DEVELOPMENTS IN
QUATERNARY SCIENCES, 6
GLACIOTECTONISM
Developments in Quaternary Science (Series Editor: Jaap J.M. van der Meer) Volumes in this series 1. The Quaternary Period in the United States Edited by A.R. Gillespie, S.C. Porter, B.F. Atwater 0-444-51470-8 (hardbound); 0-444-51471-6 paperback) - 2003 2. Quaternary Glaciations - Extent and Chonology Edited by J. Ehlers, P.L. Gibbard Part I: Europe ISBN 0-444-51462-7 (hardbound)- 2004 Part II North America ISBN 0-444-51592-5 (hardbound)- 2004 Part II1: South America, Asia, Australasia, Antarctica ISBN 0-444-51593-3 ( h b ) - 2004 .
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Ice Age Southern Andes - A Chronicle of Paleoecological Events By C.J. Heusser 0-444-51478-3 (hardbound)- 2003 Spitsbergen Push Moraines -including a translation of K. Gripp: Glaciologische und geologische Ergebnisse der Hamburgischen Spitzbergen-Expedition 1927 Edited by J.J.M. van der Meer 0-444-51544-5 (hardbound) - 2004 Iceland - Modern Processes and Past Environments Edited by C. Caseldine, A. Russell, J. Hardard6ttir, 6. Knudsen 0-444-50652-7 (hardbound) - 2005 Glaciotectonism By J.S. Aber, A. Ber Present volume
Caption for cover image." Aquinnah Cliff on Martha's Vineyard, Massachusetts, United States. The 40-m-high cliff displays multi-colored upper Cretaceous and Tertiary strata that were upthrust along the edge of the Atlantic Coastal Plain during late Wisconsin glacier advance. Aquinnah Cliff is among the most famous and frequently visited glaciotectonic sites in the world. See chapter 6 for more details; photo by J.S. Aber (2005).
Developments in Quaternary Sciences, 6
GLACIOTECTONISM
James S. Aber Earth Science Department, Emporia State University Emporia, Kansas, U.S.A. and
Andrzej Ber Department of Quaternary Geology Polish Geological Institute Warsaw, Poland
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Foreword Between 1987 and 1993 four books with 'glaciotectonics' in the title were published; three were proceedings of meetings, one edited by James Aber, the fourth volume was an edited research volume, the senior editor of which was James Aber. There is no denying that he is an authority on the subject. Since 1993 no general books devoted to the subject have been published. Does that mean that glaciotectonics have gone out of vogue? Certainly not. Checking the reference list of the present volume one will find that numerous papers have been published on all the different aspects of glaciotectonics. In fact there are close to 200 references post-1993 in the list. Not all of these are on glaciotectonics, but neither is it the complete database. These are only the references used to produce the text of the current volume. Glaciotectonics have never been away, they are more important than ever, we were just waiting for someone to pick up the pen and write the next compilation. This has now been done. One shortcoming of all the earlier titles was that relatively little attention was paid to eastern Europe, because there were so few of us in the west who could read the work, if we could get hold of it. By linking up with Andrzej Ber, himself no stranger to studies of glaciotectonics, that extensive eastern European literature has now been made accessible. Comparing the first edition of this volume with the present one will reveal the addition the eastern European literature makes to the subject. The present text will not only be the major reference work on the subject for the years to come, it will also help in defining the next set of questions: What are the major lacunae in our knowledge of the diversity, global distribution and formational processes of glaciotectonics? What are the relations to other glacial landforms, like drumlins, or sediments, like subglacial tills.
Glaciotectonism is the first volume published in Developments in Quaternary Science since my taking over from Jim Rose as Series Editor. Jim gave the series a flying start with the first five titles, one of which consists of three volumes. So far the series is doing very well, with suggestions of reprinting and/or extending some of the first titles. Naturally there is a minor break when changing editors, but now the series is fully on track again. Currently there are seven new titles under contract, and we expect three of these to be published this year. The subjects range from interglacial climates, to early glaciological studies and from China to Patagonia. This range of topics reflects the intention of the series to consider Quaternary science across different parts of the earth with respect to a wide range of Quaternary processes. Nor do we restrict ourselves to the last 2.6 million years. It is obvious that, although that point in time is recognised as a major break, developments since then are not independent of what happened before. Thus, the volume on Patagonia will be on Late Cenozoic developments, not just Quaternary. I hope that the titles published so far and the titles to be published in the near future, will make potential authors consider this book series for their work. Jaap J.M. van der Meer Series Editor
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Contents Preface
ix
Chapter 1: Nature of glaciotectonism Historical development INQUA work groups Definition of glaciotectonism Glaciotectonic terms and concepts Glaciotectonic structures Glaciotectonic landforms Case-history approach
1 1 5 6 7 10 13 15
Chapter 2: Geometric analysis . Introduction Conventional field methods Stereographic projections Subsurface methods Remote sensing Geographic Information Systems
17 17 17 20 22 26 31
Chapter 3: Kinematic analysis . Stress and strain Balanced cross sections Micromorphology Superposed deformation Kineto-stratigraphy Glaciodynamic sequence and event
33 33 35 37 38 40 44
Chapter 4: Hill-hole pair. Introduction Wolfe Lake, Alberta Herschel Island, Yukon Devils Lake, North Dakota Norwegian continental shelf
45 45 48 50 53 55
Chapter 5: Composite ridges . . Introduction MCns Klint, Denmark Dirt Hills and Cactus Hills, Saskatchewan Flade Klit, Denmark Utrecht Ridge, Netherlands Brandon Hills, Manitoba
59 59 63 67 72 75 80
Chapter 6: Cupola hills and drumlins Introduction Aquinnah, Martha's Vineyard, Massachusetts Ristinge Klint, Denmark Hvideklint, Men, Denmark Elblgg Upland, Poland Saadj~irve drumlin field, Estonia
83 83 83 88 89 93 99
viii
Aber and Ber
Chapter 7: Megablocks and rafts . Introduction Esterhazy, Saskatchewan Southern Alberta Kvarnby, Skhne, Sweden Sinim~ied, Estonia Chapter 8: Intrusions, diapirs and wedges . Introduction Atchison, Kansas Herdla Moraine, Norway Systofte, Falster, Denmark Kronowo esker, Poland
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Chapter 12: Dynamism of Glaciotectonic Deformation Fundamental cause of glaciotectonism Initiation of thrust faulting Continuation of thrust faulting Scale models of glaciotectonism
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Chapter 13: Glaciotectonic analogs Introduction Mississippi Delta mudlumps, Louisiana Thin-skinned thrusting Convergent plate boundary References .
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141 141 142 145 147
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Chapter 11: Distribution of glaciotectonism Continental distribution of glaciotectonic phenomena Regional patterns of glaciotectonism Model for lobate pattern of glaciotectonism Glaciotectonic patterns in North America Glaciotectonic patterns in central Europe Central Europe m Weichselian Glaciation Central Europe - - Saalian and Elsterian glaciations Glaciotectonic patterns in Arctic Russia
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125 125 125 127 130 133
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Chapter 10: Applied glaciotectonics Introduction Highwall failure, Highvale coal mine, Alberta Highway construction, Maymont, Saskatchewan Diatomite quarries, Fur, Denmark
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111 111 114 117 120 121
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Chapter 9: Basement and deep crustal structures . Introduction Canadian Shield and northern Appalachians Northern Scandinavia Salt displacement, Finger Lakes region, New York Polish basement structures
Index
101 101 102 104 107 109
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153 153 160 162 167 168 171 181 187 191 191 193 196 198
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203 203 203 205 210
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2 1 5
235
Preface
This book is dedicated to our glaciotectonic mentors, Profs. Asger Berthelsen, University of Copenhagen, Denmark and Hanna Ruszczyfiska-Szenajch, University of Warsaw, Poland (figs. 1 and 2). During the 1970s and '80s they led a renaissance in glaciotectonic investigations and theory in the Baltic region of northern Europe (fig. 3). Their contributions and enthusiasm for the subject were major influences for the authors and many other geoscientists. Many other colleagues were involved with our work on glaciotectonics. In particular, we wish to acknowledge the contributions and support of the following individuals. Inge Aarseth (Norway) Ojars P. Aboltins (Estonia) David L. Ackerman (Canada) Harold V. Andersen (U.S.A.) L~erke T. Andersen (Denmark) Valery Astakhov (Russia) Peter H. Banham (United Kingdom)
Figure 1. Asger Berthelsen on the beach near his country home on the island of MOn, Denmark, where he conducted much of his glaciotectonic research. Photo by J.S. Aber (1979).
John E Bluemle (U.S.A.) Julie Brigham-Grette (U.S.A.) Krzysztof Brodzikowski* (Poland) Bruce E. Broster (Canada) Chris D. Clark (United Kingdom) David G. Croot (United Kingdom) Louis F. Dellwig (U.S.A.) Linda A. Dredge (Canada) Alexis Dreimanis (Canada) Mark M. Fenton (Canada) Darek Gat~zka (Poland) Jane K. Hart (United Kingdom) Stephen R. Hicock (Canada) Howard C. Hobbs (U.S.A.) Michael Houmark-Nielsen (Denmark) Ole Humlum (Denmark) Mads Huuse (Denmark) William R. Jacobson, Jr. (U.S.A.) Peter Roll Jakobsen (Denmark) Wojciech Jaroszewski* (Poland) Carrie Jennings (U.S.A.)
Figure 2. Hanna Ruszczyfiska-Szenajch examines a freshly dug exposure at Kadyny in Elbl~g Upland, one of her many field sites for glaciotectonic investigations in Poland. Photo by J.S. Aber (1993).
x
JCrn Bo Jensen (Denmark) Flemming JCrgensen (Denmark) Volli Kalm (Estonia) Alexander Karabanov (Byelrussia) John S. Klasner (U.S.A.) Rudy W. Klassen (Canada) Alan R. Knaeble (U.S.A.) Stefan Kozarski* (Poland) Johannes Krtiger (Denmark) Dariusz Krzyszkowski (Poland) Erik Lagerlund (Sweden) E.A. Levkov* (Belarus) Tom van Loon (Netherlands) Ida LCnne (Norway) Kam Lulla (U.S.A.) Jan Lundqvist (Sweden) Holger Lykke-Andersen (Denmark) Jan Mangerud (Norway) Leszek Marks (Poland) Irene Marzolff (Germany) Andrei V. Matoshko (Ukraine) Jaap van der Meer (United Kingdom) Rod A. McGinn (Canada) David M. Mickelson (U.S.A.) Wojciech Morawski (Poland) Daniel Nyvelt (Czech Republic) Stig A. Schack Pedersen Jan Piotrowski (Denmark) Mikko Punkari (Finland) Maris Rattas (Estonia) Bertil Ringberg (Sweden) Karol Rotnicki (Poland) Joar Saettem (Norway) David J. Sauchyn (Canada)
Aber and Ber
Nicholaus M. Short (U.S.A.) Steen SjCrring* (Denmark) Everett E. Spellman (U.S.A.) Hans-Jtirgen Stephan (Germany) James T. Teller (Canada) William A. Thomas (U.S.A.) F.M. (Dick) van der Wateren (Netherlands) William A. White (U.S.A.) Richard S. Williams, Jr. (U.S.A.) Vitalijs Zelcs (Latvia) * deceased. Financial support for our efforts has come from many sources, not the least our home institutions, Emporia State University, Kansas (Aber) and the Polish Geological Institute, Warsaw (Ber). A series of grants from NASA provided Aber with experience and tools for remote sensing and GIS applications in the geosciences. Aber's experience in Estonia was funded by the U.S. National Research Council, and his field work in Poland was supported by the U.S. Council for Intemational Exchange of Scholars. INQUA provided material aid to Ber in connection with the glaciotectonic mapping project in central Europe. Other sources of support include the North Dakota Geological Survey, Danish Natural Science Research Council, and William T. Kemper Foundation - Commerce Bank, Trustee - Kansas City, Missouri. Thanks to the Swiss Geological Society for permission to reproduce copyrighted illustrations. Finally we wish to thank our wives, Susie and Maria, who patiently assisted us during preparation of this book and encouraged our long-term commitment to understanding glaciation and its role in shaping the Earth.
Figure 3. Strongly sheared and refolded chalk-till m~lange exposed in cliff near Hvideklint on the island of Men, southeastern Denmark. Small spade to right for scale. Photo by J.S. Aber (1979).
