coast curves behind Gerrans Bay, and exposure to SW waves increases. There are emerged shore platforms and segments of raised beach, then the coast curves south towards Nare Head. From Nare Head to Portholland the cliffs of Veryan Bay are cut mainly in the Roseland Breccia, a jumbled mass of brecciated slates and sandstones. There are sand and gravel beaches at Caerhayes and Hemmick, and the Dodman is a cliffed promontory of phyllites - the cliffs higher on the more exposed western coast (> Fig. 7.9.2.9). Vault Beach (> Fig. 7.9.2.10), to the east, is sandy, with patches of white quartz grit and gravel and grey phyllite

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The South Coast of Cornwall

⊡⊡ Fig. 7.9.2.8 The Fal delta, Carrick Roads, showing eroded micro-cliffs at the edge of a salt marsh and a wave-cut ramp in white china clay deposits. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.9 Slope-over-wall coast on the eastern side of Dodman Point. (Courtesy Geostudies.)

pebbles and granules, the sand having come mainly from the Head deposits exposed at the base of the coastal slope. The east-facing coast has a succession of valley-mouth coves at Gorran Haven, Portmellon and Mevagissey, separated by small promontories. Pentewan Beach (> Fig. 7.9.2.11) is wider, formed mainly by sand and brought down by the St. Austell River from a mining hinterland. To the north is Black Head, a rounded promontory, with gorse heath on slopes that decline to cliffs and shore outcrops of massive, hard greenstone. At Charlestown the harbour was built in a valley-mouth cove and the beach of sand and shingle includes pebbles of ballast from the ships that came in to collect china clay. The coast turns eastward past Carlyon Bay (>  Fig.  7.9.2.12),  where a straight and wide beach was ­augmented by the deposition of mining waste (sand and gravel) conveyed by diverting a stream through a tunnel cut in 1842 under the coastal ridge (Everard 1960). This supply has ceased, and the beach is diminishing. Beach accretion in Par Bay accelerated as the result of  the delivery of waste from tin and copper mining in the hinterland by rivers. On the Gribbin Head promontory the coastal slopes descend to cliffs and rocky shores cut in  slates of the Meadfoot Beds, and these extend eastward  past the drowned valley at Fowey to Lantivet Bay. Meadfoot Beds alternate with Dartmouth Slates, the lowest of the Devonian formations seen in Cornwall, as the coast intersects an ­anticlinal structure. The cliffs descend into the deep valley at Polperro, and Talland Bay is at the mouth of two ­converging valleys. Behind Talland Bay the coastal slope on Head deposits descends to low cliffs in

The South Coast of Cornwall

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⊡⊡ Fig. 7.9.2.10 South along Vault Beach, which contains white quartzite pebbles, derived from quartz veins in the Dodman phyllites. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.11 Pentewan Beach, built up by the accretion of sand and gravel waste produced by mining and quarrying in the hinterland, and washed down the St. Austell River. (Courtesy Geostudies.)

the  pink,  purple and grey quartz-veined Dartmouth Slates (> Fig. 7.9.2.13). Looe River flows down a long narrow valley which was invaded by the sea to form a straight inlet, now largely filled with alluvial sediment. To the east the cliffs are mostly cut in dark grey Meadfoot Beds, and remains of a submerged forest are sometimes uncovered at low tide on the Millan­draeth shore. Towards Seaton, steep cliffs of crumbling shales and slates have been cut back quickly by the sea, and there are sectors of arcuate slumping. Between

Seaton and Downderry, cliffs of earthy Head were formerly eroding and slumping because of frequent waterlogging, but the landslides have been stabilised and landscaped. Behind Long Beach there are many small caves angled into the cliffs with sloping roofs on bedding planes (> Fig. 7.9.2.14). Whitsand Bay has a 5 mile gently curving coastline that faces SW, with scrubby coastal slopes descending to low cliffs and protruding ribs of silvery grey Dartmouth Slates (> Fig. 7.9.2.15).

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The South Coast of Cornwall

⊡⊡ Fig. 7.9.2.12 The outlet from the tunnel (arrowed) through which mining waste was delivered to form a beach in Carlyon Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.13 The shore at Talland Bay, showing ridges and furrows in the pink and grey Dartmouth Beds, which strike east-west and dip steeply northward. (Courtesy Geostudies.)

Rame Head projects southward on an outcrop of hard sandstone, and at Penlee Point the coastline swings sharply northward, and becomes sheltered from the strong SW waves. Dartmouth Slates outcrop in the low cliffs and rocky shores NW to Cawsand, on a part of the coast that has forested slopes descending almost to high tide level. These continue round Mount Edgcumbe into the Tamar estuary. This is another drowned valley (ria)

system, branching into St. John’s Lake, the St. Germans estuary, the Tavy and the Plym. It is a “low wave energy coast” because wave action across the estuary is weak, even at high tide, and ineffective as the tide falls. Tidal currents shape intertidal mudbanks, which are typically convex in profile, curving down to the channels that ­persist at low tide. They are backed by segments of salt marsh, generally cliffed along the outer edges. The

The South Coast of Cornwall

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⊡⊡ Fig. 7.9.2.14 Eastward dipping Meadfoot Beds in the cliffs behind Long Beach, Tregantle, showing a cave cut out along bedding planes. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.15 The coast of Whitsand Bay at Sharrow Point, showing ribs of Dartmouth Slate, with east-facing slopes on cleavage planes. (Courtesy Geostudies.)

boundary between Cornwall and Devon follows the Tamar River upstream.

References Arber M (1949) Cliff profiles of Devon and Cornwall. Geogr J 114:191–197

Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Everard CE (1960) Mining and shoreline evolution near St. Austell, Cornwall. Trans Roy Geol Soc Cornwall 19:199–219 Ranwell DS (1974) The salt marsh to tidal woodland transition. Hydrobiol Bull 8:139–151

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Isles of Scilly

7.10  Isles of Scilly 1.  Introduction The Isles of Scilly, about 48 km SW of Land’s End, are granitic islands produced by partial marine submergence of a granite landscape much like that of the Land’s End Peninsula. The islands consist of a dissected plateau, the higher summits rising between 45 and 49 m (possibly an uplifted marine planation surface of Pliocene age), and there are subdued terraces at lower levels, notably 30–40 m above sea level, representing Pleistocene phases of marine erosion at intermediate levels. About 9,000 years ago, when sea level was 10 m lower than at present, the Isles of Scilly consisted of one large granite island, but the continuing rise of sea level submerged the low-lying parts and left the hills as numerous large and small rocky islands, with associated gravel and boulder beaches, around the inner Sounds. There are about 135 of these islands, including large rocky islets, and Broad Sound is a submerged valley or down-faulted corridor that opens to the SW. Bishop Rock with its prominent lighthouse stands far out to the south–west, beyond numerous irregular rocky skerries (Barrow and Flett 1906). The outer coasts of the islands, especially to the north and west, are exposed to frequent Atlantic storms,

and present bare rugged headlands, islands and stacks. There are many granite tors and buttresses, particularly along the coastal slopes. The granite is divided by joints, mainly vertical or sub-horizontal, but sometimes inclined, and these have influenced the shape of the rocky islands, headlands, cliffs and reefs. Shore platforms are rare, but occur on flat or gently-inclined joint planes. Sandy beaches generally occupy coves and small inlets between granitic headlands. Much of the sand is quartzose, formed by the weathering of granite and swept up from the surrounding shallow sea floor by wave action. Some sand and gravel, including shelly deposits, still arrive on these beaches, but in recent decades beach erosion has become prevalent. Some of the beaches have been shaped into tombolos, linking pairs of islands, such as the sandy barrier on which Hugh Town, on St. Mary’s, is built, linking two former granite islands, and the sandy isthmus on Samson. There are similar intertidal bars of sand and gravel, known as swashways. There are cliff exposures of periglacial Head deposits, which contain rock fragments of local origin in a matrix of sandy clay, and on White Island, St. Martin’s and locally elsewhere there are patches of glacial drift (Mitchell and Orme 1967; Scourse 1991). The glacial

⊡⊡ Fig. 7.10.1 Joint planes in the granite exposed on Inner Head, south–east of Hugh Town. (Courtesy Geostudies.)

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drift on the northern fringes of the Isles of Scilly, ­somewhat south of the general limit of Pleistocene glaciation in Britain, has been attributed to a lobe of Irish Sea ice that spread southward across an emerged sea floor and extended to this area ­during the Anglian (Mindel) Glaciation, about 200,000 years ago.

varying proportions of carbonate derived from shelly organisms. Intervening cliffed sectors have ­slope-over-wall profiles (> Fig. 7.10.3). North–east of St. Mary’s, across Crow Sound, the Eastern Isles are a group of smaller islands arranged in NW–SE groups.

2.  St. Mary’s

3.  St. Martin’s

St. Mary’s, the largest of the Isles of Scilly, consists of a western promontory, Garrison Hill, a former small island attached to the main island by a tombolo (> Fig. 7.10.1). The outlines of the island are strongly influenced by joint patterns in the granite, which also influence the shape of cliffs and tors (> Fig. 7.10.1). The south and east coasts of St. Mary’s show a series of headlands and coves excavated along joint planes, mainly NW–SE. On sheltered parts of the coast, as in Carn Leh Cove and Porth Hellick, the slope in periglacial Head descends almost to high tide level, and there are exposures in low cliffs. In Porth Hellick, the Head deposits are exposed in low cliffs and small caves have been excavated by marine erosion (> Fig. 7.10.2). Beaches on the coast of St. Mary’s are confined to bays and coves, and consist mainly of quartz and felspar derived from weathering and erosion of granite in the bordering cliffs and rocky shores and on the sea floor, together with

St. Martin’s is an elongated island rising to a plateau cut across the granite, probably by marine erosion whenthe sea stood at a higher level. White Island, linked to St. Martin’s by a natural causeway of gravel and boulders (> Fig. 7.10.4), is notable for its glacial drift deposits (till), ­indicating that Irish Sea ice at one stage extended to the northern part of the Isles of Scilly. The glacial drift is prominent in the central part of the island, notably along an incised cleft (>  Fig. 7.10.5). The archipelago of granitic islets, south of St. Martin’s, include boulder beaches, spits and tombolos.

⊡⊡ Fig. 7.10.2 Rounded caves (a metre high, a metre wide and about a metre inset) cut into Head deposits in a cliff bordering Porth Hellick, St. Mary’s, formed by abrasion as granite boulders were agitated by wave action at high tide. (Courtesy Geostudies.)

4.  Tresco Tresco is separated from St. Martin’s by shallow straits. The northern part of the island is a cliff-edged granite ­plateau, but the southern part consists of former small islands attached by sandy depositional features. The

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Isles of Scilly

⊡⊡ Fig. 7.10.3 The slope-over-wall profile on the NE coast of St. Mary’s, with low cliffs cut into the Head deposits that mantle the slope and a rocky shore developed on the underlying granite. (Courtesy Geostudies.)

⊡⊡ Fig. 7.10.4 The bouldery intertidal causeway linking White Island to St. Martin’s. Most of the boulders are of local granite, but there are cobbles and pebbles of flint, basalt and sandstone derived from the island’s glacial drift deposits. (Courtesy Geostudies.)

­ arrow strait between Tresco and Bryher contains relics of n prehistoric field systems and hut circles exposed at low tide. Grassy dunes occupy the southern part of Tresco and cap barriers, which enclose freshwater lakes, on either side of Tresco. West of Tresco, across the strait, is the island of Bryher, which has a sheltered southern coast with rocky shores and bay beaches.