Chapter 1 Nature of Glaciotectonism Historical development Glaciotectonism is an important component of modern glacial theory, but it gained widespread recognition only within the past 25 years. Glacial theory began to develop in the late 18th and early 19th centuries in the Alps of western Europe and the mountains of southern Scandinavia. Horace-Bdn6dict de Saussure was among the earliest naturalists to undertake systematic observations of glaciers in the Mont Blanc vicinity beginning in the 1760s. He observed types of glaciers, ice flow, origin of moraines, and many other aspects of glaciers. Saussure introduced the terms roches moutonn6e, s6rac, and moraine into geological usage (Carozzi and Newman 1995). James Hutton was first to recognize in 1795 that erratic granite boulders in the Jura Mountains had been transported by glaciers from the Alps (Flint 1971). In Scandinavia, Jens Esmark concluded in 1824 that glaciers once had been much larger and thicker, and had covered much of Norway and the adjacent sea floor (Andersen 1992). He attributed erratics and moraines to glacial transportation and deposition. Esmark also recognized that glaciers were powerful agents of erosion that had carved out the Norwegian fjords (Cunningham 1990). In the early 1830s, Jean de Charpentier began to marshall the scientific evidence for former alpine glaciation. He observed moraines, striations, and erratics, as well as existing glaciers of the Swiss Alps, and presented his conclusions in a persuasive manner (Teller 1983). Although a disbeliever at first, Louis Agassiz came to accept the concept of former alpine glaciation, and then carried the idea much further. In 1837 he proposed that vast sheets of ice had once covered much of the northern hemisphere. This was a radical suggestion, for at that time the modem ice sheets in Greenland and Antarctica were completely unknown. Agassiz undertook detailed studies of glacier movement on the Unteraar Glacier in Switzerland in the 1840s, and he influenced James D. Forbes to begin similar glaciologic research in the French Alps. Forbes established that glaciers move in part by internal viscous (plastic) deformation, in contrast to the more popular dilatation or regelation theories of the day (Cunningham 1990). The glacial theories of Esmark, Charpentier, Agassiz and Forbes were based on recognition of three features--large erratic boulders, moraines, and abrasion marks on boulders and bedrock, interpreted in light of observations on glacier dynamics. Taken together these features could be explained only as the results of formerly more-extensive glaciation.
Glacial theory from the beginning, thus, rested on two groups of geological field evidence: 1) features formed by glacial erosion and 2) features created by glacial deposition. The possibility that glaciers could deform shallow crustal rocks and sediments was not recognized until a few decades later. Charles Lyell (1863) was among the first geologists to discuss the origin of contorted glacial strata at Norfolk, England, in the Italian Alps, and elsewhere. He suggested three possible mechanisms for deformation: 1) pushing by stranding icebergs, 2) melting of buried ice masses, and most importantly 3) pushing before advancing glacier ice. Lyell's The Antiquity of Man (fig. 1-1) was widely read, and his comments on glacially deformed sediments must have influenced many geologists. During the late 19th century, ice-pushed structures were recognized in several now classic locations" Sk~ne, southernmost Sweden (Torell 1872, 1873; Erdmann 1873); MCns Klint, Denmark (fig. 1-2) and Rtigen, Germany (Johnstrup 1874); and southern New England islands in the United States (Merrill 1886a). The obvious structural disturbances at these places previously had been ascribed to a variety of causes, including landslide, volcanism or intrusion, mountain building, etc. Recognition of ice-shoved features from interior continental locations came about in the latest 19th and earliest 20th centuries. Gilbert (1899) noted glacier-deformed bedrock structures on the southern shore of Lake Ontario, New York. Various ice-pushed structures were described near Minneapolis, Minnesota (Sardeson 1905, 1906) and in Poland (Frech 1901, 1915). Hopkin's (1923) analysis of large iceshoved hills in eastern Alberta is perhaps one of the best early studies. He emphasized the similarity in structural style of these hills compared to the foothills of the Canadian Rocky Mountains. The first geologist who really specialized in the study of glacially deformed structures was George Slater, an Englishman. He chose the subject of ice-push deformation for his doctoral dissertation at the University of London. He studied ice-shoved hills in England (Slater 1927a, 1927b), Denmark (Slater 1927c, 1927d), Canada (Slater 1927e), the United States (Slater 1929), and the Isle of Man (fig. 1-3; Slater 1931). He was the first to use the term glacial tectonics (Slater 1926), which is now generally shortened to glaciotectonics (American) or glacitectonics (British).
2
Aber and Ber by observations of neoglacial push-moraines on Spitsbergen made by the German geologist Gripp (1929). However, when Gry moved to a position at the Geological Survey, teaching of glaciotectonics ceased at the University, and the study of Danish glaciotectonics languished for many years thereafter.
OF
THE ANTIQUITY OF MAN WITR KKKhREB ON T K l O R I B 8 0 l p
THE ORIGIN OF SPECIES BY VARIATION
BY SIR CHARLES LYELL, F.R.S. .LUTKOIt 0Y ' P ~ I K C I P L E 8 0 l t GKOLOG~," ' KLKMKNTS 0 1 GEOLOGY," ITC. ITC.
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Figure 1-1. Cover page of Lyell's "The Antiquity of Man" (1863), which contains an early discussion of contorted drift.
Unfortunately, the Danish geologist Axel Jessen, who had once helped Slater, later accused Slater in no uncertain terms of plagiarizing and misrepresenting his (Jessen's) work at LCnstrup Klint, Denmark (fig. 1-4; Jessen 1931). In spite of Slater's prolific publications, glaciotectonism was still considered an unusual or peculiar manifestation of the glacial theory. Small contortions in glacial strata were commonly recognized, but many geologists continued to deny that glaciation could create large deformations or build substantial ice-pushed landforms. The growth of glaciotectonic research in Europe during the middle portion of the 20th century was checkered by distinct national differences. The study of glaciotectonic phenomena emerged briefly as a specific field of research in Denmark during the 1930s under the leadership of Helge Gry (Univ. Copenhagen) and Jessen (Geological Survey Denmark). Gry's (1940) analysis of ice-shoved hills in the Limfjord region of northwestern Denmark demonstrated the full potential of combining structural geology with glacial stratigraphy and geomorphology. Gry was strongly influenced
Before World War II, glaciotectonic research in what is now Poland was carried out by German geologists (Bergrat and Illner 1928; Woldstedt 1932; Meister 1935; Berger 1937; Schwarzbach 1942) as well as Polish geologists (Lewifiski and R6~ycki 1929; Czajka 1931; Rr~ycki 1937). Glaciotectonics next emerged as an important subject of research in the Netherlands following World War II (Crommelin and Maarleveld 1949; de Jong 1952; Maarleveld 1953). The reason for Dutch interest in glaciotectonics is obvious: ice-shoved ridges are the most conspicuous topographic features in a country otherwise noted for its flatness. The Netherlands continues to be a center for glaciotectonic research, as shown by numerous detailed investigations (e.g. Ruegg and Zandstra 1981; van der Meer 1987; van der Wateren 1992), and Dutch scientists have played a leading role in Svalbard glaciotectonics (van der Meer 2004). Glaciotectonics continued as a significant field of research in Poland (Jahn 1956; Dylik 1961; Galon 1961; Krygowski 1965; Rotnicki 1976) during this period. Glaciotectonics is among the most active research fields in Polish glacial geology today (Ruszczyfiska-Szenajch 1979; 1985; Ber 1987, 1999; Jaroszewski 1991, 1994; Krzyszkowski 1996). Icepushed structures are increasingly important for overall Quaternary geology (Brodzikowski and van Loon 1985, 1991; van Loon and Brodzikowski 1994). Meanwhile, research in northern Germany reflected a renewed interest in ice-push deformation there (e.g. Hannemann 1970; Grube and Vollmer 1985; Stephan 1985; van der Meer 1987), and similar efforts were underway in Belarus of the Soviet Union (e.g. Lavrushin 1971; Levkov 1980; Karabanov 1987). Glaciotectonic research in Denmark was revived during the 1970s by Asger Berthelsen (Univ. Copenhagen), who applied his experience with hard-rock structural geology to unraveling glaciotectonic phenomena. He developed the method of kineto-stratigraphy (Berthelsen 1973, 1978), in which the main emphasis is placed on the study of the directional elements that reflect the movement patterns (kinetics) of former ice sheets (1978, p. 25). Berthelsen motivated students, and his kineto-stratigraphic method was highly successful for working out Weichselian glacial stratigraphy in Denmark. In fact, glaciotectonic analysis is now an integral part of geological mapping in Denmark (Petersen 1978) and plays an important role for Quaternary studies (Houmark-Nielsen 1988, 1994, 1999; Pedersen 2000).
Nature of glaciotectonism
3
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Figure 1-2. Earliest known illustration of MOns Klint, from Pontoppidan's "Danske Atlas" (1764). Chalk mass in center, Sommerspiret (B), stands in a vertical position with the pinnacle >100 m above sea level.
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Figure 1-3. Structural sections through the Bride Moraine at Shellag Point, Isle of Man, United Kingdom. Above - section from Slater (1931). Below - section measured by Thomas (1984). Note similarities in these sections described more than 50 years apart. Adapted from Thomas (1984, fig. 3). Reproduced from Boreas by permssion of Taylor & Francis AS.
4
Aber and Ber
is approximately twice that of combined Fennoscandian/ British Ice Sheet coverage (Flint 1971). Beginning in the 1950s, large-scale topographic maps and aerial photographs became widely available, which greatly facilitated geologic reconnaissance. At the same time, many states and provinces commenced systematic geologic mapping programs. The North American glaciotectonic renaissance began in western Canada on two fronts. Working in southern Saskatchewan, A.R. Byers (Univ. Saskatchewan) demonstrated the ice-thrust genesis of the Dirt Hills and Cactus Hills. These are among the largest and best-developed glaciotectonic hills in the world (Byers 1959). Walter Kupsch (Univ. Saskatchewan) expanded on Byers' work to include similar large ice-shoved hills across southern Saskatchewan and eastern Alberta (Kupsch 1962). At the Saskatchewan Research Council, reconnaissance mapping, done mainly by Earl Christiansen, resulted in much additional knowledge of glaciotectonic features (Christiansen 1961, 1971a, 1971b; Parizek 1964; Christiansen and Whitaker 1976).
Figure 1-4. LOnstrup Klint, North Sea coast of northwestern Denmark. This cliff erodes rapidly, retreating at a rate > 1 m per year. A - northward overview of cliff section which exceeds 80 m height and extends several km. B - detail of tilted Quaternary strata exposed in the cliffface near Rubjerg Knude; north toward left. A complete glaciotectonic analysis of this classic cliff section was completed recently by Pedersen (2005). Photos by J.S. Aber (2004).
Following Slater's controversial career, British interest in glaciotectonics was minimal. The classic chalk rafts and contorted drift along the Norfolk coast received some attention in the 1960s (Peake and Hancock 1961; Harland, Herod and Krinsley 1966), and a revival in glaciotectonic studies there was led by Peter H. Banham (Univ. London). Banham (1975, 1977) also applied hard-rock geologic principles to interpretation of glaciotectonic structures. Banham's work complemented that of Berthelsen along with the Soviet geologist Lavrushin (1971). Together, they built the methodologic foundation for modem glaciotectonic research. Aside from several early studies on southern New England islands (Hollick 1894; Woodworth 1897; Upham 1899)and a few other isolated investigations, little glaciotectonic research was carried out in North America until the late 1950s. Part of the explanation may be the fact that fewer glacial scientists were faced with a much larger and less developed geographic area to investigate. The glaciated region covered by the Laurentide Ice Sheet in North America
At the same time, equally impressive ice-shoved ridges on the Yukon Coastal Plain were described by J.R. Mackay (1959). These features, developed wholly within permafrost, were the basis for Mackay and colleague W.H. Mathews (both Univ. British Columbia) to develop a general theory for glaciotectonic thrusting (Mathews and Mackay 1960; Mackay and Mathews 1964). They were influenced strongly by Hubbert and Rubey's (1959) analysis of overthrust faulting in mountains. The Saskatchewan discoveries soon spilled over into North Dakota, where a county mapping program was underway. Led by Lee Clayton (Univ. North Dakota), Steven R. Moran and John L. Bluemle (both North Dakota Geological Survey), the state became a productive center for glaciotectonic research during the 1960s and '70s. Many local studies were conducted, culminating in a new Geologic Map of North Dakota (Clayton, Moran and B luemle 1980), on which a variety of glaciotectonic landforms is shown. Theoretical analysis accompanied the field mapping (Moran 1971; Clayton and Moran 1974; Bluemle and Clayton 1984). When Moran moved to the Alberta Research Council, the geographic area was expanded to include the entire glaciated Great Plains region of the United States and Canada. The result was the first attempt at continent-scale synthesis concerning ice-sheet dynamics, distribution of glacial landforms, and genesis of glaciotectonic phenomena (Moran et al. 1980). Clayton later moved to the Wisconsin Geological Survey, with the unsurprising result that glaciotectonic features are now recognized in that state. With increasing surface and subsurface information, it now appears that glaciotectonic phenomena are ubiquitous in glaciated regions underlain by sedimentary bedrock or thick
Nature of glaciotectonism
"'"~
DEFORMATION
5
+1
Figure 1-5. Triad of effects created by glaciation, on which modern glacial theory is based.
drift (Moran 1971; Moran et al. 1980) and are even common in thin drift resting on crystalline bedrock. A variety of distinctive landforms is now attributed either wholly or partly to glaciotectonic genesis. Hence, glaciotectonic features must be included with depositional and erosional features as primary field evidence for glaciation. The modern glacial theory is, therefore, supported by a triad of field evidence, including erosional, deformational, and depositional features (fig. 1-5). Although hardly recognized in older glacial geology textbooks (e.g. Flint 1971; Drewry 1986), glaciotectonism has achieved a prominent place since the 1990s (e.g. Hambrey 1994; Benn and Evans 1998; Evans 2003).
INQUA work groups International recognition of the significance of glaciotectonic phenomena came in 1982, when a Work Group on Glacial Tectonics (WGGT) was established within the International Union for Quaternary Research (INQUA). Berthelsen organized WGGT and served as its first President. The overall goals of WGGT were: to initiate and stimulate studies of glaciotectonic phenomena in both recent and ancient glacial environments, to promote interdisciplinary collaboration between scientists working in different parts of the field, and to increase glaciotectonic curriculum in academic teaching and professional training (WGGT Newsletter, 3/1987). Leadership of the work group was taken over by David G. Croot (Plymouth Polytechnic, United Kingdom) in 1987, who was succeeded in turn by Jane K. Hart (Univ. Southampton, United Kingdom) in 1995. WGGT's formal existence came to an end in 1999, when its activities were incorporated into other work groups of the INQUA Commission on Glaciation.
Figure 1-6. WGGT work group field conference at MCns Klint, Denmark. During a lunch break, A. Berthelsen (standing to right) points out a feature in the cliff exposure. Photo by J.S. Aber (1986).
During its existence, many activities and publications were sponsored under the WGGT umbrella (fig. 1-6). Special symposia and field conferences were conducted (e.g. Sjcrring 1985; van der Meer 1987; Croot 1988a; Aber 1993a; Sauchyn 1993; Warren and Croot 1994), and glaciotectonic examples have been used for laboratory exercises in structural geology (Aber 1988a). The importance of glaciotectonic structures and landforms is now well accepted among many researchers actively engaged with studies of glacial geology and geomorphology. One of the original goals of WGGT was to assemble a comprehensive bibliography of references to published reports, books, maps and other documents pertaining to glaciotectonics. This effort was pursued through a series of publications (Aber 1988c, 1993b), and with the emergence of Internet the glaciotectonic bibliography moved online in the late 1990s [http://www.geospectra.net/]. Another major goal of WGGT was to assemble regional and continental geographic databases on glaciotectonic phenomena. This effort was initiated at a work group meeting on the island of Men, Denmark in 1986 and formally adopted at the INQUA Congress in Canada (1987). Project organization began immediately in North America. At that time, a controversial decision was taken to pursue development of the map and database using Geographic Information System (GIS) methodology. In the late 1980s, GIS still was considered experimental and risky for application to such a large and complex undertaking. David J. Sauchyn and David L. Ackerman led construction of the GIS database at the University of Regina (Canada).
6 Primary compilers for glaciotectonic data were B luemle, Lynda Dredge (Geological Survey Canada), Julie BrighamGrette (Univ. Massachusetts) and Aber. They adopted a morphological approach for classifying and mapping glaciotectonic phenomena (Aber 1988b). In many cases, glaciotectonic features in North America are known mainly or only from their morphologic expression, so structural classification schemes proved ineffective.
Aber and Ber of Stig A. Schack Pedersen (Geological Survey Denmark) and then by Croot (Croot and Michalak 1993). It became apparent that compiling a European database on glaciotectonic phenomena would be more complicated than in North America. Agreement on how to proceed was more difficult to achieve among various countries with different scientific traditions in mapping glacial geology and geomorphology. Diverse points of view and working methods hampered initial progress.
An initial map and database for the northern Great Plains region was completed in 1991 (Aber et al. 1991). This first effort provided a template upon which to elaborate the whole glaciated region of North America. A preliminary dataset for the whole continent was achieved two years later (Aber et al. 1993), and a final map was published just in time for the INQUA Congress in Berlin, Germany (Aber et al. 1995). All the GIS datasets were placed online and remain available for anonymous FTP from the Data Access and Support Center (DASC) at the Kansas Geological Survey [http:// gisdasc.kgs.ku.edu/].
The demise of the Soviet empire in the early 1990s was a foremost event for integrating geological research in western, central, and eastern Europe. The emergence of newly freed countries and their scientific establishments gave renewed impetus for the glaciotectonic mapping project. Russian geologists collaborated with scientists from the west for glaciotectonic investigations, for example Valery Astakhov and Jan Mangerud (Norway). Similar transnational cooperations developed between many other European and North American countries.
Success of the GIS approach for the North American glaciotectonic map led to creation of a new work group at the Berlin congress. A work group on Geospatial Analysis of Glaciated Environments (GAGE) was approved as part of the INQUA Commission on Glaciation. GAGE was organized by Aber, who became its first president, and later he was joined by Ber as co-president. This work group continued the goal of assembling GIS datasets and producing maps of glaciotectonic structures and landforms (fig. 1-7). GAGE remained quite active until the INQUA Congress in the United States (2003), when GAGE officially completed its mission.
At the INQUA Congress in Berlin in 1995, Ber assumed the role as main coordinator of the glaciotectonic mapping project for central Europe with assistance from many individuals. The resulting glaciotectonic database spans the Baltic region from Denmark to Estonia and southward to the Ukraine and Czech Republic. Glaciotectonic maps have been prepared for individual countries (Jakobsen 2003; Ber 2003a; Ber and Krzyszkowski 2004; Rattas and Kalm 2004), and a first version of the glaciotectonic map of central Europe was presented in 2003 at the INQUA Congress in the United States (Ber 2003b; Ber and Aber 2003).
Meanwhile, organization of the European glaciotectonic mapping project began to develop, first under the leadership
Definition of glaciotectonism Glaciotectonism refers to the processes of glaciotectonic deformation. However, considerable uncertainty surrounds the exact meaning and use of the term glaciotectonic, and confusion is both semantic and conceptual. Semantically it is both an adjective and noun. In English, the word glaciotectonic is an adjective that should be used to modify a noun, for example glaciotectonic landform. In noun forms, glaciotectonics refers to features or results of glaciotectonic deformation, whereas glaciotectonism refers to the processes of glaciotectonic deformation.
Figure 1-7. GAGE work group field conference, Slovakia. Group prepares to take a raft tour down the Dunajec River in the Tatra Mountains. Juraj Janocko (standing to right) was the conference organizer. Photo by J.S. Aber (2000).
Conceptual confusion arises because various deformed structures are common both in glacier ice and in glacial deposits. Although Slater (1926) did not define the phrase glacial tectonics, he used it in reference to structural disturbances in both drift and bedrock as well as deformations in glacier ice. Similar, if somewhat vague, references have been given by other geologists in the years since. An important review by Occhietti (1973) included five categories of glacially related deformation.
Nature of glaciotectonism •
• • •
•
- deformation of pre-existing substratum (drift and bedrock) by active ice movement. glaciodynamic- primary structures (such as till fabric) produced within ground moraine by active ice. glaciostatic - deformation of ground moraine and substratum by static ice loading. g l a c i o k a r s t i c - deformation accompanying freezing and thawing of buried dead-ice masses. i c e b e r g d r i f t i n g - deformation of sea or lake sediments by grounded icebergs. glaciotectonic
The latter two categories may be eliminated from further consideration, leaving threeAglaciotectonic, glaciodynamic, and glaciostatic, which are the primary topics of this book. is an all-encompassing term defined here as glacially induced structural deformation of bedrock or sediment masses as a direct result of glacier-ice movement or loading (based on Moran 1971; Aber, Croot and Fenton 1989). Maltman, Hubbard and Hambrey (2000) adopted a quite similar definition with the further stipulations that structures within glacier ice are not included and primary depositonal structures (such as till fabric) are excluded. Glaciotectonics
The word "bedrock," as used in this definition, implies any type of pre-existing consolidated or unconsolidated materials ranging from crystalline basement rock, to lithified sedimentary rock, to unlithified or loose sediment; in other words, the substratum and foreland over which the glacier advances. Glaciotectonic deformation also may involve sediments deposited during the same glaciation that subsequently deformed those strata. A glacier or ice-sheet may induce deformations of substratum or foreland m~terial as a result of its forward (dynamic) movement or its vertical (static) loading. Both causes of deformation often operate simultaneously, and the effects of each cannot usually be separated. Therefore, glaciodynamic and glaciostatic structures are considered as joint manifestations of secondary deformations produced during glaciation. Depression and rebound of the crust and lithosphere, as a consequence of glaciation and deglaciation, are forms of glaciotectonism, long recognized and much studied (Walcott 1969; Andrews 1970; Tushingham 1992). However, the emphasis in this book is placed on shallow (<1 km deep) structures and landforms. Lithospheric and crustal depression and rebound are considered here insofar as they may bear on the genesis of shallow crustal structures and seismicity (Thorson 2000). The identification of glaciotectonic structures and landforms is based on two fundamental criteria: 1) presence of recognizable, pre-existing bedrock or sediment bodies and 2) presence of glacially induced deformations within those bodies. Deformed structures, such as folds, faults, breccia,
7 mylonite, cleavage, slickensides or other disturbances, may be produced essentially in situ (autochthonous) or during the transportation and deposition of a detached mass (allochthonous). Such deformations must be the result of glacially imposed stresses and are not merely pre-existing structures inherited from the parent material. Glaciotectonism can be viewed in some cases as a kind of incomplete or partial glacial erosion of bedrock. Conversely, it can also be considered as a form of glacial deposition in other cases. However, neither approach is entirely satisfactory; glaciotectonic features represent those situations where glacier ice is capable of deforming pre-existing strata without completely removing or destroying the rock or sediment beyond recognition. Whether glaciotectonic features may develop is controlled to a large extent by the physical conditions of the substratum material. The same dynamic conditions that elsewhere cause conventional glacial erosion and deposition create glaciotectonic features in appropriate kinds of bedrock or glacial sediments. Glaciotectonic
terms and concepts
Contrasting glaciotectonic concepts and terminology have varied among geologists and geographers over the years. Many ideas and terms have been proposed, some of which have proven useful while others have led to confusion, as noted below. Ruszczyfiska-Szenajch defined glaciotectonics as glacial deformations of unconsolidated sediment or lithified rock of the glacier substratum and the immediate foreland, caused by mechanical action of the glacier, that is by dynamic or static pressure, or by ice movement (translated from Polish, 1983, p. 503). She restricted glaciodynamics to processes operating within the glacier and made a distinction between materials dislocated by deformation and materials transported by freezing onto the underside of the glacier. Van der Wateren (1992, p. 9) defined glaciotectonics simply as deformations of the upper horizon of the lithosphere caused by glacial stresses. He qualified this broad definition by excluding crustal depression and rebound, reactivation of large crustal faults, deformations within glacier ice, and glacial erosion. He specifically included glaciodynamic deformations within till. Houmark-Nielsen (1988, p. 92) stated that glaciotectonic structures are glacier-induced tectonic structures developed above a surface of dgcollement in soft sediments in connection with the advance, overriding, and decay of ice sheets. He included proglacial, subglacial and supraglacial zones of deformation, and he restricted glaciodynamics to the ice mass (fig. 1-8). A basic distinction in structural geology is the difference between primary and secondary structures. P r i m a r y s t r u c t u r e s are ones that originate during the formation of sediments before they become lithified; some primary
8
Aber and Ber I
GLACIER ADVANCE
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Figure 1-8. Schematic diagram of glaciotectonic and glaciodynamic phenomena, including proglacial, subglacial and supraglacial zones of deformation. Notice the migration of these zones through time during a glacier advance. Takenfrom Houmark-Nielsen (1988, fig. 2). structures originate from deposition and others are created by deformation (Davis and Reynolds 1996). S e c o n d a r y structures are those created later, when the lithified rock suffers strain as a consequence of changing stress conditions. This concept is relatively straighforward where preQuaternary rocks are deformed by Quaternary glaciation, but is difficult to apply to many Quaternary strata. In the most problematic case, subglacial deposition and deformation may be almost synchronous or penecontemporareous, so that a separation between the two processes is nearly impossible. A distinction has emerged between glaciotectonic terminology derived from structural geology, on one hand, and s e d i m e n t o l o g y on the other (Table 1-1). The term "glaciotectonite" best represents the structural school. On the sedimentology side, "deformation till" is the most widely used term. To deal with the dual structural and sedimentary nature of many glaciotectonic structures, Banham (1975, 1977) introduced and elaborated the terms endiamict, exodiamict and glaciotectonite. •
•
•
primary, usually small structures including tight to isoclinal, sheared-out folds, axial-plane foliation and pebble lineation (fabric) that are developed in till penecontemporaneous with its deposition. Equivalent to "domainal" structures (Berthelsen 1978).
geologists and are illustrated further in Fig. 1-9. The principal lithological boundary is between till (D - endiamict) and penetratively deformed bedrock (C - exodiamict). In contrast, the principal structural boundary is between penetratively deformed materials (C - glaciotectonite) and nonpenetratively deformed bedrock (B). Banham (1977) argued that the structural boundary is most important, as it lies at the lower limit of penetrative deformation within the whole sequence. He cited several practical reasons for this position. • Penetrative structures may be identified and measured readily in the field. • A d6collement often separates the penetratively deformed allochthonous rock body above from non-penetratively deformed autochthonous substratum. • Penetrative structures in bedrock are closely related to those in overlying till; they both formed in response to the same stress regime. • Penetrative deformation of mixed bedrock could produce material that is similar to or transitional in character with overlying till.