5.  St. Agnes St. Agnes, in the SW of the Isles of Scilly, has rocky northern, western and southern coasts exposed to strong wave action and a more sheltered eastern coast, including the island of Gugh, attached by a sandy tombolo.

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⊡⊡ Fig. 7.10.5 A cliff on White Island exposing glacial drift containing angular fragments of local granite and a variety of erratics including sandstones, flint and volcanic rocks in a yellow-brown earthy matrix. The erratic stones have been derived from regions far to the north and carried here by ice. (Courtesy Geostudies.)

References Barrow G, Flett JS (1906) The Geology of the Isles of Scilly. Memoirs of the Geological Survey, Geological Survey Publication, HMSO, London

Mitchell GF, Orme AR (1967) The Pleistocene deposits of the Isles of Scilly. Quart J Geol Soc London 123:59–92 Scourse JD (1991) Glacial deposits of the Isles of Scilly. In: Ehlers J et al (ed)  Glacial deposits in Great Britain and Ireland. Rotterdam, the Netherlands, pp 291–300

7.11  South Devon The South Devon coast consists of a southern part which resembles the south coast of Cornwall, developed on overthrust Devonian formations that extend from the Tamar estuary across to Torbay, with some headlands of greenstone and older rock formations, notably the Palaeozoic and Pre-Cambrian schists of Start Point and Bolt Tail, which resemble those of the Lizard in Cornwall. The central section is dominated by red cliffs of Triassic sandstone and mudstone within Torbay and from Teignmouth round to Seaton, and the eastern part by the Cretaceous rocks at the head of Lyme Bay. Much of the coast is steep and cliffed. On the southern part there are slope-over-wall profiles similar to those of Cornwall, with vegetated slopes mantled by periglacial Head deposits and basal cliffs cut in hard rock formations, but the receding cliffs on Triassic sandstone are generally vertical (often with outlying stacks), and the coastal slopes in Cretaceous rocks, irregular as the result of landslides. The steep and cliffed coasts are interrupted by incised valleys that have been submerged by the sea to form rias, which have been in varying degree filled with sediment to produce intertidal sandy and muddy

areas, salt marshes and reedswamp, and alluvial valley floors. Wave action is often strong from the SW, and occasionally from the SE and east, particularly in Lyme Bay. Mean spring tide ranges are generally between 4 and 5 m. Beaches are generally sandy, with some gravel in rocky coves, but the shores of Start Bay are bordered by a gently curving beach of brown shingle, predominantly flint pebbles. This includes the barriers that enclose Slapton Ley and smaller lagoons at Beesands.

The South Devon Coast Plymouth Sound is a major ria at the mouth of the River Tamar and its tributaries, notably the Tavy, Lynher and Plym. As the tide falls, wide muddy smooth slopes are exposed, crossed by minor creeks. The muddy sediment has been derived largely from fine-grained material in the Head deposits on bordering slopes, exposed in low cliffs beside the estuary. The entrance to Plymouth Sound is sheltered by a long breakwater.

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⊡⊡ Fig. 7.10.5 A cliff on White Island exposing glacial drift containing angular fragments of local granite and a variety of erratics including sandstones, flint and volcanic rocks in a yellow-brown earthy matrix. The erratic stones have been derived from regions far to the north and carried here by ice. (Courtesy Geostudies.)

References Barrow G, Flett JS (1906) The Geology of the Isles of Scilly. Memoirs of the Geological Survey, Geological Survey Publication, HMSO, London

Mitchell GF, Orme AR (1967) The Pleistocene deposits of the Isles of Scilly. Quart J Geol Soc London 123:59–92 Scourse JD (1991) Glacial deposits of the Isles of Scilly. In: Ehlers J et al (ed)  Glacial deposits in Great Britain and Ireland. Rotterdam, the Netherlands, pp 291–300

7.11  South Devon The South Devon coast consists of a southern part which resembles the south coast of Cornwall, developed on overthrust Devonian formations that extend from the Tamar estuary across to Torbay, with some headlands of greenstone and older rock formations, notably the Palaeozoic and Pre-Cambrian schists of Start Point and Bolt Tail, which resemble those of the Lizard in Cornwall. The central section is dominated by red cliffs of Triassic sandstone and mudstone within Torbay and from Teignmouth round to Seaton, and the eastern part by the Cretaceous rocks at the head of Lyme Bay. Much of the coast is steep and cliffed. On the southern part there are slope-over-wall profiles similar to those of Cornwall, with vegetated slopes mantled by periglacial Head deposits and basal cliffs cut in hard rock formations, but the receding cliffs on Triassic sandstone are generally vertical (often with outlying stacks), and the coastal slopes in Cretaceous rocks, irregular as the result of landslides. The steep and cliffed coasts are interrupted by incised valleys that have been submerged by the sea to form rias, which have been in varying degree filled with sediment to produce intertidal sandy and muddy

areas, salt marshes and reedswamp, and alluvial valley floors. Wave action is often strong from the SW, and occasionally from the SE and east, particularly in Lyme Bay. Mean spring tide ranges are generally between 4 and 5 m. Beaches are generally sandy, with some gravel in rocky coves, but the shores of Start Bay are bordered by a gently curving beach of brown shingle, predominantly flint pebbles. This includes the barriers that enclose Slapton Ley and smaller lagoons at Beesands.

The South Devon Coast Plymouth Sound is a major ria at the mouth of the River Tamar and its tributaries, notably the Tavy, Lynher and Plym. As the tide falls, wide muddy smooth slopes are exposed, crossed by minor creeks. The muddy sediment has been derived largely from fine-grained material in the Head deposits on bordering slopes, exposed in low cliffs beside the estuary. The entrance to Plymouth Sound is sheltered by a long breakwater.

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South Devon

The coast at Plymouth is on Devonian limestone, which also outcrops on Mount Batten Point. From this Point coastal slopes rise to Staddon Point, where there are cliffs and basal rock ledges cut in the Lower Devonian Staddon Grits. To the south the underlying Meadfoot Beds emerge in slaty cliffs and irregular shore platforms in Bovisand Bay, and the Dartmouth Slates outcrop in cliffs and a wide shore platform extending round to Wembury Point. The slope-over-wall coastline on Dartmouth Slates then trends eastward into the Yealm estuary below Wembury. The Yealm estuary is another ria, joined by a tributary at Newton Ferrers. On the southern side, the Headmantled coastal slopes descend almost to high tide level, but basal cliffs in Dartmouth Slates develop past Mouthstone Point and increase to Gara Point. A series of coves and rocky promontories continues eastward to the mouth of the Erme estuary. The Erme estuary (> Fig. 7.11.1) has much inwashed sand. Only a small meandering river channel remains between sandflats at low tide, and upstream the sand gives place to muddy sediment. Beacon Point marks the beginning of Bigbury Bay, where slope-over-wall profiles border little promontories and sandy valley-mouth coves, as at Westcombe and Challaborough. Bigbury-on-Sea is linked to Burgh Island, a small island with Head-mantled slopes and rocky shores, by an intertidal sand bar. The cliffs are cut in Dartmouth Slates. Sand has been washed onshore from Bigbury Bay,

forming dunes at Bantham. There are offset spits at the mouth of the River Avon. A low slope-over-wall coast extends along the eastern side of Bigbury Bay, where the Dartmouth Slates disappear beneath Meadfoot Beds, and at Thurlestone there is a coastal outcrop of Triassic sandstone running out to Warren Point. Rocky ledges on Meadfoot Beds extend along the cliffs of Great Ledge to the two coves, Outer Hope and Inner Hope. The coast steepens from Hope Cove to Bolt Tail, the beg­­inning of a high section of steep coast on Palaeozo­ icand Pre-Cambrian mica schists that runs SE to Bolt Head, with bolder promontories on harder hornblende schist. There are tors on the coastal slope above the cove at Fernyhole Point. The Warren is a broad scrubby coastal slope with slumping in the Head deposits, and at Bolt Head the coast swings northward and steep, high cliffs extend past Sharp Tor and successive valley-mouth coves opening to sandy beaches, as at South Sands. Salcombe is built on a slope bordering a deep ria, extending inland as the Kingsbridge estuary. To the east the bold coast resumes, with steep, often irregular Head slopes and basal cliffs in hornblende schists out to Prawle Point. The coast then faces SE, and consists of Headaproned terraces backed by bluffs marking degraded Pleistocene cliffs and fronted by low cliffs and wide rocky shore platforms. Periglacial drift (Head) is exposed in bluffs behind Lannacombe Bay and in conical stacks.

⊡⊡ Fig. 7.11.1 The sanded estuary of the River Erme at low tide. (Courtesy Geostudies.)

South Devon

An incised valley opens on to Lannacombe Beach and to the east, the coastal terraces fade into a steepening slope, undercut by cliffs along to Start Point, a craggy headland of mica schist, with slopes with periglacial Head descending to cliffs which are higher on the more exposed southern side than on the northern flank facing Start Bay (> Fig. 7.11.2). The steeply sloping coast curves north from Start Point past the ruined village of Hallsands (> Fig. 7.11.3). In the nineteenth century the village was protected by a high, wide shingle beach, but cliff recession has been rapid here following the dredging of shingle between 1897 and 1902 to obtain gravel for the construction of Plymouth dockyard. The Start Bay beaches are relict, receiving no shingle under present conditions, so that the losses have not been compensated, and erosion is continuing. The floor of Start Bay is sandy, but there is shingle out  on Skerries Bank, which may be a submerged barrier  beach. On the coast the shingle beach is interrupted  by  Tinsey Head, but resumes to enclose a small lagoon at Beesands. To the north is a quarried headland of  greenstone, Limpet Rocks, with the village of Tor­­ cross beyond. A barrier beach up to 140 m wide, consisting of  fine flint shingle encloses a lagoon, Slapton Ley (> Fig. 7.11.4), which has a controlled outlet at the southern end. The Ley is almost fresh, receiving salt water and spray only during easterly storms, when waves may wash over the barrier beach and sweep shingle across into the lagoon. The barrier beach is thus migrating landward, narrowing the lagoon. There is no local source of flint ⊡⊡ Fig. 7.11.2 Coastal slopes on Head deposits, looking towards Start Point. (Courtesy Geostudies.)

7.11

pebbles, which may have been swept shoreward from the  sea floor  during  the Late Quaternary marine transgression. The northern part of Slapton Ley is occupied by reedswamp. At Pilchard Cove the Meadfoot Beds come to an end as the underlying Dartmouth Beds rise. There are rocky ribs and coves, the largest of which contains Blackpool Sands, and occasional headlands of greenstone, as on Matthews Point, on to the mouth of the River Dart. The long estuary of the River Dart is a ria cut into Dartmouth Slates, and is tidal to above Totnes. As it opens to the sea the northern side of the ria has steep wooded slopes extending seaward past Mill Bay Cove and New­ foundland Cove to headlands on igneous rock at Inner and Outer Froward Points. A steep coast with slope-­ over-wall profiles then continues northward with many small headlands and coves such as Scabbacombe Sands (> Fig. 7.11.5). Man Sands is a lagoon enclosed by a ­scrubcovered barrier that was breached during a flood in 2008. The coast then runs out to a small Devonian limestone promontory at Sharkham Point, and St Mary’s Bay is cut in slates. Berry Head rises to a plateau 60 m above sea level and its northern cliff has been cut back by quarrying. The coast runs westward along the northern side of the Brixham peninsula, with wooded bluffs on the southern coast of Tor Bay. Tor Bay has been excavated in soft Permo-Triassic sandstone between harder Devonian rocks to the south and north. Elbery Cove contains a beach of well-rounded cobbles, and Broad Sands are bordered by cliffs cut in grey

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South Devon

⊡⊡ Fig. 7.11.3 The ruined village at Hallsands. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.4 Slapton Ley and the barrier beach north of Torcross. (Courtesy Geostudies.)