Endiamict-
- secondary, usually large structures comprising all types of folds, thrusts and other faults developed within sequences of glaciogenic sediment or bedrock. These may be the same age as endiamict structures in overlying till. Equivalent to "extra-domainal" structures (Berthelsen 1978). Exodiamict
Glaciotectonite - tectonite created by glacier action that imposed planar and/or linear fabric on pre-exisiting rock. Penetrative structures (cleavage, foliation or lineation) form the dominant fabric within the deformed rock or sediment body at the scale of observation (hand specimen).
These terms have achieved wide acceptance among glacial
Table 1-1. Terminology applied to glaciotectonic structures and sediments. Group A is based primarily on concepts from structural geology; group B reflects ideas from sedimentology. Both groups share the usage of petrographic terms, such as mylonite. Based on Kozarski and Kasprzak (1994). Terminology
Preferred terms
Group A Structural Criteria
Group B Sedimentologic Criteria
Glaciotectonite Endiamict glaciotectonite Exodiamict glaciotectonite Glacial mylonite
Deformation till Communition till
Glaciotectonic breccia Glaciotectonic boudinage Intra-domain deformation Extra-domain deformation Sand-mylonite striation Boudinage (torpedo) structure Glaciodynamic melange G lacio-cataclasite
Ground moraine dynamic stratification Gneiss-like ground moraine Glaciomylonite Shear till
. . . . . . . .
Other terms
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Nature of glaciotectonism
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Figure 1-9. Schematic sections in till and bedrock• A - undeformed bedrock, B - weakly deformed bedrock (exodiamict), C - penetratively deformed bedrock (exodiamict), D - lodgement till (often endiamict). 1. principal structural boundary, 2. principal lithologic boundary. Adapted from Banham (1977, fig. 2). Reproduced from Boreas by permssion of Taylor & Francis AS. As introduced and elaborated by Elson (1961, 1989), deformation till is defined as weak rock or unconsolidated sediment that has been detached from its source, the primary sedimentary structures distorted or destroyed, and some foreign material admixed. The only significant difference between this definition and glaciotectonite is the presence of "some foreign material" within the deformation till. Deformation till is created in the subglacial zone of shearing by intermingling of locally deformed substratum material that has moved only a short distance with sparse erratics transported from afar. The idea of a deforming glacier bed has become quite popular (e.g. Boulton and Jones 1979; Boulton 1987; Hart 1994), and has been demonstrated beneath ice streams of modern ice sheets (Alley et al. 1986, 1987; Rooney et al. 1987). Deformation till is postulated to extend over tens of thousands of klTl2 across broad lowlands and plains (Alley 1991; Hicock and Dreimanis 1992). Van der Meer, Menzies and Rose (2003) have taken the matter one step further by introducing the term teetorniet and claiming that all subglacial tills are formed by tectonic activity. Hart (1998) drew a distinction between the deforming layer below and in front of a glacier and deformations within the debris-rich basal ice of the glacier (fig. 1-10). The deforming layer has high potential for
preservation in the geologic record, whereas the latter is rarely preserved (Dreimanis 1993). The debris-rich basalice layer serves to store and transport sediment which could be fed via melting into the deforming layer below or in front of the advancing glacier. Thus, a close genetic relationship is possible between these two zones of deformation. The concept of deformation till has proven quite attractive and useful, but it also has generated considerable controversy. According to a well-known definition, till is a sediment that has been transported and is subsequently deposited by or from glacier ice, with little or no sorting by water (Dreimanis and Lundqvist 1984; Dreimanis 1989). While the role of glacier drag is recognized, the emphasis of this definition is on glacial sedimentation. Ruszczyfiska-Szenajch (1998, 2001) argued that deformation of pre-existing materials cannot produce new sediment. Such deformation may bring about substantial structural changes, including severe strain and dislocation of the rock or sediment bodies, but it does not generate new sedimentary deposits. Hence, she argued the term deformation till is misleading and liable to be misunderstood. Pedersen (1989) and Kozarski and Kasprzak (1994) likewise recommended to avoid the confusing term deformation till, because deposition and deformation are different processes.
debris-rich basal ice layer
Figure 1-10. Schematic illustration of two zones of glacial deformation. The deforming layer includes the substratum below and immediately in front of the glacier. Deformation also takes place in debris-rich basal ice of the glacier. Adapted from Hart (1998, fig. 2); reprinted with permission from Elsevier.
10
Aber and Ber Table 1-2. Listing of ductile (fold) and brittle (fracture) structures commonly e n c o u n t e r e d in g l a c i o t e c t o n i c disturbances. Based in part on Brodzikowski and van Loon (1985), Dreimanis (1993), Kozarski and Kasprzak (1994), and others. All terms can be given a glaciotectonic connotation by adding the adjective "glacial" or the prefix "glacio" or "glaci."
TILL --
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Ductile
Figure 1-11. Idealized sketch of structural zones formed beneath a glacier. Exodiamict glaciotectonite: 1 -jointing and fracturing, 2 - brecciation and shearing, 3 - shearing and grinding. Endiamict glaciotectonite (4) - local, deformation or lodgement till varieties. Adapted from Pedersen (1989, fig. 2).
As the foregoing discussion suggests, certain terms applied in glaciotectonics have controversial meanings, and considerable disagreement exists as to their proper use (or misuse). This book is not intended to settle such debates, which often serve to sharpen thinking and arguments concerning glaciotectonism. In general, the terminology of Banham (1975, 1977), as noted above, is employed in the following chapters. However, other terms and concepts will be utilized wherever appropriate in discussions of glaciotectonic landforms, structures, and processes.
I. ....................
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thrust and reverve faults normal fault overthurst fault imbrication joints and fracture breccia fissure and wedge cleavage and fissility slickensides decollement shear zone boudin, raft, scale, floe cataclasite
buckle fold kink fold drag fold isoclinal, recumbent fold nappe translational flexure diapir, flow fold intrusive structure liquifaction structure hook fold mylonite, melange foliation (schistose) foliation (gneissic)
Dreimanis (1989) emphasized the polygenetic origin of till and noted a continuum of structures and deposits beneath a glacier, ranging from weakly deformed bedrock, to glaciotectonite, to local or deformation till, to classic lodgement till (fig. 1-11). Thus, glaciotectonic structures in bedrock may have gradational or abrupt transitions with glaciodynamic structures in overlying till, as noted by Houmark-Nielsen (1988).
lO00 km
Brittle
structures
Glaciotectonic structures
Glaciotectonic structures range in scale from microscopic to continental. Deformed material includes mostly sedimentary strata varying from unconsolidated to poorly and moderately consolidated in character. These are typically Cenozoic or Mesozoic in age, including gravel, sand, silt and clay, mudstone, sandstone, chalk, lignite and coal, etc. Less commonly, well-consolidated sedimentary strata and even
SCALE IO0 m
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ICE- PUSHED HILLS, MORAINES B RIDGES
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FOLDS 8t OVERTHRUST FOLDS . . . . . . . . . . . . . . . . CRUSTAL
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. . . . . . . . . . . . . . . . . . . . . . . .
Figure 1-12. Chart of common glaciotectonic structures arranged according to their typical horizontal scales (logarithmic). Ductile structures toward bottom; brittle structures toward top. Based on Occhietti (1973, Tab. II).
Nature of glaciotectonism crystalline rocks also have suffered glaciotectonic deformation (Kupsch 1955; Babcock, Fenton and Andriashek 1978; Lloyd Davies 2004). These types of rocks are usually older (Paleozoic, Proterozoic and Archean), such as strongly lithified limestone and dolostone, sandstone, gneiss and schist, granite and other plutonic rocks. It should be understood, however, that the conditions existing at the instant of deformation may have been quite different from what we observe today. For example, many deformations took place under conditions of high confining pressure or while the rock body was permafrozen. A complete listing of all glaciotectonic structures is probably not possible, simply because so much variety exists in type, style and size. Nonetheless, several common glaciotectonic structures can be recognized (fig. 1-12). In a general way, glaciotectonic structures can be divided in two broad categoriesmductile and brittlemdepending on the nature of deformation (Table 1-2). Ductile deformation takes place by internal creep or flow of material in a plastic or fluid manner. During ductile deformation, the rock mass has essentially no internal strength, so that quite small pressure differences may result in substantial changes in the size or shape of the mass. Ductile structures are most typical of unconsolidated or fine-grained strata, such as clay, silt, shale or chalk, that were deformed under high confining pressures. Various folds, intrusions, diapirs, and contortions represent ductile deformation. Among the most distinctive ductile glaciotectonic structures is chalk-till m61ange (Aber 1982) found in the western Baltic and eastern England (fig. 1-13).
ll permanent deformation. Brittle structures are most characteristic of consolidated or coarse-grained strata, such as sand, gravel, sandstone, limestone or quartzite, that were deformed under low confining pressure. Joints, faults, breccia, fissures and other fractured structures exhibit brittle deformation (fig. 1-14). Among the most unusual brittle glaciotectonic structures are caves beneath Montreal, Canada. Hundreds of meters of passageways were created by glacial shearing of well-consolidated Ordovician limestone layers along shale seams, which caused underground fissures to open in the limestone (Schroeder, Beaupr6 and Cloutier 1986). Ductile and brittle structures are often intimately associated within the same sequence of deformed strata. This is because strength varies for each stratum depending mainly on lithology and thickness of the layer, as well as ground-water pressure and the presence or absence of permafrost. Within a stratified sequence, layers of differing lithology and thickness each respond to deforming pressure differently. As a result, some layers develop brittle structures, and others show ductile deformation (fig. 1-15). Also, a particular rock that displays brittle deformation in some places may show ductile structures elsewhere due to variations in pressure, temperature or fluid content during deformation.
Brittle deformation results when rock masses fail by fracturing along discrete planes. Deformation is accomplished by movement or adjustment along fracture planes, whereas the internal fabric of the rock between fractures shows no
Some geologists have attempted to classify glaciotectonic structures according to size, style, degree of deformation, or mode of genesis. Brodzikowski and van Loon (1985) distinguished three groups: (1) fold-type structures, (2) faults and fissures, and (3) shear zones and related structures. Kozarski and Kasprzak (1994) also proposed three classesm glaciotectonite, glaciocataclasite, and glaciomylonite, representing increasingly strong degrees of deformation. Moran (.1971) identified three types of glaciotectonic structures.
Figure 1-13. Photograph of chalk-till m(lange from West Runton, England. Chalk and till were sheared and mixed together in ductile fashion by strong glacial deformation. Pocket knife for scale. Photo by J.S. Aber (1988).
Figure 1-14. Photograph of fractured flint nodules within chalk, Weybourne, England. Ice pressure transmitted through the chalk led to brittle fracturing of hard flint. Pocket knife for scale. Photo by J.S. Aber (1988).
Aber and Ber
12
Structures of the first and second types are distinguished mainly by size and amount of dislocation, rather than by any real differences in their modes of origin. The mechanism of transportational stacking has much in common with Hart's (1998) model of a deforming layer beneath a debris-rich layer in the basal ice (see fig. 1-10). As active shear planes shift up and down, slices of material are transfered from the basal ice into the deforming layer and visa versa. Like repeatedly shuffling a deck of cards, the migrating shear zone mixes up slices of glaciogenic sediment and substratum material. This structure can be demonstrated by vertical repetition, duplication, or discontinuities of textural or compositional attributes within the till. In other words, the till is a multilayer composite, which contrasts with the traditional idea of homogeneous till properties.
Figure 1-15. Upper portion of Hanklit, Mars, northwestern Denmark. Detached and folded mass of Fur Formation (A Eocene diatom#e) is enclosed by gravel (B) and rests on till (C). Note brittle fracturing and faulting in core of fold (left) and plastic stretching in nose of fold (right) Dimensions of view ~30 x 50 m. Photo by J.S. Aber (1979).
Many scientists have pointed out the striking similarity in deformational style between ice-shoved hills and true mountains (Hopkins 1923; Berthelsen 1979; van der Wateren 1985; Pedersen 2005). All manner of hard-rock structures described from mountains and shields has been recognized in glaciotectonic settings (Banham 1977). Glaciotectonic structures often mimic those seen in igneous and metamorphic rocks. In fact, ice-shoved hills may be regarded as natural scale-models of mountains (Croat 1987).
• Simple in-situ deformation - small folds and faults suffering only minor displacements involving bedrock and glaciogenic strata. The structures are relatively simple, in close proximity to the site of d6collement, and common near the bedrock-drift contact (fig. 1-16).
Understanding the architecture of mountains and ice-shoved hills involves detailed structural analysis, which has three stages: descriptive analysis, kinematic analysis and dynamic analysis (Davis and Reynolds 1996).
• Large-scale block inclusion - large, intact blocks of substratum (bedrock and glaciogenic strata) incorporated into the ice and transported away from the site of d6collement. Such blocks may be emplaced as individual, flat-lying rafts or as imbricated thrust masses (fig. 1-17).
• D e s c r i p t i v e analysis ~ geometry, orientation and structural patterns within deformed bodies.
• Transportational stacking - shear planes within a single till sheet, formed by repeated, episodic, differential movement of basal ice along nearly horizontal shear planes. Slices of till are rearranged, and the sequence of slices is thoroughly disordered within the till sheet.
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Figure 1-16. Simple in-situ glaciotectonic structures seen in a coal strip-mine, Illinois. Scale given in feet; note vertical exaggeration. Taken from Moran (1971, fig. 1).
Nature of glaciotectonism
13 EAST
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Figure 1-17. Idealized section through Thunder Hill, eastern Saskatchewan, based on test-drilling data, showing thrust stacking of Cretaceous bedrock blocks with glacial drift. Note large vertical exaggeration; adapted from Moran (1971, fig.
Methods for geometric description and analysis of deformed rock or sediment are presented in chapter 2. Structural data are collected in the field using a variety of direct measurements, subsurface prospecting, and remote-sensing techniques at various scales of observation. These data are then plotted in maps, cross sections, and stereonets to display and analyze 2- and 3-dimensional patterns of structural elements. The resulting descriptive analysis forms the basis for the next stage, kinematic analysis, which is the focus of chapter 3. The third stage, dynamic analysis, is discussed further in a later chapter (12).
Glaciotectonic landforms comprise a variety of types, including hills, ridges, buttes and plains, all of which are constructed wholly or partly of pre-existing bedrock masses or glaciogenic strata. Many kinds of glaciotectonic landforms have been described, and a great many terms have been used for ice-shoved hills, including (Kupsch 1962; Prest 1983): Stauchrficken, Stauchmor~nen and Stauchendmor~nen (German); stuwmorenen and stuwwallen (Dutch); moraine de chevauchement, crates de chevauchement and moraine de pouss6e (French); pseudo-moraine, push moraine, icethrust moraine, ice-pushed ridges, etc. (English).
Glaciotectonic landforms
The general term ice-shoved hill is used here to refer to any glaciotectonic landform of constructional nature. Depressions or basins formed by glacial erosion will be considered insofar as they may relate to ice-shoved hills. The term pushmoraine is restricted here to those ice-shoved hills that are composed largely or wholly of deformed glaciogenic strata. Moran (1971, p. 137) emphasized that glaciotectonism has
Glaciotectonic landforms are the surface or morphologic expressions of subsurface structures resulting from glacial deformation of bedrock and glaciogenic strata. The landforms may display their original or primary morphology as initially created during glaciation. This is particularly true of young (late Wisconsin/Weichselian) features that were little modified by subsequent glaciation or by postglacial erosion and deposition. In such cases, the morphology may be a direct expression of subsurface structures. In other situations, glaciotectonic landforms were significantly altered by later glacial or nonglacial events. The present landforms may be little more than erosional ruins of the original forms or may be largely hidden beneath younger deposits. In these areas, the morphology is only a subdued reflection of subsurface structures. In all cases, some knowledge of subsurface structure and stratigraphy is invaluable for properly interpreting the landforms.
probably played a major role in the formation of many if not most of the classical end moraines of midwestern United States. Thus, many ice-shoved hills have been interpreted mistakenly as end moraines. He noted that great care should be taken for interpreting the glacial sequence in such situations. The terms floe, raft, scale, and megablock have all been used for the individual dislocated masses of bedrock and drift that make up ice-shoved hills. Floe is a general term for any kind of dislocated and deformed mass, and scales are blocks thrust into an imbricated or overlapping position. The terms raft and megablock both refer to large, comparatively thin masses lying in more-or-less horizontal positions.
14
Aber and Ber
A
E B
B
Figure 1-18. Schematic block diagrams showing hilldepression landforms. A - single block of dislocated material B - multiple blocks in imbricated thrusts. Size and shape of such forms varies greatly. Adapted from Bluemle and Clayton (1984, fig. 4). Reproduced from Boreas by permssion of Taylor & Francis AS.
Clayton, Moran and Bluemle (1980) and Bluemle and Clayton (1984) recognized three types of ice-shoved hills in North Dakota: 1) hill-depression forms, 2) transverse-ridge forms, and 3) irregular forms. This scheme is significant, because it was the first attempt to classify glaciotectonic landforms for a large region, and because the hill-depression form was recognized as the fundamental type of glaciotectonic landform (fig. 1-18). A more elaborate classification is possible based on subdivision of these three types, inclusion of other landforms, and consideration of building materials. An expanded morphological classification for constructional glaciotectonic landforms includes four types (Aber 1988b): 1) hill-hole pair, 2) composite ridges, 3) cupola hill, and 4)
Table 1-3. Basic characteristics of constructional glaciotectonic landforms arranged in order of decreasing topographic prominence from the top down. See chapters 47 for further discussion. Landform
Height (m)
Area (km2)
Primary Material
.......
Hill-hole pair
Primary Morphology ...................
ridged hill with source depression
20 to 200
<1 to >100
variable
I00 to 200+
20 to > 100
bedrock
20 to 100
1 to 100
Quaternary strata
Cupola hills
20 to > 100
<1 to 100
variable, till cover
smoothed dome to elongated drumlin
Megablocks
zero to <30
<1 to 1000
bedrock
buttes, plateaus, or irregular hills
(large)
Composite ridges (small)
subparallel ridge-and-valley system, looped or arcuate in plan
Figure 1-19. Schematic block diagrams showing progressive change in morphology during glacier advance over a hilldepression landform. A -freshly created hill-hole pair, B minor rounding and molding, C- advanced streamlining into cupola-hill. Size and shape of such forms varies greatly. Adapted from Bluemle and Clayton (1984, fig. 5). Reproduced from Boreas by permssion of Taylor & Francis AS. fiat-lying megablock. Each class represents an ideal genetic type within a continuous spectrum of glaciotectonic landforms (Table 1-3). Intermediate, transitional, or mixed landforms exist between these ideal types (fig. 1-19). For example, Herschel Island (chap. 4) is the hill portion of a large hill-hole pair. Part of the island displays large composite ridges, whereas the rest of the island resembles a cupola hill. Such combination landforms are relatively frequent. Nonetheless, the four ideal types are distinctive enough to justify special recognition of each as an end-member class. The materials of which glaciotectonic landforms are built may be divided into three categories: 1) pre-Quaternary bedrock that is usually, but not always, consolidated to some extent; 2) pre-existing Quaternary strata, both glacial and nonglacial, that are usually unconsolidated; and 3) penecontemporaneous drift, deposited and subsequently deformed during the same glaciation, that is normally unconsolidated. Here, as with the landform types, these
Nature of glaciotectonism
15
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tt
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/
Stream!ined T e r r a i n
t
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Ice Advance
f
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-50 km l
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Figure 1-20. Sketch map of glacial landform features in east-central North Dakota. Ice-thrust terrain is concentrated along the limits of glacial advances, and streamlined terrain is located in upglacier positions. Adapted from Moran et al. (1980, fig.
5). categories represent end members with many landforms being composed of more than one type of material. Regardless of their present state of consolidation, any of these materials may have been permafrozen at the time of deformation. Glaciotectonic deformation did not take place randomly across glaciated terrains. Geomorphology of the pre-existing landscape had a strong influence on the types and locations of glaciotectonic landforms. Certain geomorphic settings are well known sites for glaciotectonism, as summarized by Banham (1975). • Valley side - flanks of valleys that contained ice lobes or tongues advancing along the valley axis (Krygowski 1965; Brykczyfiski 1982). • Scarp face - margins of escarpments or coteaus facing against the direction of glacier advance. • Island and peninsula - plateaus, buttes, or isolated high obstacles around and over which ice moved. • Down-gradient slope - tensile (rift) deformation on lowrelief plain sloping in downglacier direction. Regional mapping and classification of glaciotectonic landforms was undertaken by Moran et al. (1980) across the northern Plains of the United States and western Canada. They recognized two categories of glacier-bed landforms: 1)
glacial-thrust blocks and source depressions, and 2) streamlined terrain. The first includes hill-hole pairs and composite ridges typically associated with known ice-margin positions. The latter is found behind (upglacier) from ice margins and includes cupola hills and drumlins with cores of ice-thrust blocks. On this basis, the spatial distribution of glaciotectonic landforms is an important line of evidence for reconstructing dynamics of paleo-glaciers (fig. 1-20).
Casehistory approach The case-history methodology of teaching, which is well established in law, medicine, engineering, and other fields, is adopted for this book. In the following chapters, actual case examples are selected to illustrate the typical or salient characteristics of each morphological form or structural type. These examples are representative of glaciotectonic analysis and demonstrate both its possibilities and limitations. Some of these case examples are readily accessible, others are remotely located; some are thoroughly investigated, others have received only cursory study; a few are famous for their scenic beauty (fig. 1-21). In any event, they collectively display structures and geomorphic features that represent the full spectrum of glaciotectonic phenomena.
16
Aber and Ber
Figure 1-21. Aquinnah Cliff and lighthouse on the island of Martha's Vineyard, Massachusetts, northeastern United States. The multi-colored cliffs have attracted attention since prehistoric time. Each year thousands of tourists visit the cliffs, which are famous for their stunning appearance at sunset. See Aquinnah cupola-hill in chapter 6. Photo by J.S. Aber (2005).
Chapter 2 Geometric Analysis Introduction
Conventional field methods
The methods employed for geometric analysis of glaciotectonic structures and landforms are much the same as in other branches of earth science. These methods were developed in most cases for other purposes in exploration and mapping surface or subsurface geology and then were applied to glaciotectonic investigations. Glaciotectonic structures are really no different from folds, faults, and dislocations created by other means of tectonic deformation. In fact, glaciotectonic structures may be regarded as natural scale-models of true mountains (Berthelsen 1979). Thus, the techniques used by structural geologists can be applied directly to analysis of glaciotectonic features. These techniques fall in two general categories (Abet 1988a).
The search for glaciotectonic phenomena often begins with review of existing large-scale topographic, geologic, and soil maps, as well as airphotos of a study region. Such review may bring to light landforms or geologic conditions suggestive of glaciotectonic features. Typical glaciotectonic landforms are described in subsequent chapters and are a guide for what to seek in a preliminary map review. Such review may (or may not) reveal promising locales for further investigation.