Devonian limestone, passing to red Triassic sandstones on  the northern side. Roundham Head is fringed by cliffs of red sandstone which continue past Paignton and Hollicombe to Torquay. The Torquay peninsula runs out to Hope’s Nose, where the coastal slopes descend to low cliffs and shore platforms in Devonian limestone, with shingle beaches in coves, passing to greenstone at Black Head and Long Quarry Point. On the northern side of the peninsula Permian sandstones and breccias begin below Babbacombe, where

they are faulted against pale weathered Upper Devonian sandstones. Shingle beaches are derived from the breccias, and the shore is littered with fallen blocks of sandstone north of Oddicombe Beach. The Devonian limestone returns at Petit Tor Point and Watcombe Head has red sandstone cliffs. At the Ness south of Teignmouth, rounded Permian sandstone blocks have been washed out of a sandy matrix in the cliff. At Teignmouth the Teign estuary has wide muddy sandflats and some gravelly shoals at low tide, patches of salt

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⊡⊡ Fig. 7.11.5 Scabbacombe Sands occupy a cove on the northern side of Scabbacombe Head, near Dartmouth. (Courtesy Geostudies.)

marsh, minor beaches and forelands, and reed swamp upstream near Newton Abbot. A spit has grown south almost across the mouth of the Teign estuary. North of Teignmouth red sandstone and conglomerate cliffs extend to Dawlish and Dawlish Warren - the coast having been modified by the construction of the Great Western Railway. Dawlish Warren is an unusual double spit, consisting of the Inner and Outer Warrens with sand dunes and Greenland Lake occupying the intervening swale (Kidson 1950). There are shoals of sand off Exmouth and sand and gravel in the Exe estuary to the north. East of Exmouth cliffs of red sandstone begin at Orcombe and extend past Sandy Bay to Littleham Cove. This marks the start of a beach of ovoid cobbles that extends to Budleigh Salterton, and has been derived from cobble beds that descend through the cliff (> Fig. 7.11.6). East of Budleigh Salterton, the River Otter estuary was once a long narrow marine inlet, but is now occupied by salt marsh with much Spartina, backed by rushes and reeds. Bluffs (degraded cliffs) run up the eastern side of the Otter valley, and the shallow river mouth is constricted by a recurved shingle spit. Cliffs with shore platforms on the overlying red Triassic sandstone start at Otterton Point, where Otterton Ledge is a protruding sandstone bench. At Ladram Bay, thick almost flat bedded red sandstones form vertical cliffs 20–25 m high, and have been dissected along joint planes to form stacks, surrounded by a sandy beach with ­scattered

pebbles and coves with pocket beaches of shingle. The sandstone cliffs increase in height rapidly to High Peak (157 m) and Peak Hill (156 m), which are truncated by steep cliffs with angular vertical grooves and spurs on the cliff face (> Fig. 7.11.7). High cliffs of red Triassic marl and sandstone with buttresses and a basal apron of slumped sediment run behind the beach to the promontory at Sidmouth. The shingle beach here had become much depleted by 1995, when the shore was protected by boulders overlain by gravel brought from Woodbury quarry to form a ­nourished beach, which was held between three groynes and sheltered by two en echelon offshore breakwaters at the western end, built to break up SW storm waves. The River Sid percolates through a shingle bank to the sea. To the east are cliffs in red marl, sandstone and breccia, dipping gently eastward and fronted by shingle beaches interrupted by salients of sandstone and boulders known locally as ebbs. Rock falls occur frequently. Successive headlands have a capping of Upper Greensand, and are separated by deep valleys, as at Weston Mouth and Branscombe. Shingle beaches, consisting mainly of flint from the Chalk and chert from the Upper Greensand, occupy compartments along the coast, and typically show lateral gradation, with larger pebbles and cobbles towards the eastern ends (Bird 1989). To the east the cliff rises, the slope steepening as it passes from Triassic mudstone to the overlying Upper

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⊡⊡ Fig. 7.11.6 Gravel beds in the cliff west of Budleigh Salterton are the source of distinctive ovoid beach cobbles. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.7 The gullied cliff at Peak Hill, Sidmouth. (Courtesy Geostudies.)

Greensand and Chalk. The shingle beach ends at a landslide under Hooken, forming a hummocky undercliff with buttresses and blocks that have fallen from the upper cliff of Chalk. This descends to the coast at Beer Head, where the cliffs turn back into the shingle cove at Beer (> Fig. 7.11.8). In Seaton Bay a shingle beach fronts cliffs of red and grey marl, interrupted by a wooded chine, and then the esplanade begins. Seaton has a wide shingle beach, and the

sea wall forming the esplanade has tidal gates built in 1980 after a 1979 storm surge flooded the seafront and swept shingle over the road. At the eastern end of Seaton esplanade the River Axe reaches the sea, having been diverted eastward by a longshore shingle spit. The Axe estuary is broad at  high tide, and fringed by salt marshes backed by reedswamp. Triassic red marls form the cliffs from Axmouth to  Haver Cliff, the pink rock declining beneath white

South Devon

l­ imestones, then Blue Lias and Black Ven Marls. Cretaceous formations (Chalk, Upper Greensand and Gault) ­outcrop above, behind a subsided fringe. This is the start of the great Axmouth-Dowlands Landslide (Landslip), which took place at Christmas 1839 (Arber 1940). Masses of permeable Chalk and Upper Greensand slid seaward over Gault clay when the interface was lubricated and slurried by percolating groundwater. Behind the subsided area, an upper cliff of Chalk forms an abrupt edge to the coastal plateau. Luxuriant scrub ⊡⊡ Fig. 7.11.8 Chalk cliffs at Beer, looking towards Beer Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.9 The cliffs at Pinhay Bay, looking west towards the Landslip at Humble Point. (Courtesy Geostudies.)

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and woodland vegetation occupies a subsided area of dislocated blocks and chasms, with numerous fractures that are active, especially after wet weather, indicating that subsidence is continuing. The shore in front of the Axmouth-Dowlands Landslide is strewn with boulders, has irregular outcrops of tilted strata and segments of gravelly beach as in Charton Bay. Tumbled rocks extend along the shore past Humble Point to Pinhay Bay, where there is a basal cliff in Rhaetic White Lias (> Fig.  7.11.9). Steeply-sloping cliffs

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run on towards Seven Rock Point, where the basal cliff is in the underlying Blue Lias. The cliffs decline into scrubby bluffs at Devonshire Head which formerly protruded seaward and the shore  platform shows limestone layers forming scarplets between shale bands. The Devon-Dorset boundary crosses the coast half a mile west of The Cobb at Lyme Regis.

References Arber MA (1940) The coastal landslips of south-east Devon. Proc Geol Assoc 51:257–271 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Kidson C (1950) Dawlish Warren: a study of the evolution of sand spits across the mouth of the River Exe in Devon. Trans Inst Br Geogr 16:67–80

7.12  Dorset 1.  Introduction The Dorset coast is dominated by Jurassic formations, overlain by Cretaceous rocks including the Chalk, which forms cliffs between White Nothe and Lulworth Cove, in Worbarrow Bay, and on Ballard Down. The Chalk outcrop narrows eastward as the dip increases to form the Purbeck Ridge, which runs along the northern fringe of the Purbeck Anticline. To the north Lower Tertiary (Eocene) formations occupy the lowland around Poole Harbour and along the coast of Bournemouth Bay (Bird 1995). The present coastline of Dorset is largely the outcome of erosion and deposition during the past 6,000 years, since the sea reached its present level as a sequel to the worldwide Late Quaternary marine transgression. There are only a few surviving Pleistocene features, such as the Portland Raised Beach. It is likely that, as the Holocene sea rose to its present level, high coastal sectors presented steep, formerly periglaciated slopes, but these were soon undercut by marine erosion and converted into receding cliffs, the crests of which undulate across the truncated hills and valleys. In general, promontories occur where harder rocks outcrop at sea level, while coves and bays have been excavated in softer outcrops, particularly where these have already been lowered by stream erosion. In detail, lines of weakness such as joints and faults have been exploited to form clefts and caves, and locally to isolate stacks. Cliff profiles are related to varying rock resistance and permeability, so that the Chalk, Greensands and Jurassic limestones and sandstones all form bold outcrops, especially where they have been hardened as the result of intensive folding. However, in addition to geological factors it is necessary to take account of the effects of weathering and rainwash, as well as marine processes and con­­trasts between sheltered sectors and those exposed to high wave energy.

Cliff recession has been accompanied by landslides of various kinds, where masses of rock, weathered material and soil have slipped from high to low ground (Arber 1973; Brunsden and Jones 1972). They have occurred ­particularly where a capping of Chalk and Upper Green­ sand has collapsed over underlying outcrops of the less permeable Gault and Lias clays, lubricated to mudslides by seeping groundwater, as on Ware Cliffs and White Nothe. In a similar way the Portland Beds have slid down over Kim­meridge Clay, as at St. Alban’s Head, Hounstout Cliff, Gad Cliff and on the northern slopes of the Isle of Portland. The movements have been over a clearly definable surface, either translational (down a plane) or rotational (over a concave curve). On soft fine-grained sediment, mudslides usually lobate and elongated, move slowly downslope, torn by shearing and tension cracks. As they subside they leave arcuate cliff-backed hollows. More rapid movements, especially during very wet weather, include mudflows (where unconsolidated materials soaked with water move as a viscous mass), topples (where lumps of rock and soil break away and subside) and rock falls (when detached material drops from a cliff and forms a scree of talus). Shore platforms have been cut by marine erosion as cliffs recede. Some are seaward-sloping intertidal abrasion platforms while others are subhorizontal structural ledges on the surface of resistant strata outcropping at the cliff base. West of Portland Bill the Dorset coast is exposed to ocean swell from the Atlantic Ocean and strong wave action generated by the prevailing south-westerly winds; to the east it is at first sheltered by the Isle of Portland, but exposure increases again beyond White Nothe. There is occasionally strong wave action from the south and southeast, most evident on east-facing shores in Weymouth Bay, and between Durlston Head and Poole. Mean spring tide ranges are smaller on the Dorset coast than in south-west or south-east England. At Lyme Regis

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run on towards Seven Rock Point, where the basal cliff is in the underlying Blue Lias. The cliffs decline into scrubby bluffs at Devonshire Head which formerly protruded seaward and the shore  platform shows limestone layers forming scarplets between shale bands. The Devon-Dorset boundary crosses the coast half a mile west of The Cobb at Lyme Regis.

References Arber MA (1940) The coastal landslips of south-east Devon. Proc Geol Assoc 51:257–271 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Kidson C (1950) Dawlish Warren: a study of the evolution of sand spits across the mouth of the River Exe in Devon. Trans Inst Br Geogr 16:67–80

7.12  Dorset 1.  Introduction The Dorset coast is dominated by Jurassic formations, overlain by Cretaceous rocks including the Chalk, which forms cliffs between White Nothe and Lulworth Cove, in Worbarrow Bay, and on Ballard Down. The Chalk outcrop narrows eastward as the dip increases to form the Purbeck Ridge, which runs along the northern fringe of the Purbeck Anticline. To the north Lower Tertiary (Eocene) formations occupy the lowland around Poole Harbour and along the coast of Bournemouth Bay (Bird 1995). The present coastline of Dorset is largely the outcome of erosion and deposition during the past 6,000 years, since the sea reached its present level as a sequel to the worldwide Late Quaternary marine transgression. There are only a few surviving Pleistocene features, such as the Portland Raised Beach. It is likely that, as the Holocene sea rose to its present level, high coastal sectors presented steep, formerly periglaciated slopes, but these were soon undercut by marine erosion and converted into receding cliffs, the crests of which undulate across the truncated hills and valleys. In general, promontories occur where harder rocks outcrop at sea level, while coves and bays have been excavated in softer outcrops, particularly where these have already been lowered by stream erosion. In detail, lines of weakness such as joints and faults have been exploited to form clefts and caves, and locally to isolate stacks. Cliff profiles are related to varying rock resistance and permeability, so that the Chalk, Greensands and Jurassic limestones and sandstones all form bold outcrops, especially where they have been hardened as the result of intensive folding. However, in addition to geological factors it is necessary to take account of the effects of weathering and rainwash, as well as marine processes and con­­trasts between sheltered sectors and those exposed to high wave energy.

Cliff recession has been accompanied by landslides of various kinds, where masses of rock, weathered material and soil have slipped from high to low ground (Arber 1973; Brunsden and Jones 1972). They have occurred ­particularly where a capping of Chalk and Upper Green­ sand has collapsed over underlying outcrops of the less permeable Gault and Lias clays, lubricated to mudslides by seeping groundwater, as on Ware Cliffs and White Nothe. In a similar way the Portland Beds have slid down over Kim­meridge Clay, as at St. Alban’s Head, Hounstout Cliff, Gad Cliff and on the northern slopes of the Isle of Portland. The movements have been over a clearly definable surface, either translational (down a plane) or rotational (over a concave curve). On soft fine-grained sediment, mudslides usually lobate and elongated, move slowly downslope, torn by shearing and tension cracks. As they subside they leave arcuate cliff-backed hollows. More rapid movements, especially during very wet weather, include mudflows (where unconsolidated materials soaked with water move as a viscous mass), topples (where lumps of rock and soil break away and subside) and rock falls (when detached material drops from a cliff and forms a scree of talus). Shore platforms have been cut by marine erosion as cliffs recede. Some are seaward-sloping intertidal abrasion platforms while others are subhorizontal structural ledges on the surface of resistant strata outcropping at the cliff base. West of Portland Bill the Dorset coast is exposed to ocean swell from the Atlantic Ocean and strong wave action generated by the prevailing south-westerly winds; to the east it is at first sheltered by the Isle of Portland, but exposure increases again beyond White Nothe. There is occasionally strong wave action from the south and southeast, most evident on east-facing shores in Weymouth Bay, and between Durlston Head and Poole. Mean spring tide ranges are smaller on the Dorset coast than in south-west or south-east England. At Lyme Regis

Dorset

the range is 3.7 m, and at Bridport 3.5 m. The tides diminish east of Portland Bill to 1.9 m at Portland, 2.0 m in Lulworth Cove and 1.7 m between Swanage and Bournemouth. There are the complications of prolonged low tide in Portland Bay and prolonged high tide in Bournemouth Bay - the latter facilitating wave attack on cliffs. The effects of marine erosion and deposition have been to simplify coastal outlines, so that only the most resistant rocks persist as promontories, such as the massive limestones of Portland Bill and the southern shores of Purbeck. Lyme Bay has evolved from a series of small bays with irregular headlands to the relatively smooth coastline that now runs from Torquay to Portland Bill as the result of cliff recession and the deposition of Chesil Beach. The coastline has become adjusted to fit the patterns of waves that move in across Lyme Bay. Between Portland Bill and the Isle of Purbeck there are still minor rocky promontories between small curving bays with beaches, but in Bournemouth Bay the cliffed coastline cut in softer Eocene rocks has become more open and smoothly curved, with some sandy deposition on the western shores. As these cliffs receded, small streams entrenched steep-sided valleys known as chines (Bury 1920). The beaches of Dorset are generally of local derivation, consisting of sand and gravel produced by erosion of nearby cliffs and shore outcrops, and distributed along the coast by wave action. But, as is evident on Chesil Beach, some beach material has been washed up from the sea floor. Most of the beaches are dominated by flint and chert shingle, with locally derived limestone pebbles and sandy material from weathered sandstones, especially in Bou­ rnemouth Bay, where the sandy beaches have been derived from Eocene formations. Some beaches are still receiving sand and gravel eroded from nearby cliffs or the sea floor but others, like Chesil Beach now have little if any sediment supply. Very little beach material has come from rivers, which have been mainly delivering fine-grained sediment (silt and clay) to the coast. In general the predominance of wind and wave action from the south-west has caused beach material to drift eastward, as shown by the accretion of beaches on the western sides of natural obstacles such as landslide lobes under Golden Cap and man-made ob­­ stacles such as breakwaters and groynes, like the Cobb at Lyme Regis. Estuarine inlets have been formed by Late Quaternary marine submergence of valley mouths, notably in Poole and Christchurch Harbours, which have been modified by Holocene deposition to alluvial plains and areas of salt marsh, mudflats and shoals.

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2.  The Dorset Coastline East of Devonshire Head the cliffs have been cut in the layered strata of the Blue Lias and the intertidal area consists of broad structural ledges of limestone where the intervening blue clay horizons have been washed out (>Fig.  7.12.1). Chippel Bay occupies a gentle syncline, with ledges rising again towards the Cobb, the stone breakwater that shelters the ancient harbour of Lyme Regis. Early in the nineteenth century the coast at Lyme Regis had a wide curving beach of brown shingle in front of scrubby slumping bluffs (Wanklyn 1922). A stone and concrete promenade was built at Cobb Gate, and later extended to protect the coastal slopes from wave erosion, and there has been recurrent slumping in the coastal slopes of Middle Lias. Attempts are being made to stabilise the coastal slopes in Lyme by draining off groundwater. Eastward drifting of beach material has been prevented by the Cobb, and the shingle beach of the Lyme seafront has been gradually depleted. It was restored artificially in 2007. At the eastern end of Lyme seafront the coastline curves behind Church Cliffs towards Black Ven. The Blue Lias is overlain by marls, and there is an upper, receding cliff in Gault and Upper Greensand. There have been frequent landslides, the Lias marls and capping Gault having subsided irregularly in front of the receding cliff of Upper Greensand (>Fig. 7.12.2). Cracks develop at the top of the cliff, which recede as blocks calve away and collapse (>Fig. 7.12.3). The cliffs descend towards Charmouth where the mouth of the River Char is encumbered by shingle, and there is a looped intertidal shingle bar off the river mouth. The cliffs ascend steeply towards Stonebarrow Hill, which has a high arcuate landslide back-scar (a curved slipplane) in Upper Greensand, overlooking an amphitheatre of slumped material, Cains Folly, in which rotational slides are moving down towards a 40–60 m lower cliff, where mudslides spill over and down to the shore (>Fig.  7.12.4). This receding coast may be in ‘dynamic equilibrium’ over periods of a century or so, meaning that there have been similar associations of upper and lower cliff and intervening slumped topography, maintaining the transverse profile as the coast retreated (Brunsden and Jones 1976, 1980). Golden Cap (191 m) is the highest coastal cliff in southern England. Its summit is one of many flat-topped ridges, dissected remnants of the West Dorset plateau, and the cliff has yellow-brown Upper Cretaceous rocks over the grey Liassic clays and limestones. There is landslide topography, with slumping slopes in Lias beneath the Upper Greensand capping. Mud lobes have descended to the shore, separating pocket beaches of shingle in small

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⊡⊡ Fig. 7.12.1 Limestone ledges and beach boulders on the shore at Devonshire Head, looking east towards the Cobb at Lyme Regis. The dark marl horizon can be seen below a limestone layer in the scarplet. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.2 The slumping cliff at Black Ven, showing the 2008 landslide undercut at the base. (Courtesy Geostudies.)

coves, and the slumped material has been cut back to leave boulder festoons on the shore. At Seatown the cliffs decline to the mouth of the River Winniford, which percolates out through stormpiled shingle on a beach that is low, flat and sandy with only sparse fine shingle to the west but a higher, steeper and coarser shingle beach to the east. This lateral

g­ radation is characteristic of the beaches of Lyme Bay (Bird 1989). Pebbles have been quarried from this beach for use in ornamental walls and pavements (>Fig. 7.12.5). The cliff east of Seatown is of slumping blue clay, and as it rises towards Ridge Hill the slope steepens on overlying sandstones. Doghouse Hill (132 m) has a small ­capping

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⊡⊡ Fig. 7.12.3 Capping of weathered (oxidised) Upper Greensand on the receding upper cliff at Black Ven. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.4 A basal fan of slumped debris that spilled over the lower cliff east of Charmouth during the wet spring of 2001. It was subsequently cut back behind an abrasion ramp cut in shale. (Courtesy Geostudies.)

of yellow Bridport Sands. Each of the hill crests carries cherty gravel in a brown earthy matrix, and the cliff crest shows that the slopes are mantled by similar material washed downhill. A coombe separates Doghouse Hill from the higher Thorncombe Beacon (157 m), and the little valley at Eype opens between low slumping clay cliffs. Recurrent slumping on West Cliff has been countered by the building of a promenade. The West Bay harbour breakwaters were built as piers in 1823–1825 and extended as a marina in 2003.

The vertical East Cliff and Burton Cliff, on either side of the Bride valley below Burton Bradstock, are cut in Bridport Sands, here almost horizontal calcareous sandstone layers with intervening softer yellow sand. Where there have been recent cliff falls there are basal talus cones, and the cliff face is smooth, but intervening areas show protruding structural ledges of calcareous sandstone etched out by weathering (>Fig. 7.12.6). At Burton Beach the cliffs end abruptly, but Chesil Beach continues as a shingle bank about 150 m wide,

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⊡⊡ Fig. 7.12.5 Seatown, looking west along a shingle beach towards Golden Cap. (Courtesy Geostudies.)

fronting marshy meres, and later the lagoon known as The Fleet (>Fig. 7.12.7). Its crest gradually increases in height from about 6 m above low tide level at Burton Cliff to 14 m at Chiswell, at the SE end. Chesil Beach is a vast accumulation mainly of flint and chert shingle, with some Triassic quartzites, a few pebbles of granite, quartz and metamorphic rock, all derived from formations that outcrop in South Devon. At the south-eastern end there is an increasing proportion of limestone pebbles and cobbles derived from the cliffs of the Isle of Portland. It is essentially a relict formation (no longer receiving much beach material) and probably originated as an offshore barrier which was driven landward during the Late Quaternary marine transgression in such a way that the inner shores of The Fleet have never been exposed to erosion by the waves of the open sea. There is lateral grading of the beach material, the longest diametre of pebbles at West Bay being about 1 cm, increasing gradually to 2–3 cm at Abbotsbury and 10–15 cm at Portland. This is due largely to sorting of the

a­ vailable shingle by size, for attrition of flint and chert pebbles is very slow. Long­shore drifting occurs when waves arrive obliquely to the beach, and it may be that the sorting is the outcome of the contrast between stronger westerly and SW waves, which can move all grades of pebbles, and the weaker SE waves, which take back only the smaller material. This has proceeded over a long period, during the landward migration of Chesil Beach in Holocene times (Carr and Blackley 1974). The Fleet is a shallow, brackish tidal lagoon, separated from the sea by Chesil Bank and having a tidal entrance from Portland Bay at the SE end. It is about 8 miles (13 km) long, from Abbotsbury Swannery to the tidal entrance at Small Mouth, and varies in width from about 75 m at The Narrows to nearly a kilometre in Butterstreet Cove. Tide range increases south-eastward to Small Mouth, where there is prolonged ebb because sea seeps in through Chesil Beach while the tide is still falling in Portland Harbour, especially when sea level in Lyme Bay is raised by westerly winds. At low tide extensive shoals of muddy sand, bearing seagrasses, are exposed in the southeast, while the north-west, towards Abbotsbury Swannery, is less tidal and somewhat freshened by inflowing streams, so that reedswamp has invaded the shallows. Chesil Beach comes to an end below West Weare, and vertical cliffs in Portland Limestone extend south towards Portland Bill (>Fig. 7.12.8). At Portland Bill storm-washed structural ledges of Portland Stone end in cliffs that descend directly into deep water, and Pulpit Rock is a joint-bounded stack rising from a structural ledge. The Raised Beach, much disturbed by quarrying, is banked on a ledge of Portland Limestone, and forms a gently sloping terrace that rises from 7 to 18 m above present sea level and extends for about a mile to the NE. It was formed when sea level was about 12 m higher than it is now, and is thought to comprise two emerged beaches, formed 210,000 years and 125,000 years ago (Davies and Keen 1985). Offshore is the Tide Race, where the ebb and flow of tides interacts with waves to produce rough water over the shingle shoal known as The Shambles. The east coast of the Isle of Portland is more sheltered from strong wave action than the west coast, but the steep sharp-crested coastal slope of Portland Stone increases in altitude northward to East Weare. Portland Bay has been enclosed by breakwaters to form Portland Harbour, and piers have been built on either side of the harbour at the mouth of the River Wey. Weymouth promenade, to the east, is fronted by a wide curving sandy beach, the upper part of which becomes gravelly near the clock tower. The shingle beach then becomes wider and higher, usually with several berms,