• Descriptive analysis - - emphasizing three-dimensional geometry of structural elements. • Terrain a n a l y s i s - emphasizing geomorphic expression of structural features. Structural analysis of deformed bedrock and glaciogenic strata has three objectives (Billings 1972). First, the geometry of the structure must be determined--its size, shape, depth or thickness, and orientation. This is best done in the field using a variety of techniques ranging from simple surveying methods for surface exposures to well logging and geophysical measurements in the subsurface. The resulting geometric data are then displayed graphically in the form of maps, cross sections, block diagrams, or stereographic projections for further analysis. From this geometry, the second and third objectives of analysis--determining the age and genesis of the structureufollow. Terrain analysis for structural purposes is based on the assumption that landscape topography, drainage, vegetation, and soils often faithfully reflect the nature of subsurface structures. In fact, folds, faults, and other structures are often mapped on the basis of topographic expression, even where the bedrock is itself not visible at the surface. The so-called geobotanic approach relies on interpretation of vegetation and its relationship to geologic conditions (Punkari 1982, 1985). Vegetation integrates many geomorphic variables, such as sediment and soil, land slope and aspect, and drainage, which in turn reflect underlying substratum lithology and structure. Thus, vegetation patterns, both natural and artificial, often reflect geomorphic conditions. This approach is particularly useful for interpreting remotely sensed images, such as airphotos and satellite pictures.
In the field, available outcrops, excavations and exposures are sought out, including coastal and lake cliffs, stream banks, pits dug for sand and gravel, coal and clay mines, highway and railroad cuts, building sites, water wells, and so forth. These may be considered exposures of opportunity. Some larger, natural cliff sections may be available more or less permanently (see fig. 1-3); whereas, others may last only a few days or weeks during a construction project. It always pays to talk with local farmers, ranchers and other land owners, as well as water-well drillers, civil engineers and others engaged in construction projects. Most are quite cooperative and may grant access to study good sites on private land. In the authors' experience, one large, deep, well-exposed section is far more valuable than many small, shallow, poorly revealed exposures. The large, deep section allows much more thorough investigation of all lithologic units and examination of structures at all scales and levels within the geologic setting. Where such a large exposure of opportunity is not available, a study section may be excavated at a specific site chosen to maximize the potential for revealing subsurface strata and structures. In this case, the orientation, depth, and other characteristics of the excavation may be designed for best opportunity to study the buried structures and sediment. It is often possible to negotiate "add on" excavations in connection with established commercial enterprises. For example, a small payment may be sufficient to have extra trenches or sections dug at clay pits, coal mines, building sites, etc. (fig. 2-1). Likewise, water-well drillers often are amenable to drilling deeper in order to sample lower levels or reach bedrock beneath sediment cover. Basic geologic field equipment includes a survey compass (fig. 2-2), tape measure, survey markers, and assorted survey devices. Global positioning system (GPS) has become standard in recent years for quick positional data, and altimeters provide rapid elevation measurements. The level
18
Aber and Ber structures, and making geologic maps. Figure 2-3 shows a typical measured section for the coastal cliff profile at RCgle Klint, Denmark. Each body of sediment is identified and shown on the section in schematic style to represent its position in the actual exposure. Dashed lines suggest the likely subsurface extension of units as well as portions that have been removed by erosion prior to deposition of the discordant till. Structural features, thrust faults in this case, are marked.
Figure 2-1. Mine excavator has just finished removing the overburden from a section of glaciogenic strata (background) in a clay pit. The operator was hired to work "afierhours" to accomplish this preliminary opening of the exposure. Kadyny, Poland. Photo by J.S. Aber (1993).
RCgle Klint is an example of a key section in which normally buried strata were uplifted by glaciotectonic deformation and are exposed at the surface for first-hand study (HoumarkNielsen 2004). The section (fig. 2-3) demonstrates a basic structural principle, namely distinguishing unconformities within a glaciotectonic sequence (Houmark-Nielsen 1988, 1999). In this case, units 1-8 are all disturbed by folds and faults and truncated by an unconformity. Unit 9 rests on top of the unconformity and is itself undisturbed. The overall section, thus, consists of a dislocated portion below and a d i s c o r d a n t portion above that are separated by an unconformity. Glacier advance from the northeast was quite likely responsible for deformation of the dislocated strata, erosion of the unconformity, and deposition of the discordant till (SjCrring 1983a). Recognition of such relationships in the field is crucial for further structural analysis. Structural measurements are derived from two geometric elements of rock bodies--linear and planar features. A great variety of primary and secondary structures may be treated Table 2-1. Table of some common linear and planar elements associated with glaciotectonic structures. Many other linear and planar features may exist in some circumstances.
Figure 2-2. Silva compass of the type often used for geological field measurements. The needle is housed in a liquid-filled chamber to dampen its movements. The compass may be adjusted to compensate f o r local magnetic declination, and a small clinometer measures tilt angles. Seen here on a slickensided fault surface within brecciated clay, Kadyny, Poland. Photo by J.S. Aber (1993).
Linear
Planar
striation
bedding plane
trough cross bed
planar cross bed
groove & channel
joint, fracture, fissure
drumlin & flute trend
fault plane
augen & boudin axis
dike & vein
..................
of spatial accuracy necessary for most geologic field work is generally in the 1-10 meter range, which is provided by lowcost, relatively simple survey equipment and methods. For greater spatial accuracy (<1 m), more sophisticated techniques may be employed, generally with greater investment of time and cost of equipment.
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Among the fundamental field activities are measuring and sampling sections, collecting data on the orientation of
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Figure 2-3. RCgle Klint, island of Fyn, Denmark. The illustration presents a vertical section exposed in a coastal cliff. Note the several components: measured profile, scale bar, legend of symbols, and direction indication. 1 - Eocene clay, 2 Oligocene clay, 3 - Elsterian till, 4 - glaciofluvial sediment (Elsterian), 5 - marine interglacial clay (Holsteinian), 6- glaciofluvial sediment, 7 - till (Saalian or Weichselian), 8 - Weichselian till, 9 - discordant till (Weischelian). Position of thrust faults indicated by X-X. Ice movement from left. See Fig. 2-6for location of ROgle Klint; adapted from SjCrring (1983a, fig. 177).
as lines or planes for purposes of collecting orientation data (Table 2-1). The orientation of a plane relative to the Earth's surface is determined by two measurements~strike and dip. Strike is defined as the compass direction of a horizontal line in the plane. Compass directions are customarily measured in azimuth degrees (0 ° to 360°). Dip is the direction and angle of maximum downward tilting, which is always perpendicular (90 °) to strike. The orientation of a line relative to the Earth's surface is likewise determined by two measurements~trend and plunge. Trend is defined as the compass direction of the line, and plunge is the direction and angle of downward tilting of the line (fig. 2-4). Detailed investigations at many sites are the basis, then, for constructing regional maps. Such maps vary considerably in kinds of information and styles of presentation. Conventional maps of surficial geology usually portray glaciogenic sediments first by textural composition--till, sand, gravel, etc. Morphology may be considered also, such as end moraine, sandur, drumlin, esker, etc. Glaciotectonic
Figure 2-4. Pen rests on excavated axis of small fold in upper portion of photograph. The fold is situated just above a thrust fault in glaciogenic strata. Island of Herdla, near Bergen in western Norway. Photo by J.S. Aber (1987).
phenomena may not be apparent, unless they are included specifically in the map classification (fig. 2-5). For purposes of glaciotectonics, features of importance are ice-shoved hills, source basins, and related glacial landforms such as meltwater spillways, tunnel valleys, drumlins, eskers, etc. In addition, directional indicators are often included as clues to glacier movements (fig. 2-6). Among the most important of these directional indicators are folds, faults, and other structures associated with icepush deformation. As a rule of thumb in structural geology, larger structures are more significant for interpretation of kinematics than are smaller features. For example a thrust and folded mass of bedrock several 100 meters in dimensions
Figure 2-5. Portion of Geologic highway map of North Dakota. Qct- ice-thrust hills, Qco - outwash sediment, Qcl - lacustrine sediment, Q o d - wind-blown sediment, Qcg glacial sediment (till - ground moraine), and Qce - glacial sediment (till - end moraine). Lineations in southwestern portion are long drumlins. Taken from Bluemle (1988).
Aber and Ber
20
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. . M a i n StationaryLine -,--,,- Presumedposition of recessional ice margin Figure 2-6. Map of glacial features associated with the main phase of Weichselian glacier advance in central Denmark. Large arrows indicate directions of glacial movement based on glaciotectonic structures. Note the general distribution of tunnel valleys, eskers, drumlins and other glacial features in relation to glaciotectonic deformation and the overall pattern of glacier advance. Asterisk shows location of RCgle Klint (see fig. 2-3); Modified from SjCrring (1983a, fig. 194). reflects greater and longer-acting glacier influence than does a single striation on a boulder pavement. Thus, glaciotectonic structures of large size should be given greater weight for field investigation and interpretation of glacier dynamics than are striations, till fabrics and other indicators of relatively small size. Stereographic projections
A stereographic projection, or more simply a stereonet, is a powerful method for displaying and manipulating the 3dimensional geometry of lines and planes (Davis and Reynolds 1996). The orientations of lines and planes can be plotted relative to the center of a sphere, called the projection sphere, as shown at the top of Fig. 2-7. The intersection made by the line or plane with the sphere's circumference defines the 3-dimensional orientation of the line or plane. The projection sphere is divided in half by a horizontal plane called the projection plane. Compass (azimuth) directions are indicated on the projection plane. All measurements are made relative to horizontal, so only the lower hemisphere of the projection sphere need be used.
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The lower stereonet hemisphere can be projected in various manners onto a flat graph. However, stereographic projections distort the distances and spatial relationships of features. Stereonets are preferable to maps or cross sections for solving many geometric problems involving lines and planes that are common in structural geology. Two popular stereographic projections are the equal-angular or Wulff stereonet and the equal-area or Schmidt stereonet. The Wulff stereonet is perhaps the easier of the two projections to visualize. Imagine that your eye is positioned at the zenith of the projection sphere (fig. 2-7). The dipping plane intersects the lower hemisphere forming an arc that your eye visualizes on the horizontal projection plane. The resulting projection-plane arc is, in effect, a perspective view of the dipping plane as seen from the top of the sphere. The Wulff stereonet consists of a series of circular arcs that form meridians running N-S and parallels running E-W. Although the meridians and parallels resemble longitude and latitude lines on a map, they are not; the stereonet is not a map. Angles are preserved on the Wulff stereonet; note that meridians and parallels always intersect at fight angles. However, size and shape are distorted.
Geometric analysis
21 10 "" I0 20 ~ ' ~ ~ I ~ ~ , ~
PROJECTION SPHERE
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30
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features may be made, and all the data can be plotted on a single Schmidt stereonet to display the overall structural pattern. Because it maintains size, the Schmidt stereonet is visually more pleasing than the equal-angular stereonet, and the Schmidt stereonet has proven more popular for routine use.
PROJECTION OF DIPPING PLANE Figure 2-7. Dipping plane plotted relative to the center of a projection sphere (above), and horizontal projection of that plane (below). The plane has a strike/dip of 500/45 °. Illustration adapted from Aber (1988a, fig. 6-1). The Schmidt stereonet is a different projection on which size is preserved, but angles and shape are distorted (fig. 2-8). Each quadrangle on the stereonet is equal in size or area, whereas shape of quadrangles varies from almost square near the center to narrow, curved rectangles near the edge. Unlike the Wulff stereonet, on which meridians and parallels are circular arcs, the meridians and parallels of the Schmidt stereonet are oval curves. The Schmidt stereonet has one important a d v a n t a g e - - i t preserves area. Therefore, the density of plotted data can be analyzed. In many situations, dozens or even hundreds of measurements of linear or planar
Plotting of lines and planes on either stereonet projection is carried out following exactly the same procedures. Typically linear elements, such as fold axes, are plotted as points on the stereonet, and planar elements, such as faults and joints, are plotted as arcs. In the case where many planes must be plotted on a single stereonet, visual confusion could result. To simplify plotting planes on the stereonet, it is convenient to plot the pole to each plane. The pole is a line running through the center of the projection sphere and perpendicular to the plane (fig. 2-9). The pole forms a 90 ° angle with the strike line and a 90 ° angle with the dip line. Thus, the pole will always be found in the opposite quadrant of the stereonet from the dip of the plane. A c o m m o n structural p r o b l e m in g l a c i o t e c t o n i c s is determining the axis of a fold. The axis is a linear geometric element that represents the line of maximum curvature in a folded surface. It is often possible to measure strike and dip of many planar surfaces (strata) on the fold limbs, but the fold axis itself may not be exposed or is inaccessible for direct measurement. In this case, measurements on the fold limbs may be adequate to determine the so-called constructed fold axis. Each planar limb measurement is plotted on the stereonet as a pole. The poles tend to align along an arc (fig.