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⊡⊡ Fig. 7.12.6 East Cliff, near Bridport, is cut in horizontally bedded sandstones. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.7 The crest of Chesil Beach near Abbotsbury, with the Fleet lagoon on the left. (Courtesy Geostudies.)

continuing in front of the concrete sea wall that borders the Lodmoor lowland. The beach has been augmented by dumping shingle to protect the sea wall. There are cliffs in Oxford Clay to the east, and Redcliff Point is bordered by mudflows, which descend calving, crumbling, flaking and slurrying to form lobate tongues that advance over the shingle beach on the eastern side. This beach consists of brown flint and chert shingle,

g­ enerally coarser to the west and finer to the east, in contrast with Chesil Beach and the Lyme Bay beaches. This is probably because this part of the coast is sheltered from the prevailing westerly and SW waves by the Isle of Portland, so that here the SE waves are dominant, and have moved the larger pebbles westward. Kimmeridge Clay outcrops in the upper part of the sloping cliff at Black Head, and a slumping cliff declines to

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⊡⊡ Fig. 7.12.8 Cliffs in Portland Limestone on the coast of Portland Bill. (Courtesy Geostudies.)

the valley at Osmington Mills. Cliffs continue, subsiding behind the shingle beach at Ringstead, and then rising again to Burning Cliff (so-called because a fire ignited spontaneously in the Kimmeridge oil shale in 1826, and persisted for several years). This overlooks a widening area of scrubby landslide topography on the Kimmeridge Clay, which forms low cliffs and concave ramps immediately behind the beach. Below Holworth House the Kimmeridge Clay is overlain by a disintegrating cliff of Portland Sand, capped by Portland Stone then Purbeck Beds, all dipping north-west. These Upper Jurassic formations are overlain unconformably by the more gently south-eastward dipping Gault, Upper Greensand and Chalk, which descend through the upper cliffs to White Nothe. The cliffs to the east are cut in frost-shattered Chalk, rising and falling as they intersect ridges and coombes. Beyond Swyre Head, is the irregular promontory of Durdle Door. At the eastern end of St. Oswald’s Bay the Chalk passes inland behind a valley cut in the underlying softer formations. This climbs behind the Purbeck and Portland ridge on Dungy Head to meet another valley that descends into Lulworth Cove. This occurs where converging valleys cut into the soft Wealden sands and clays have been scoured out by marine erosion behind a coastal ridge of steeply dipping Purbeck and Portland Limestone. It is backed by a steepened slope cut into the Chalk escarpment after the outer rampart of Purbeck and Portland Limestone had been breached by the sea (>Fig. 7.12.9). Along the coast to the east, a steep escarpment of Portland Stone is capped by Purbeck Beds. The outer ­rampart of Portland Stone breaks up into Mupe Rocks,

and in Mupe Bay the cliffs curving behind a beach of sand and gravel show a sequence from the Purbeck Beds through the crumbling Wealden to the Lower Greensand, partly obscured by Chalk talus, the Gault and Upper Greensand and then a sudden rise to the high Chalk cliff cut into the escarpment. The cliff crest descends just as abruptly to Arish Mell, where cliff recession has intersected the mouth of a steep-sided but gently declining dry valley to form a gap. Flower’s Barrow is an Iron Age hilltop camp truncated by cliff recession, and the cliff at Cow Corner rises sharply, with ribs of chalk jutting out between clefts and caves. Worbarrow Bay is an enlarged and more open version of Lulworth Cove, exposing the Wealden Beds cut back through the Greensands and Gault to the Chalk. At the eastern end of the bay is Worbarrow Tout, a steep promontory much like Durdle Door. The Chalk ridge runs inland as the Purbeck Hills, and from Pondfield Cove the Portland and Purbeck Beds ascend rapidly to outcrop high in Gad Cliff. At the eastern end is Tyneham Cap, where the Portland Stone becomes a well-defined escarpment that runs behind a broad amphitheatre of Kimmeridge Clay, and round to Swyre Head. Behind Brandy Bay are cliffs cut in the layered limestones and blue shales of the Kimmeridge Clay, fronted by structural ledges developed on resistant limestone layers, as at Broad Bench (>Fig. 7.12.10). Undulating cliffs to the east are cut in the Kimmeridge Clay dipping eastward. Successive limestone bands des­ cend to the cliff base, forming slight salients, running out  across the shore as ledges, some of which have

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⊡⊡ Fig. 7.12.9 Lulworth Cove. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.10 Broad Bench, a structural ledge. (Courtesy Geostudies.)

­ isintegrated into boulder beaches. The deeply incised d Encombe valley is followed by cliffs rising to Hounstout, then descending to Chapman’s Pool, a rounded cove formed where the sea has scooped out soft Kimmeridge Clay between bouldery spurs at the mouth of a valley. The coast turns southward as cliffs run out to St. Alban’s (St Aldhelm’s) Head, (108 m), then eastward as a coastal slope descends to cliffs up to 60 m high in hard stratified Portland Stone towards Durlston Head. The more exposed southern shore of Durlston Head is boldly cliffed, but to

the north, in Durlston Bay, the east-facing coast consists of steep, crumbling slopes exposing the sequence of Purbeck Beds rising northward. Swanage Bay has been cut into the Wealden Beds at the eastern end of the lowland that extends through the Isle of Purbeck. At the northern end of Swanage Bay in Punfield Cove the steeply dipping Wealden Beds are overlain by Lower Greensand, Gault and Upper Greensand beneath the Ballard Down Chalk, which forms an escarpment cliff out to Ballard Point (>Fig. 7.12.11).

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⊡⊡ Fig. 7.12.11 Escarpment cliff in Chalk at Ballard Down. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.12 Chalk stacks at Handfast Point, on the south coast of Studland Bay. (Courtesy Geostudies.)

On the east-facing coast of Ballard Down the vertical cliffs decline to Handfast Point. In contrast to the similar coast at the eastern end of the Isle of Wight there is no shore platform. Towards Handfast Point the Chalk cliff becomes strongly crenulate, with bays and promontories cut along joint planes, erosion having separated an elongated peninsula into the stacks known as The Pinnacles and Old Harry Rocks (>Fig. 7.12.12). At the southern end of Studland Bay an eroded surface in Upper Chalk is overlain by the Reading Beds. Along the

shore are some large blocks of ferruginous sandstone and a beach of smooth, rounded chalk cobbles and pebbles and mainly unworn or broken flint nodules. The chalk cobbles and pebbles diminish rapidly northward and disappear as the result of abrasion (May and Heeps 1985). North of Redend Point the beach becomes sandy, backed by parallel dune ridges on the South Haven Pen­ insula. Beyond is Shell Bay, a curving NW to South Haven Point at the entrance to Poole Harbour, which has been constricted by formation of bordering spits, South Haven

Hampshire

Peninsula to the south and the Sandbanks peninsula extending SW from Bournemouth beach. Poole Harbour is a tidal estuary fed by several rivers, the largest of which are the Frome and the Piddle, entering at the western end. It shows the branched configuration typical of a valley system partly drowned by the Late Quaternary marine transgression, interfluvial ridges persisting as promontories and high islands. A century ago Poole Harbour had relatively narrow fringing salt marshes and wide soft mudflats exposed at low tide, but the ­salt-tolerant rice (cord) grass, Spartina anglica, became established in 1899 and quickly spread across intertidal areas to form wide swards (Bird and Ranwell 1964). The Spartina marshland is generally wider on shores sheltered from strong wave action, as on the eastern sides of islands and peninsulas, but thins out where tidal channels run close to the shore (May 2006). Sandbanks spit, formerly dune-capped, is now built over. East from Sandbanks former cliffs have been almost entirely stabilised as artificial vegetated slopes extending past Bournemouth to Branscombe. Cliffs then rise to Hengistbury Head. Offshore mining of ironstone nodules in the nineteenth century deepened the water and allowed stronger waves to attack the cliffs, which then began to retreat rapidly. In 1938 a groyne was built at the  end of the promontory, and this trapped eastward-­ drifting sand and gravel to form a widening beach, backed by low dunes in front of a cliff that became weathered back to a gentler bluff, with basal fans of sand and gravel no longer reached by the sea. The eastern cliff of Hengistbury Head then lost its beach, and was retreating with basal heaps of talus and ironstone, until recently, when it was stabilised behind an artificial beach held by several boulder groynes. A spit runs out northward to the entrance to Christchurch Harbour. Fed by the Rivers Stour and Avon,

7.13

Christchurch Harbour is a smaller version of Poole Harbour, with salt marshes, including some Spartina, bordering tidal mudflats and creeks. From Mudeford the coast is protected by sea walls, and the former cliffs in soft clay and sand have been largely stabilised as sloping vegetated bluffs. A series of massive groynes between sand and shingle beaches continues eastward in front of the stabilised coastal slopes, and this largely man-made coastal scenery runs on to the steep-sided chine at Chewton Bunny where the Dorset coast comes to an end at the Hampshire border.

References Arber MA (1973) Landslips near Lyme Regis. Proc Geol Assoc 84:121–133 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Bird ECF (1995) Geology and scenery of Dorset. Ex-Libris Press, Bradford on Avon, Wiltshire, pp 207 Bird ECF, Ranwell DS (1964) The physiography of Poole Harbour, Dorset. J Ecol 52:355–366 Brunsden D, Jones DKC (1972) The morphology of degraded landslide slopes, south-west Dorset. Quart J Engrg Geol 5:205–222 Brunsden D, Jones DKC (1976) The evolution of landslide slopes in Dorset. Philos Trans R Soc Lond A 283:605–631 Brunsden D, Jones DKC (1980) Relative time scales and formative events in coastal landslide systems. Zeitschrift für Geomorphologie, Supple­ ment band 34:1–19 Bury H (1920) The chines and cliffs of Bournemouth. Geol Mag 57:71–76 Carr AP, Blackley MWL (1974) Ideas on the origin and development of Chesil Beach, Dorset. Proc Dors Arch Soc 95:1–9 Davies KH, Keen DH (1985) The age of Pleistocene marine deposits at Portland, Dorset. Proc Geol Assoc 96:217–225 May VJ (2006) The ecology of Poole Harbour. Elsevier, Oxford May VJ, Heeps C (1985) The nature and rates of change on chalk coastlines. Zeitschrift für Geomorphologie, Supplement band 57:81–94 Wanklyn F (1922) History of Lyme Regis. Humpreys, London

7.13  Hampshire 1.  Introduction Hampshire occupies the northern part of a broad syncline that runs from west to east through The Solent and Spithead. The Chalk of the Dorset Downs and Salisbury Plain subsides beneath Tertiary formations comprising the Bagshot Beds (Poole Formation), overlain by the Lower Oligocene Barton and Bracklesham Beds and the Upper Oligocene Hamstead and Bembridge Beds: a succession of sands, clays, marls and some limestones. In

Late Tertiary times the so-called Solent River was an eastward extension of the Dorset Frome eastward through a valley which ran along the syncline, and the geological formations are sequences of marine, estuarine, deltaic and valley floor deposits that formed as sea level rose and fell (marine transgressions and regressions) along this valley (Everard 1954). The southern side of the valley was for long the Chalk ridge of the Purbeck Hills extending eastward through the centre of the Isle of Wight. This was breached by marine erosion to form the 24 km wide gap

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Peninsula to the south and the Sandbanks peninsula extending SW from Bournemouth beach. Poole Harbour is a tidal estuary fed by several rivers, the largest of which are the Frome and the Piddle, entering at the western end. It shows the branched configuration typical of a valley system partly drowned by the Late Quaternary marine transgression, interfluvial ridges persisting as promontories and high islands. A century ago Poole Harbour had relatively narrow fringing salt marshes and wide soft mudflats exposed at low tide, but the ­salt-tolerant rice (cord) grass, Spartina anglica, became established in 1899 and quickly spread across intertidal areas to form wide swards (Bird and Ranwell 1964). The Spartina marshland is generally wider on shores sheltered from strong wave action, as on the eastern sides of islands and peninsulas, but thins out where tidal channels run close to the shore (May 2006). Sandbanks spit, formerly dune-capped, is now built over. East from Sandbanks former cliffs have been almost entirely stabilised as artificial vegetated slopes extending past Bournemouth to Branscombe. Cliffs then rise to Hengistbury Head. Offshore mining of ironstone nodules in the nineteenth century deepened the water and allowed stronger waves to attack the cliffs, which then began to retreat rapidly. In 1938 a groyne was built at the  end of the promontory, and this trapped eastward-­ drifting sand and gravel to form a widening beach, backed by low dunes in front of a cliff that became weathered back to a gentler bluff, with basal fans of sand and gravel no longer reached by the sea. The eastern cliff of Hengistbury Head then lost its beach, and was retreating with basal heaps of talus and ironstone, until recently, when it was stabilised behind an artificial beach held by several boulder groynes. A spit runs out northward to the entrance to Christchurch Harbour. Fed by the Rivers Stour and Avon,

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Christchurch Harbour is a smaller version of Poole Harbour, with salt marshes, including some Spartina, bordering tidal mudflats and creeks. From Mudeford the coast is protected by sea walls, and the former cliffs in soft clay and sand have been largely stabilised as sloping vegetated bluffs. A series of massive groynes between sand and shingle beaches continues eastward in front of the stabilised coastal slopes, and this largely man-made coastal scenery runs on to the steep-sided chine at Chewton Bunny where the Dorset coast comes to an end at the Hampshire border.

References Arber MA (1973) Landslips near Lyme Regis. Proc Geol Assoc 84:121–133 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Bird ECF (1995) Geology and scenery of Dorset. Ex-Libris Press, Bradford on Avon, Wiltshire, pp 207 Bird ECF, Ranwell DS (1964) The physiography of Poole Harbour, Dorset. J Ecol 52:355–366 Brunsden D, Jones DKC (1972) The morphology of degraded landslide slopes, south-west Dorset. Quart J Engrg Geol 5:205–222 Brunsden D, Jones DKC (1976) The evolution of landslide slopes in Dorset. Philos Trans R Soc Lond A 283:605–631 Brunsden D, Jones DKC (1980) Relative time scales and formative events in coastal landslide systems. Zeitschrift für Geomorphologie, Supple­ ment band 34:1–19 Bury H (1920) The chines and cliffs of Bournemouth. Geol Mag 57:71–76 Carr AP, Blackley MWL (1974) Ideas on the origin and development of Chesil Beach, Dorset. Proc Dors Arch Soc 95:1–9 Davies KH, Keen DH (1985) The age of Pleistocene marine deposits at Portland, Dorset. Proc Geol Assoc 96:217–225 May VJ (2006) The ecology of Poole Harbour. Elsevier, Oxford May VJ, Heeps C (1985) The nature and rates of change on chalk coastlines. Zeitschrift für Geomorphologie, Supplement band 57:81–94 Wanklyn F (1922) History of Lyme Regis. Humpreys, London

7.13  Hampshire 1.  Introduction Hampshire occupies the northern part of a broad syncline that runs from west to east through The Solent and Spithead. The Chalk of the Dorset Downs and Salisbury Plain subsides beneath Tertiary formations comprising the Bagshot Beds (Poole Formation), overlain by the Lower Oligocene Barton and Bracklesham Beds and the Upper Oligocene Hamstead and Bembridge Beds: a succession of sands, clays, marls and some limestones. In

Late Tertiary times the so-called Solent River was an eastward extension of the Dorset Frome eastward through a valley which ran along the syncline, and the geological formations are sequences of marine, estuarine, deltaic and valley floor deposits that formed as sea level rose and fell (marine transgressions and regressions) along this valley (Everard 1954). The southern side of the valley was for long the Chalk ridge of the Purbeck Hills extending eastward through the centre of the Isle of Wight. This was breached by marine erosion to form the 24 km wide gap

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Hampshire

now seen between Ballard Down in Dorset and The Needles at the western end of the Isle of Wight, and thus opens up Bournemouth Bay, Christchurch Bay and The Solent along the line of the submerged valley (Melville 1982). The coast of Hampshire is generally low-lying and fringed mainly by sand and shingle beaches and salt marshes. There are a number of estuarine inlets, as at Lymington and the Beaulieu River, and on a larger scale Southampton Water (the drowned mouths of the Test, Itchen and Hamble valleys), Portsmouth Harbour, Lang­stone Harbour and Chichester Harbour. Each of these is the outcome of marine submergence during the Late Quaternary (Flandrian) marine transgression, and they remain relatively open because the rivers draining into them have carried only small sediment loads. Mean spring tide ranges are generally small, with double high tides in the Solent, Southampton Water and Spithead. The range at Christchurch is 1.4 m and on Hurst Castle Spit 2.2 m. Lymington has 2.5 m, Calshot Castle 3.8 m and Southampton 4.0 m. At Lee on Solent the tides attain 3.7 m and at Portsmouth 4.1 m.

2.  The Hampshire Coast The coast of Christchurch Bay consists of steep bluffs and cliffs in the yellow and white Highcliffe Sands at the top of the Bracklesham Beds. At various times a spit grew

a­ longshore from Mudeford as far as Highcliffe Castle, deflecting the outlet from Christchurch Harbour eastward, but the harbour entrance has been stabilised by breakwaters. East of Highcliffe the Bracklesham Beds dip below sea level and an upper cliff of gravel-capped Barton Clay develops behind a slumped foreground, where mudflows and clay lobes have spread out across the shore (> Fig. 7.13.1). The cliffs rise to more than 35 m at Barton-on-Sea. There has been general recession of the cliffed coastline of Christchurch Bay during the past century. The crest of the upper cliff at Barton-on-Sea has retreated up to a metre per year, and several buildings have been lost. Sectors of the coastline have now been stabilised (> Fig. 7.13.2). The shingle beach continues along the coast, much of which is now modified by walls or wooden revetments. Near Milford on Sea headwater valleys of Dawes Stream have been beheaded by cliff recession to form cliff cols (similar to the valleuses of the Normandy coast). The cliffs decline east of Milford-on-Sea and the shingle beach continues as a high bank that ends in the recurved spit at Hurst Castle Point. In its lee are extensive salt marshes and mudflats exposed at low tide off Keyhaven, where the lowlying hinterland is protected by a high tide embankment at Sea Grass Lane. There is an intricate system of tidal creeks and some shelly cheniers on the salt marsh, which has an eroding microcliff (up to 1.5 m high) along its seaward margins (Pye and French 1993). Introduction of Spartina anglica late in the nineteenth century resulted in rapid vertical accretion and widening of the salt marsh terrace, but since 1930 there has been die-back of Spartina ⊡⊡ Fig. 7.13.1 Slumping clay cliffs at Barton, Christchurch Bay. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.13.2 Stabilised coast at Barton-on-Sea. (Courtesy Geostudies.)

(­ preceding erosion) along the seaward margins and beside tidal creeks, and in pans within the marshes (Bradbury 1996). Hurst Castle spit grew from the west, supplied with shingle drifting along the shore from Christchurch Bay (Lewis 1931; King and McCullach 1971). It has been shaped by SW waves, which have piled up Hurst Beach and driven it landward (over salt marsh) during storm surges (as in 1989–1990, when it moved up to 80 m), and easterly waves from the Solent, which have shaped the successive parallel recurves and built beach ridges at the eastern end (> Fig. 7.13.3). Narrow beaches back salt marshes along the coast east from the Lymington estuary, and there are successive spits at Warren Farm and Needs Oar (Ore) Point. At Inchmery Quay the salt marsh is being dissected by erosion (> Fig. 7.13.4). Brown flint shingle dominates Calshot spit, which curves out to Calshot Castle. It shelters salt marshes which extend NW along the shore of Southampton Water to Hythe. NW of Hythe former salt marshes have been ­extensively embanked and reclaimed for port and industrial development, but a cliffed salt marsh persists at Totton. There has been much land reclamation north of Southampton Water for quays and docks and for urban or industrial use, but south of the Itchen estuary, narrow sand and shingle beaches front low bluffs and cliffs in Bracklesham Beds at Netley and Warsash. Exposure to SW wave action increases towards Lee-on-the-Solent, where there has been cliff recession; but sea walls have been built to halt erosion here and in Stokes Bay. A broad

lobate shingle foreland has formed at Gilkicker Point, but the coast is now largely walled at Gosport and Portsmouth, with extensive docks. Within Portsmouth Harbour there are salt marshes and intertidal mudflats. Portsea Island is largely urbanised, with reclaimed areas along its coasts. The southern (Spithead) coast has Southsea promenade, and is fringed by a shingle beach (largely submerged at high tide) retained by a lattice pattern of groynes. The entrance to Langstone Harbour is bordered by inwashed paired shingle and sand spits. Glacial erratics have been found at the entrance to Langstone Harbour, possibly deposited from melting icebergs in Pleistocene times. Langstone Harbour is much like Portsmouth Harbour, with salt marshes, mudflats and tidal creeks converging southward, and much of the coastline walled at high tide level. The bordering beaches at Eastney and Hayling Island are mainly of brown and white subangular flint shingle, with white shelly berms. Hayling Island is low-lying and covered with brown Pleistocene brickearth, part of the wind-deposited loess that extends eastward along the Sussex coastal plain. There are minor bluffs, fringed by salt marshes and parallel shingle beach ridges on scrubby Sinah Common, in the SW of the island, which are surmounted by dunes. Salt marshes are extensive on the eastern shore. A shingle beach runs along the south coast of the island, with a sea wall and numerous groynes, to Eastoke Point. The entrance to Chichester Harbour is also bordered by paired spits, and the Sussex boundary runs down Emsworth Channel to the west of Thorney Island.

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⊡⊡ Fig. 7.13.3 Shingle beach ridges at the end of Hurst Castle spit. (Courtesy Geostudies.)