Aber and Ber
22
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Figure 2-10. Example of a Pi diagram showing constructed fold axis based on several measurements of bedding planes on a series of cylindrical folds. Strike and dip of bedding planes are plotted as poles, which define an arc on the stereonet. The pole to the arc represents the constructed fold axis. Illustration adapted from Aber (1988a, fig. 8-2). S
Figure 2-9. Pole perpendicular to dipping plane in projection sphere (above), and the resulting stereographic projection of the plane and its pole (below). Illustration adapted from Aber (1988a, fig. 8-1).
4-km-long sequence of exposures in western Michigan (fig. 2-11). They concluded that deformation had come from the west by an ice lobe that expanded from the Lake Michigan basin. Subsurface methods
2-10). The pole to this arc represents the constructed fold axis. Stereonets are a powerful way to present and analyze complex structural data. Nowadays computer software is employed normally for plotting stereonets, calculating geometric relationships, and determining statistical distribution of features. Stereonets may be added to stratigraphic profiles, cross sections, or maps to provide a 3-dimensional basis for interpretation of the structural patterns. Rieck et al. (1991) used this approach effectively in their analysis of structural features in glaciotectonically disturbed peat and gyttja for a
Exploring the subsurface is a pursuit that has occupied geologists for many practical and theoretical applications, and thus a great deal of effort has been expended on developing methods for sampling and probing the Earth's interior. These methods fall in two general categories--direct sampling and geophysical prospecting. Direct sampling is based in most cases on drilling and physical recovery of cores or cuttings. Various geophysical techniques may be undertaken either from the surface or within drill holes. For purposes of glaciotectonics, the shallow subsurface (<300 m deep) is most usually of interest.
Geometric analysis
23
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pole to bedding constructed fold axis
8&9
Figure 2-11. Map of glaciotectonic sites in peat and gyttja exposed along the Pine River in western Michigan. Stereonets display poles to bedding, constructed fold axes, measured fold axes, and fold asymmetry (Z symbols). Adapted from Rieck et al. (1991, fig. 2). Reproduced from Boreas by permssion of Taylor & Francis AS. Test drilling for core or cutting recovery is quite common for geological exploration and mapping programs, as conducted by state, provincial or national geological surveys. On-site logging by an experienced geologist normally provides good records for such wells, and cores or cuttings may be archived for later laboratory study. In most countries, commerical water-well drillers also are required to file well logs, although the quality and usefulness of such data are quite variable. In the authors' experience, such information should be viewed with caution.
structural interpretations. The most telling indication of the potential for glaciotectonic dislocation is repetition or omission of strata (Christensen 197 la), particularly the occurrence of glacial strata beneath pre-Quaternary bedrock (fig. 2-12). Jakobson (1996) listed two situations in which glaciotectonic structures may be suspected on the basis of well data (fig. 2-13).
Down-hole geophysical methods are often used in combination with direct sampling of test wells. Electrical and radioactive logs provide an important supplement to the actual samples and may provide a more-continuous down-hole record than do recovered samples (Ehlers and Iwanoff 1983). The test drilling program of the Saskatchewan Research Council is a good example of this combined approach (Christensen 197 la). A rotary drilling rig with 300 m depth capability was utilized. Cuttings were collected from drilling fluid, and a special sidewall sampler allowed additional collection at specific depths. Electrical logging included spontaneous potential (SP) and electrical resistance (R).
• Repetition of a distinctive, well-defined sedimentary unit, such as a marine interglacial clay.
Well data are essentially linear, one-dimensional profiles through three-dimensional geologic bodies. As such, considerable care should be exercised for stratigraphic or
• Presence of pre-Quaternary strata above glacial sediments, such as Paleogene bedrock over till.
In order for either situation to be applied, the local stratigraphy must be well understood and the units in question should be relatively continuous within the stratigraphic sequence throughout the region. The finding of Quaternary glaciogenic strata beneath pre-Quaternary bedrock is strong evidence for glaciotectonic deformation. However, repetition or omission of strata may be explained in some cases by other mechanisms, so caution is advised. In general, a single well record may be suggestive of possible glaciotectonic dislocations, but is not conclusive. Additional surface or subsurface evidence should be sought before reaching a glaciotectonic interpretation.
Aber and Ber
24 spontaneous potential (SP)
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SAND TILL SHALE Figure 2-12. Representative well log from near Steen, Saskatchewan, showing lithologic column, spontaneous potential, and electrical resistance. Scale bar = 30 m (100 feet). A layer of till is present underneath a block of shale 33 m thick. This is evidence for glaciotectonic dislocation of the shale block. Adapted from Moran (1971, fig. 4). Among the newer methods for shallow subsurface prospecting is ground penetrating radar (GPR), a geophysical method that has much in similar with seismic-reflection techniques. In the case of GPR, short pulses of high-energy microwaves are transmitted from the surface into the ground. The microwaves propagate through the ground, interact with subsurface materials in complex ways, and may be scattered or reflected from interfaces with dielectric contrasts. Microwave frequency (wavelength) determines the depth of penetration as well as spatial resolution of the resulting data. Longer wavelengths (lower frequency, 25 to 200 MHz) result in deeper penetration, but lower resolution. Such GPR is useful for determining deeper and larger geologic structures. Shorter wavelengths (higher frequency, 300 to 1000 MHz) are better for identifying shallow, small objects such as buffed pipelines or individual boulders. Microwave interactions are influenced strongly by water content of soil and sediment; the best subsurface conditions are either dry or frozen materials. LCnne and Lauritsen (1996) utilized a portable field GPR unit for investigating a push-moraine complex in front of the small valley glacier Scott Turnerbreen near Longyearbyen
Till Tit Eemian marine clay
Figure 2-13. Representative lithologic well logs from Denmark. Left - presence of Paleogene bedrock above glaciogenic strata, western Limfjord district. Right repetition of marine clay unit, island of ~Ere. Taken from Jakobsen (1996, fig. 5). in Svalbard. The push moraine was formed during one or more glacier surges, most recently in 1930. They operated the GPS unit in two modes: 50 MHz, 1000 V, 1600 pulses per second (ps), and 200 MHz, 400 V, 800 ps. The transmitter and receiver were spaced 1 m apart in both cases, and the unit was positioned on snow or ice cover 1-2 m deep. Two people were sufficient to move the radar unit along a path across the push moraine complex in a longitudinal traverse (fig. 2-14). On the resulting profiles, depth is given in twoway travel time (nanoseconds), which was converted to linear depth according to the method of Annan and Davis (1976). The most distinctive features revealed on the profiles are tops of buffed ice blocks, sediment bedding planes, and numerous, closely spaced, imbricated thrust faults that dip in the upice direction (fig. 2-15). They concluded that proglacial thrusting had taken place in permafrozen substratum. In the Netherlands, Bakker and van der Meer (2003) utilized GPR to reveal the internal structure of the Veluwe pushmoraine complex. A dense grid of ~250 km of GPR transects was collected at 50 MHz, which achieved a maximum penetration depth of 45 m. This depth was possible because
Geometric analysis
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the sediments are relatively coarse (mostly sand and gravel) and the water table is relatively deep (20-50 m). GPR data were supplemented with test drilling, cone-penetration, and shallow seismics (Bakker 2004). These ancillary observations demonstrated that coherent radar reflectors consist of finegrained layers, namely silty fine sand (Bakker 2002). Based on these examples, GPR seems to hold great promise for elucidating shallow glaciotectonic structures in many situations.
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Another relatively new geophysical technique is the transient eleetrornagnetie (TEM) method, developed for shallow subsurface investigations (Auken, JCrgensen and SCrensen 2003; JCrgensen, Sandersen and Auken 2003). TEM is based on the response of subsurface materials--solids and fluids-to an electromagnetic impulse. In the conventional mode, TEM is done with a 40 m x 40 m transmitter loop carrying a steady current of 3 A and a receiving coil in the center of the loop (fig. 2-16A). The ground response is measured when the transmitted signal is turned off for intervals ranging from 9 gs to 9 ms. In this mode, ground penetration is about 100 m to 130 m depth. In the high-moment mode, a 75 A current is passed through a 30 m x 30 m transmitter loop, which can
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Figure 2-15 (below). Ground penetrating radar profiles (above) and interpreted sections (below). A - 50 MHz frequency with ground penetration ~30 m. B - 200 MHz frequency with ground penetration ~10 meters. Upice direction (northeast) to left; adapted from LOnne and Lauritsen (1996, fig. 6).
Figure 2-14. Sketch map of Scott Turnerbreen showing position of ground penetrating radar (GPR) profile in the push-moraine complex in front of the glacier. Adapted from LCnne and Lauritsen (1996, fig. 2B). 0
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26
Aber and Ber
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in areal extent and depth, and have lower dip angles compared to results from conventional, land-based geological observations. This reflects the gridded areal coverage and depth of penetration of seismic data as well as the inability to resolve steeply dipping features in seismic sections (Andersen 2004). Shallow seismic exploration may yield equally valuable results on land, as demonstrated by Krzywiec et al. (2004) in the vicinity of Ortowo, Warmia, northern Poland. This region, long-known for complicated glaciotectonic deformation and uplift of Tertiary strata, came under coordinated shallow geophysical investigations during the early years of the 21st century (Morawski 2004a). Results from other methods of geophysical surveying, particularly geoelectric profiles (Twarogowski 2004), were integrated with seismic and gravity data plus borehole information to provide precise 3dimensional spatial interpretation of glaciotectonic and stratigraphic conditions. Despite difficult topographic conditions, seismic profiles clearly imaged complex glaciotectonic structures, their substratum and overburden (fig. 2-19). Dislocations reach depths of 300 m and seem to involve both Miocene and Oligocene formations, which were thrust generally towards the northeast. This direction is almost opposite to the main direction of advancing glaciers (Morawski 2004b); however, a detailed explanation of this apparent contradiction should be regarded at the moment as an open question requiring further studies.
Remote sensing Remotely sensed images of the Earth's surface span visible, infrared, and microwave portions of the spectrum and range in scale from 1"100 to 1"1,000,000. It is well beyond the
27
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scope of this text to describe the plethora of remote sensing techniques that might be applied to glaciotectonic situations. For further details about remote sensing, see Jensen (2000) and Lillesand, Kiefer and Chipman (2004). Following are brief descriptions of some common methods.
Aerial photography has been a mainstay in geological exploration and mapping programs for more than half a century. In single frames or stereopairs, conventional airphotos are essential for all types of modern geological and geomorphological investigations, including study of glacial
Aber and Ber
28
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Figure 2-19. Seismic profile from the Ortowo vicinity, southern Warmia, Poland. Q - Quaternary, N - Neogene (Miocene), E Paleogene (Oligocene, Eocene, Paleocene), K - Cretaceous. Note slices of Oligocene strata (E~) thrust and folded within Miocene (N) toward right side of profile. Adapted from Morawski (2004b, fig. 10). -
Figure 2-20. Conventional panchromatic airphoto of the Crestwynd vicinity, southern Saskatchewan, Canada. Large composite ridges cross the scene from NW to SE. Photograph A21639-7 (1970); original photo scale = 1:80,000. Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada.
Geometric analysis
29
Figure 2-21. Landsat multispectral scanner (MSS) scenes from southern Saskatchewan. A - spring image, 21 May 1978, MSS band 7 (near-infrared). The Dirt Hills and Cactus Hills are large ice-shoved ridges on the northeastern margin of the Missouri Coteau upland. B - winter image, 9 Dec. 1972, MSS band 7 (near-infrared). Note the "pen-and-ink" quality of this snowcover view in which topography is quite evident. This is among the earliest Landsat images. Scanned from original image prints at scale 1:1,000,000; acquired from U.S. Geological Survey, EROS Data Center.