⊡⊡ Fig. 7.13.4 Eroding salt marsh at Inchmery Quay. (Courtesy Geostudies.)

References Bradbury A (1996) Western Solent salt marsh study. National Rivers Authority, Bristol, UK Everard CE (1954) The Solent river: a geomorphological study. Trans Inst Br Geogr 20:41–54 King CAM, McCullach MJ (1971) A simulation model of a complex recurved spit. J Geol 79:22–36

Lewis WV (1931) Effect of wave incidence on the configuration of a ­shingle beach. Geogr J 78:131–148 Melville RV (1982) The Hampshire Basin and adjoining areas. British Regional Geology, 4th edn. HMSO, London Pye K, French PW (1993) Erosion and accretion on British salt marshes. Cambridge Environmental Research Consultants, Cambridge, UK

Isle of Wight

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7.14  Isle of Wight 1.  Introduction The Isle of Wight is a diamond-shaped island measuring 37.8 km from east to west and 21.4 km from north to south, with a coastline about 90 km long. Its central feature is a ridge of steeply-dipping Chalk extending from The Needles in the west to Culver Cliff in the east. To the north are Tertiary formations, Eocene and Oligocene: the Eocene strata stand almost vertical in Alum Bay and Whitecliff Bay, immediately adjacent to the Chalk ridge, but the dip diminishes rapidly northward and the Upper Oligocene strata are almost horizontal. To the south the Chalk escarpment looks across the outcrops of Lower Cretaceous formations exposed in anticlinal zones to a southern plateau of Chalk dipping gently southward (Bird 1997). Much of the coast is cliffed across these various formations. The drainage divide is generally close to the southern coast, so that relatively long rivers drain to the north coast estuaries, as at Newtown Harbour, the Medina at Cowes, Wootton Creek and Bembridge Harbour. Shorter streams drain to the south coast, several of them having cut deep, narrow steep-sided valleys known as chines. Beaches are of sand and shingle, the proportion of shingle increasing in the vicinity of sandy cliffs, and salt marshes are found in the estuaries. Mean spring tide range is small around the Isle of Wight, where double high tides occur on the Solent and Spithead shores. The range at Cowes is 3.6 m, diminishing to 2.6 m at Yarmouth and 2.2 m in Totland Bay on the west coast. On the south coast the tide range at  Freshwater is 2.1 m, increasing to 3.1 m at Ventnor and 3.5 m at Sandown. On the east coast Brading Harbour has 3.1 m, and the largest tides are at Ryde, 3.8 m. The coast is described and illustrated in an anticlockwise sequence beginning and ending at Cowes in the north.

2.  The Isle of Wight Coastline The Medina estuary opens northward from a steep-sided, almost straight drowned valley north from Newtown to open to the sea beside Cowes on its west bank. Cowes stands on a hill where slopes in Bembridge Marls descend to Bembridge Limestone on the shore. The promenade, with a sea wall and a narrow beach of flint gravel, curves round to Egypt Point. To the west vegetated coastal bluffs rise behind Gurnard Bay. The coastal slope is a degraded

escarpment, with hummocky slopes on the Bembridge Marls and patchy outcrops of Bembridge Limestone along the shore. In Thorness Bay the wide muddy foreshore exposed at low tide has ledges of Bembridge Limestone, and is backed by sandy beach ridges on a spit that deflects the little stream eastward. Cliffs rise to the west and the beach becomes a shingle spit, capped by dunes, on the western side of Newtown Harbour. This is a ria formed by Late Quaternary marine submergence, which was followed by the growth of paired spits bordering the marine entrance. West of Newtown Harbour, cliffs and bluffs extend to Yarmouth, where the Western Yar has an estuary bordered by tidal mudflats and salt marshes with areas of Spartina grass, threaded by tidal creeks. At Sconce Point the coast turns southward, and there are cliffs cut in almost horizontal marls and limestones, past Totland Bay and beneath Headon Hill. In the northern part of Alum Bay the beach is sandy and the cliffs show chutes of slumping clay and lobes of truncated Headon Beds. The strata curve upward and are soon almost vertical as the pale yellow and grey Barton Sand emerges north of Alum Bay Chine. There follows a descending sequence of Eocene and Palaeocne strata standing almost vertical in the cliffs. The cliff face has been dissected into ravines on soft strata and ribs on harder rock in the almost vertical strata. At the southern end the basal Reading Beds dip northward at about 80°, immediately overlying eroded Upper Chalk, which emerges as steep cliffs along the southern coast of Alum Bay. The Upper Chalk cliffs run out westward, and have grassy slope segments and buttresses and a narrow basal ledge. A rock fall in 1993 produced a white scar (an exposed nearly vertical bedding plane), with a large talus apron, which was vegetated and undercut by a basal cliff in 2006: the scar was then grey (> Fig.  7.14.1). The grassy slope has been cut back from the plane of the sub-Palaeocene surface, and is not strictly a dip-slope. The Needles are stacks of very hard and splintery Upper Chalk which has become resistant largely as the result of compression and recrystallisation accompanying the intense folding. Three large stacks running out to the red and white lighthouse are shards of this chalk, about 30 m high (>Fig. 7.14.2). The Chalk cliffs run behind Scratchell’s Bay, where the northward-dipping Upper Chalk is crossed by numerous steeply inclined parallel flint horizons marking the ­bedding-planes. There is a beach of blue-black flint shingle  eroded from the cliffs, which undulate and decline ­westward to Watcombe Bay, below the steep convex grassy

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⊡⊡ Fig. 7.14.1 Slumped fan of Chalk undercut by marine erosion on the south coast of Alum Bay. (Courtesy Geostudies.)

escarpment of Tennyson Down. Watcombe Bay is a small cove cut into the head of a dry valley, and gravelly Coombe Rock in the cliff indicates that a deeper valley was excavated, then infilled with frost-shattered chalk and flint sludged down the bordering slopes under periglacial conditions in the Late Pleistocene. Cliffs cut in hard Upper Chalk descend into Freshwater Bay, a broad cove at the head of the Western Yar valley. It is similar to Lulworth Cove in Dorset, but with a wider entrance, and cut entirely in Chalk. The eastern half of the cove has a beach of flint shingle in front of cliffs cut in shattered Upper Chalk, the flint layers having been much disturbed by frost heaving and lateral sludging. This is overlain by Coombe Rock and a reddish brown Brickearth with some interbedded gravel. There are residual stacks, Stag Rock, Mermaid Rock and the foundations of Arch Rock, on the shore platform below the cliff to the east. The Chalk cliffs continue into Compton Bay, passing into the red ferruginous sands of the Upper Greensand (> Fig. 7.14.3). The shore is littered with chalk and greensand boulders. The Upper Greensand recedes behind a landslide in Atherfield Clay as the coast curves out to Hanover Point. The SW coast of the Isle of Wight has cliffs cut in Lower Cretaceous (Wealden) marls and clays with beds of sandstone forming shore ledges. They are incised by several small V-shaped chines, notably Alum Chine to the SE. An undercliff terrace develops near the cliff crest on the eastern side of Walpen Chine and widens to 130 m as it

descends eastward. Springs and seepages of groundwater from thick overlying sandstones issue from the cliff base, especially after wet weather, and this outflow has been res­ ponsible for the undermining, collapse and recession of the sandstone cliffs, as well as the consequent disappearance of Blackgang Chine. Cliffs in sandrock and sandstone stand behind a subsided undercliff (> Fig.  7.14.4) and recurrent landslides have damaged and destroyed cottages and chalets at Blackgang (Hutchinson et al. 1981). To the east the high vertical upper cliff swings seaward and the undercliff terrace comes to an end. Almost verti­­cal  cliffs in Upper Greensand extend past Rocken End. Above Rocken End the irregular topography of the Niton Landslide (> Fig. 7.14.5) ascends to the base of Gore Cliff, a vertical upper cliff in evenly-bedded horizontally stratified Upper Greensand, with protruding ledges of hard nodular cornstones and etched out shelves in softer sandstone. It is capped by the Lower Chalk, which forms a receded brow behind a wind-scoured bench. Recession of the upper cliff has interrupted the old undercliff road at Windy Corner. The steep coast east towards Ventnor consists of wooded slopes on talus and slumped rock, with basal cliffs cut into slipped masses of Chalk and Upper Greensand, often with a landward dip, indicating rotational slumping, extending past St. Catherine’s Point (Hutchinson et al. 1991). Ventnor originated as a seaside resort and spa because of a spring near the base of the slumped cliffs. Sea walls of varying design have been built eastward to Horseshoe Bay,

Isle of Wight

⊡⊡ Fig. 7.14.2 Scratchells Bay, looking towards The Needles. The sea is milky with chalk powder in suspension. (Courtesy Geostudies.)

⊡⊡ Fig. 7.14.3 Steeply dipping Chalk underlain by Upper Greensand, Gault and Lower Greensand (Ferruginous Sands) in Compton Bay. (Courtesy Geostudies.)

7.14

fronting cliffs in slumped material, some of which have been graded to grassy slopes. The Undercliff to the east is a broad strip of descending terraces of Chalk and Upper Greensand that have subsided over the Gault, overlooked by an upper cliff in which layers of Upper Greensand outcrop, were overlain by the Lower Chalk. There is a wide perched basin behind Luccombe Bay floored with Gault, and overlooked by an escarpment of Upper Greensand and Lower Chalk. Luccombe Chine is a deeply incised wooded ravine cut through the lower cliff of Sandrock Beds, and above Shanklin Point is bold Knock Cliff - the lower part showing dark Ferruginous Sands, the upper part massive grey Sandrock Beds capped by yellow bedded sandstones (> Fig. 7.14.6). From the top of Knock Cliff the ground drops across to hummocky, slumped topography on the Gault, and the cliffs decline to Appley Beach at the western end of Shanklin. The cliff, draped with vegetation, runs behind the Chair Lift at Shanklin, where the Ferruginous Sands show a gentle northward dip on the southern side of the gentle Shanklin Anticline. The coast between Shanklin and Sandown consists of almost vertical cliffs 20–30 m high in Ferruginous Sands, with basal talus covered by grasses and scrub, fronted by a sea wall and undercliff walk. Sandown is built on a plateau of Ferruginous Sands, and the coast curves gently round to the east. Sandown Bay has been excavated in the Wealden Beds which occupy the core of the unroofed Sandown Anticline. Beyond the last of the many wooden groynes at San­ down, built to prevent the sand drifting away eastward,

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⊡⊡ Fig. 7.14.4 Blackgang Undercliff. (Courtesy Geostudies.)

⊡⊡ Fig. 7.14.5 Receding cliffs in Upper Greensand above the landslide at Rocken End. (Courtesy Geostudies.)

slumping cliffs in Wealden Beds and Atherfield Clay pass into vertical cliffs in the more resistant Ferruginous Sands and the Sandrock Beds. The cliff crest declines through a hollow across the head of a dry valley that descends to Yaverland. There is a hollow on a narrow outcrop of Gault clay, followed by the Upper Greensand, somewhat ob­­ scured by slumping, and then the Lower Chalk descending from the cliff crest. Culver Cliff has been cut into the south-facing Chalk escarpment by marine erosion. At the eastern end, Whi­ tecliff Point is a vertical cliff in hard, massive Upper Chalk,

with irregular flint layers including some large tabular nodules along the bedding planes, and seams of marl. At the southern end is the jutting spur called the Anvil, formed of massive white chalk with few marl seams, its grassy northern slope following a bedding-plane which has been exposed by erosion along a weaker overlying horizon. The shore platform is strewn with chalk bould­­ ers, and shows steeply-inclined ribs that extend eastward between wave-scoured furrows, some with elongated pools at low tide. There are protruding layers and nodules of flint.

Isle of Wight

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⊡⊡ Fig. 7.14.6 Knock Cliff. (Courtesy Geostudies.)

On the southern side of Whitecliff Bay the cliff is cut into Upper Chalk, high and partly vegetated with talus slopes. It is very steep, with a large notch at the base, and a shore platform, cut across steeply dipping chalk strata and submerged at high tide. This part of the coast is much less exposed to storm waves than Alum Bay and the top of the Chalk is obscured by grassy slumped aprons. The Eocene formations in Whitecliff Bay are similar to those in Alum Bay at the western end of the Isle of Wight, showing the same sequence of vertical strata, the dip declining to the north. At the northern end of Whitecliff Bay the Bembridge Limestone appears in the top of the cliff and descends northward. Low cliffs and ledges of brown Bembridge Limestone continue along the shore, the inner and outer reefs separated by a corridor of gravelly sand which declines into a low tide channel known as The Run. Off The Foreland, the Bembridge Ledges are parallel outcrops of Bembridge Limestone. Low slumping cliffs cut in dark Bembridge Marls, capped by gravels in a clay matrix, extend NW past Bembridge. Bembridge Point is a broad, blunt dune-capped spit with a shingle beach fronted by sand on its seaward side. The Yar estuary which has been formed by submer­­gence  of a river valley cut out along the eastern part of the Bouldnor syncline widens into a bay - Bembridge Harbour. To the north the Bembridge Limestone reappears in a series of rocky spurs, between which are little beaches of sand and flint pebbles. An outcrop of hard calcareous sandstone is

responsible for the Horestone Point promontory, which declines to a shore where the Nettlestone Grit has disintegrated into large reddish-brown and ­ferruginous blocks. At Seaview shore ledges are on the outcrop of this grit, which at Nettlestone Point can be traced laterally from a shelly limestone into a hard marl, and then calcareous sandstones with flint pebbles. The coast curves out to Puckpool Point, and there are walled sectors westward to Ryde. An extensive sandy area is exposed at low tide in front of the Ryde promenade. Much of the coast between Ryde and Cowes consist of low wooded bluffs with minor basal cliffing and narrow gravel beaches, interrupted by the estuary of Wootton Creek, another drowned valley, with paired spits bordering the entrance. At Old Castle Point the sea wall and coast road begin, curving round to the mouth of the Medina estuary at Cowes.

References Bird ECF (1997) The Shaping of the Isle of Wight. Ex-Libris Press, Bradford on Avon, Wiltshire Hutchinson JN, Chandler MP, Bromhead EN (1981) Cliff recession on the Isle of Wight south-west coast. In: Proceedings of the 10th internatio­ ­nal  conference on soil mechanics and foundation engineering, Stockholm, Sweden, pp 429–434 Hutchinson JN, Brunsden D, Lee EM (1991) The geomorphology of the landslide complex at Ventnor, Isle of Wight. In: Chandler RJ (ed) Slope stability engineering. Thomas Telford, London, pp 213–218

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7.15  Sussex 1.  Introduction The western part of the Sussex coast is low-lying, bordering a coastal plain from the Hampshire border to Brighton, where it is intersected by the South Downs. Chalk cliffs extend from Brighton to Eastbourne interrupted by gaps at the mouths of the Ouse and Cuckmere valleys, and beyond Eastbourne the coast cuts across the southern part of the Wealden dome. Pevensey Levels are alluvial lowlands at the eastern end of the Sussex Low Weald (the Weald Clay vale), and were swamps and fens until they were embanked and reclaimed. At Bexhill, the Tunbridge Wells Sandstone reaches the coast in low cliffs and shore reefs, and beyond Hastings the high sandstone cliffs are cut in Ashdown Sand and Fairlight Clay in the central anticline of the High Weald. Another alluvial lowland at Winchelsea and Rye passes behind the large shingle complex at Dungeness Foreland. Mean spring tide range on the Sussex coast increases gradually eastward from 4.2 m at the entrance to Chichester Harbour to 4.7 m at Selsey Bill, 4.9 m at Pagham, 5.0 m at Bognor Regis, 5.2 m at Littlehampton, 5.4 m at Worthing, 5.5 m at Shoreham, 5.9 m at Brighton, 6.1 m at Newhaven, 6.5 m at Eastbourne and 6.8 m at Hastings.

2.  The Sussex Coast The Hampshire-Sussex boundary runs down Emsworth Channel in Chichester Harbour, to the west of Thorney

Island, which has been attached to the mainland by land reclamation. It is bordered by sea walls overlooking salt marshes and mudflats exposed at low tide (> Fig.  7.15.1), the salt marshes including much Spartina. On the southern shore the salt marsh is much dissected and eroding, and there are some large boulders of distant origin, thought to have been deposited from melting icebergs Late Pleistocene times. The streams draining into Chichester Harbour carry very little sediment, and the mud that has accreted has been derived largely from the brickearth deposits on the coastal plain. Paired spits border the entrance to Chichester Harbour, the one on the eastern (Sussex) side being East Head, a recurved sand and shingle spit surmounted by dunes up to 4 m high and backed by salt marshes and a shallow bay at high tide. The Chidham peninsula is similar to Thorney Island, and on its eastern side Bosham Channel is a tributary of the large Chichester Channel. At Bosham high tide flooding is frequent, especially at the equinoxes (> Fig. 7.15.2). East of the entrance to Chichester Harbour the coast of Bracklesham Bay has a shingle beach and numerous groynes. Wave action is limited by the short fetch across Spithead, but there are some low vertical cliffs in Brickearth at West Wittering. At Selsey Bill erosion has been rapid, but sectors are now protected by boulder ramparts. The shingle beach continues northward to the paired spits that border the entrance to Pagham Harbour. This bay has extensive salt marshes and mudflats at low tide, and receives only minor freshwater inflow from streams and drainage ditches. ⊡⊡ Fig. 7.15.1 Dissected salt marsh, Thorney Island. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.15.2 High spring tide flooding at Bosham. (Courtesy Geostudies.)

Gravelly shoals are extensive off Pagham Harbour, and to the east sandstone reefs outcrop offshore at Barn Rocks and Bognor Rocks. The seafront at Bognor Regis has been walled and armoured with dumped boulders, and numerous groynes have been built to prevent beach shingle drifting away to the east. There has been artificial replenishment of the shingle beaches at Felpham and Middleton- on- Sea, where nearshore breakwaters have been constructed. At Atherington (Climping Beach) dunes back the shingle beach and a sandy foreshore is exposed at low tide, with green weed on gravel patches and in gravelly pools. The dunes continue east to the Arun mouth, where the beach widens because of accretion alongside the harbour breakwaters. East from Littlehampton the shingle beaches continue, with many groynes, along the edge of the urbanised Sussex coastal plain to Worthing. At Shoreham-by-Sea a shingle bank, built by eastward drifting, has forced the Adur outlet to the east. The beach at Hove has been depleted as the result of interception of drifting shingle by the breakwaters at the Adur mouth. The seafront has been walled from Hove past Brighton to Black Rock, and the shingle beach retained by numerous groynes. At Black Rock the Chalk cliffs begin. There is a section showing an emerged shingle beach overlain by frost-­ shattered chalk and flint gravel (Coombe Rock) behind the Black Rock marina, built in 1974. Recession of the Chalk cliffs to the east was rapid until an undercliff walk was built in the 1930s between Black Rock and Saltdean, serving as a high tide wall. The undercliff walk was subsequently extended eastward from Saltdean, but the Chalk

cliff is still receding at Peacehaven. It is fronted by a slightly concave shore platform up to 150 m wide, with transverse gradient declining seaward below low spring tide level. This shore platform has been cut across Chalk strata that dip gently seaward, forming minor scarps and mesas, particularly on flint layers and dissected into chalk boulders. It has been cut partly by abrasion, as waves move flint fragments to and fro and partly by solution of the Chalk in rainwater and seaspray: there are often clouds of fine calcium carbonate in suspension offshore, possibly formed as dissolved chalk is precipitated. Frost shattering during cold winters has contributed to weathering on the shore platform (Robinson and Jerwood 1987) (> Fig. 7.15.3). Shingle drifting eastward has accumulated on the beach that widens towards the Newhaven breakwater (> Fig. 7.15.4) built in 1731 and later extended. Behind the harbour, the Chalk is overlain by the Reading Beds at Castle Hill, and the former cliff, protected by the breakwater and reclaimed land, has developed a distinctive profile: an upper, convex slope on soft slumping Eocene sediment, a cliff face of vertical Chalk and a lower concave slope of downwashed chalk, sand and clay. The shingle beach continues to Seaford, where it was much depleted as a consequence of the interception of eastward drifting by Newhaven breakwater. Groynes failed to retain it, and a sea wall along the promenade had to be repaired after storms. In 1987 a shingle beach was renourished here. At the eastern end of the promenade frost-shattered Chalk rises in cliffs to Seaford Head (85 m). These diminish on the eastern side towards Cuckmere Haven, at the

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⊡⊡ Fig. 7.15.3 Shingle accumulation west of the Newhaven breakwater. The white strip is scoured Chalk exposed after waves pushed back the base of the beach. (Courtesy Geostudies.)

⊡⊡ Fig. 7.15.4 Chalk cliffs and intertidal shore platform, looking east towards Beachy Head. (Courtesy Geostudies.)

mouth of a wide valley cut through the South Downs. Beyond Cuckmere Haven the Seven Sisters are Chalk ridges and dry valleys truncated by vertical cliffs up to 70 m high which have been dissected along joints and flint-lined bedding planes. The Chalk has a very gentle southward dip, truncated in broad fronting intertidal shore platforms (> Fig. 7.15.4). The cliff face has a series of prominent buttresses (> Fig. 7.15.5). At Birling Gap the 12 m high cliffs show solid Chalk on either side of up to 5 m of Coombe Rock over

f­ rost-shattered Chalk beneath a broad dry valley. Cliff recession has truncated the row of nineteenth century fisherman’s cottages, and the steep stairway down to the beach has been repeatedly replaced. To the east the cliffs rise to Beachy Head (160 m), where the escarpment of the South Downs is truncated by the coastline. Cliff recession here is by way of successive rock falls, producing aprons of Chalk rubble that are gradually consumed by abrasion and solution. A major rock fall occurred in January 1999, producing a talus fan that extended out to

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⊡⊡ Fig. 7.15.5 The Chalk cliffs of the Seven Sisters, showing prominent vertical buttresses behind an intertidal shore platform. (Courtesy Geostudies.)

the lighthouse, and has since been trimmed back several metres (Williams et al. 2004). From Beachy Head the Chalk escarpment swings inland to face eastward, and the cliffs decline from the Upper Chalk of the escarpment crest through Middle and Lower Chalk to Upper Greensand and Gault, fronted by shingle and boulders at Holywell. Sea walls line the Eastbourne waterfront, with shingle beaches contained between groynes. To the east the shingle widens to ­multiple beach ridges on a blunt lobate foreland - The Crumbles out to Langney Point. Between Eastbourne and Hastings numerous groynes retain shingle within beach compartments, often angled in response to longshore drifting and the dominant SW waves. The shingle bank is backed by the wide lowlands of Pevensey Levels, former marshes reclaimed as pastureland. At Bexhill there are low cliffs and shore reefs cut into the Tunbridge Wells Sand Formation, and the shingle beach continues along the promenades of St Leonard’s and Hastings. From Hastings past Fairlight to Cliff End are high cliffs in the Hastings Beds of the Wealden Series, exposed in the core of the Wealden dome, where the High Weald comes to the coast.