30
Aber and Ber
Figure 2-22. Manned space-shuttle photograph of southern Denmark and northernmost Germany. Low-oblique view toward north; compare with Fig. 2-6. Late Weischelain ice lobes followed Baltic Sea channels between the islands and mainland. Derived from color-visible photograph, STS68-153-038, Sept. 1994. Image obtained from NASA Johnson Space Flight Center, courtesy of K. Lulla.
landforms (Way and Everett 1997). This mature method has expanded in recent decades upward to space and downward to low-level platforms in order to provide diverse kinds of imagery. Conventional airphotos are medium-scale, panchromatic, large-format (23 cm), vertical views (fig. 2-20). Glaciotectonic landforms have been pictured in airphotos for geologic atlases, for example Gravenor, Green and Godfrey (1960) and Prest (1983). Such airphotos are indispensible for regional mapping and assessment for all manner of environmental conditionsmsoils, geology, water resources, vegetation, etc. Aerial photographs of this type are available in principle for most northern countries in which glaciotectonic landforms are conspicuous; however, the age, quality and cost of such airphotos vary greatly. At a nominal scale of 1:40,000, for example, a vertical airphoto covers approximately 85 km 2 ground area, and 0.1 mm on the airphoto represents 4 m on the ground. The usual resolution limit for conventional airphotos at scales 1:20,000 to 1:25,000 is 1-2 m; smaller objects cannot be discerned
unless they have high contrast with the surroundings. The panchromatic (black-and-white) nature of conventional airphotos also limits their use for interpreting and classifying vegetation cover, which is often an indicator for underlying geologic conditions. Color-visible or color-infrared aerial photographs are available for selected areas in only a few countries. A recent trend is conversion of hardcopy airphotos into digital orthophotos, which have been scanned and rectified to fit a cartographic grid system. Orthophotos may be imported and used directly in GIS databases (see below).
Space-based remote sensing for civilian, scientific applications began with the manned Skylab Missions and the unmanned Landsat satellite series in the early 1970s (fig. 2-21). Landsat instruments provide a wealth of moderateresolution datasets for all manner of earth-science applications. Many Landsat datasets are in the public domain and are available from the U.S. Geological Survey EROS Data Center or from other remote sensing centers at modest cost. The potential of Landsat imagery for glacial geomorphology was recognized early by Morrison (1976) and Slaney (1981). In regions of low relief, winter Landsat images
Geometric analysis
31
impart a strong morphologic aspect to the scene due to low sun angle, snow cover, and lack of active vegetation (Skoye and Eyton 1992). Manned space photographs likewise have given regional overviews for visualizing glacial landscapes (fig. 2-22). Other satellite systems of particular interest are Radarsat, operated by the Canadian Centre for Remote Sensing, and Ikonos, a commerical, high-resolution multispectral scanner. The spatial resolution of the latter rivals traditional airphotos. Many other systems and types of space-based remotely sensed imagery are widely available nowadays from various governmental and commercial sources. Remote sensing from space provides several advantages compared to conventional aerial photography (Short and Blair 1986). • Synoptic, big-picture views that portray large areas of the Earth's surface. • Multispectral data--visible, infrared and microwave portions of the spectrum. • Repetitive, global coverage throughout the year and spanning several decades. • Low cost compared to collecting same type and volume of data on the ground. The application of space-based remote sensing has revolutionized geomorphology, and indeed has led to a new subdiscipline known as megageomorphology, which is concerned with geomorphic description and analysis of large areas (Baker 1986). This approach has extended to glacial geomorphology (Williams 1986), in which remote sensing has provided the data for large-region to continent-scale mapping and interpreting of glaciated landscapes. Punkari (1980) was among the earliest to use this approach for interpreting glacial dynamics in Finland, which he subsequently elaborated to include all of Scandinavia and northwestern Russia (Punkari 1996). Similar investigations were carried out in North America (e.g. Boulton and Clark 1990a, 1990b; Clark 1993). At the other end of the height and scale spectrum, smallformat aerial photography (SFAP) has emerged as a means to acquire large-scale, high-resolution imagery for detailed site investigations (Warner, Graham and Reed 1996; Bauer et al. 1997). This method is based on 35- or 70-mm film cameras as well as compact digital or video cameras operated from low-height, manned or unmanned platforms. Many types of lifting platforms may be employed, such as manned aircraft (Light 2001), and unmanned model airplanes (Quilter and Anderson 2000), blimps (Ries and Marzolff 2003) or kites (Aber et al. 1999). Vertical photographs typically have ground resolution of 2-10 cm (linear pixel size), which is suitable for mapping and analysis in the microstructural scale range, 1:100 to 1:1000 (Masing 1998).
Figure 2-23. Kite aerial photograph of Feggeklit, a small ice-shoved hill on the island of Mors, northwestern Denmark. The cliff section on the eastern (right) side exposes deformed Eocene Fur Formation uplifted from the floor of the Limfjord estuary in the background. For more information see chapter 3. Photo by S.W. Aber (2005). Relatively low cost combined with high portability and ease of use in the field make SFAP an excellent method to acquire large-scale airphotos for site-specific geomorphic investigations (Marzolff and Ries 1997). Glacial landforms, such as eskers and drumlins, are portrayed especially well in SFAP (Aber and Gat~tzka 2000; Aber and Kalm 2001), and the method is well suited for depicting glaciotectonic landforms (fig. 2-23). SFAP has proved effective for accurately mapping complex glaciotectonic structures exposed in high cliff sections. The Danish approach involves low-height, oblique airphotos taken in overlapping sequence along the cliff face. The photographs are acquired from a small, manned airplane that flies parallel to the cliff at a distance of 200-300 m. Ground control points are derived from standard vertical airphotos of the study site. Multi-model photogrammetric analysis is applied to oblique stereopairs based on the method developed by Dueholm (1992). This technique was utilized for construction of detailed geological sections for two of Denmark's most famous glaciotectonic sitesnMCns Klint (Pedersen 2000) and LCnstrup Klint (see fig. 1-4). In the latter case, Pedersen (2005) produced a section 6 km long at scale 1:500 with a spatial resolution of ~25 cm. Detailed geometric data of this type are invaluable for elaborating the glaciotectonic genesis of such key sites.
Aber and Ber
32 An important approach for remote sensing is the multiconcept, including multilevel, multispectral, and multitemporal datasets (Avery and Berlin 1992). Each type of remote sensing has particular strengths for application to glacial geomorphology, and used in combination these systems provide data spanning all spatial, spectral and temporal dimensions of observation.
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Geographic information systems The multiconcept may be expanded to include many other types of relevant data, such as field observations, subsurface logs, topographic maps, soil maps, digital elevation models, etc. Multiple kinds of evidence can be integrated in geographic information systems (GIS), which allow display, analysis, and modeling of complex landscapes based on many "layers" of input data. GIS has emerged rapidly during the past two decades to become a powerful tool of choice throughout academic, governmental and commercial enterprises dealing with earth science data and applications (Burrough 1986; Faust, Anderson and Star 1991; DeMers 2003; Clarke 2004). The potential for utilizing digital elevation models (DEM) was demonstrated by Lidmar-Bergstr6m, Elvhage and Ringberg (1991) for display and interpretation of glacial landforms in Skhne, southern Sweden. At the time, their DEM was a classified military database, but now moderate- to highresolution DEM datasets are available to the public for many portions of the world. Aber (1999) employed a shaded-relief map derived from a DEM as a background for portraying preIllinoian glacial landforms in the central United States. Colgan, Mickelson and Cutler (2003) expanded the GIS approach for the southern Laurentide ice sheet from the Dakotas to New England. Clark (1998) realized the potent capability of GIS for reconstructing and modeling paleo-ice sheet dynamics over continent-sized areas. GIS was adopted for compilation of glaciotectonic maps in North America (Aber et al. 1991, 1995) and in central Europe (Ber and Abet 2003). As an example, Jakobsen (1996) employed a GIS approach for mapping the density of concealed (subsurface) glaciotectonic structures in Denmark based on well records (fig. 2-24). The distribution of concealed glaciotectonic structures reveals some distinctive patterns. In many cases, areas with high densities of concealed structures coincide with well-known glaciotectonic landforms (various ice-shoved hills). But in some other cases, zones with intense
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Figure 2-24. Map of Denmark showing density of concealed (buried) glaciotectonic structures in 5-km-square cells. Values for each cell indicate number of wells with glaciotectonic features on a per mille (0.1%) basis. Zones with high densities of glaciotectonic features: 1 - northwestern Jylland, 2 southeastern Jylland, 3-southern Fyn, ~Ere andAls, 4-Great Belt, 5- central Sjcelland, and 6-Men. Adapted from Jakobsen (1996, fig. 8).
concealed glaciotectonics have few or no associated landforms. He determined that regional distribution of glaciotectonic phenomena in Denmark corresponds at a first order to lithology of the pre-Quaternary substratum (Jakobsen 2003). On this basis, GIS has become fully integrated into the working method for glaciotectonics at local, regional, and continental scales. GIS facilitates merging of geographic datasets collected in various ways and over a range of scales and resolutions, which allows sophisticated geometrical analysis and numerical modeling of glacial dynamics and landscape systems.
Chapter 3 Kinematic Analysis
Stress and strain The goals of kinematic analysis are to measure the strains that have taken place within the rock or sediment body and to relate these strains to the stress field that induced the strain. Stress is simply force per unit area, and can be expressed in such common units as pounds per square inch (psi), kg/cm 2, or atmospheres (1 atm. ~ 1 kg/cm 2) pressure. A stress unit often used in structural geology is the bar (b), which is 1 standard atmosphere pressure (1 kilobar = 1000 bars). The high pressure imposed on deeply buried rocks due to weight of the overburden is uniform and equal in all directions. Such stress is called lithostatic pressure, and it is determined solely by the thickness and density of the overburden. A similar, uniform pressure is imposed by glaciers and ice sheets on the substratum and is glaciostatic pressure (see chap. 1). It is compressive or positive stress, in contrast to tension which is a negative stress. Lithostatic pressure may p r o d u c e strain in rocks due to simple compaction, but it probably cannot account for most folding and faulting. For this, differential stress is required. Differential stress may be developed by unequal normal stresses or by shear stress (figs. 3-1 and 3-2). Some common variables in kinematic analysis are listed below. cr normal stress ( n ) ~ compressive (or tensile) stress applied normal (perpendicular) to a surface. "c shear stress m stress applied parallel to a surface within a geologic body. 0 angle of failure between the fracture plane and the maximum (c~) stress vector. "c° cohesive strength ~ measure of rock or sediment resistance to fracturing. 0 angle of friction - - measure of resistance to internal deformation (slip). e s t r a i n - percentage shortening (or lengthening) of body subjected to stress.
Strain in rocks is accomplished by dilation, which is a change only in size, or by distortion, which is a pure shape change, or commonly by both. The amount of strain is, of course, related to the magnitude of stress, and the strain response of various rocks to stress can be tested in laboratory apparatus. Kinematic analysis, in fact, relies heavily on laboratory testing of sample rock and sediment under various physical conditions in order to determine how geologic materials fail
under stress. Thus a particular type or pattern of failure (fracture or fold) may be related to particular conditions of stress application. Laboratory testing of rock and sediment samples reveals certain basic characteristics of material strength. First, geologic materials are relatively strong when subjected to simple lithostatic compression. They are much weaker when differential stress is applied, because this creates shear stress within the rock body. Rocks and sediments are weaker still when exposed to tension. Rocks and sediments do not fail structurally because of compression; they fail when subjected to shear stress or tension. Second, rocks and sediments fail by fracturing at certain characteristic angles relative to the stress field (fig. 3-3). • Shear fracture ~ plane of failure oriented about 30 ° from 0~, 60 ° from o 3, and parallel to o 2. Most common type; two crossing sets may form a conjugate pattern. • Extension fracture - - plane of failure parallel to ~yi and ~Y2, and perpendicular to o 3. Not common geologically; develop only where there is little or no confining pressure. • Release fracture - - plane of failure perpendicular to o~, and parallel to o 2 and c~3. Develop when stress load is released, as when buried rocks are unloaded by removal of an ice sheet. Among these types, conjugate shear fractures are most important, because they form a recognizable "X" pattern that can be used to establish orientation of the o stress vectors within the rock or sediment body. The typical angle (0) for shear fractures, 30 ° to 35 ° , holds for all manner of hard and soft rocks under a wide range of physical conditions (Kulhawy 1975). Furthermore, a geometric relationship exists between 0 and ¢, such that ¢ = 90 - 20 or 0 = (90 - ¢)/2. Thus, most rocks have an angle of friction (¢) also equal to about 30 °. These factors may be combined into the Mohr stress equations to yield stress conditions operating within rock or sediment bodies along planes at different angles (0). In most cases, it is sufficient to analyze o 1and o 3 values, as they represent the maximum in differential stress. n = (c~l + (y3)/2 - ((Yl - o3)/2 " (cos 20) "c = (cy~ - cr3)/2 • (sin 20)
34
Aber and Ber MINIMUM
STRESS
ELLIPSOID
F i g u r e 3-1. The stress ellipsoid. Three n o r m a l (perdendicular) stress axes define the total stress applied to a rock body. cr1 is the maximum compressive stress, cr2 is intermediate, and cr3 is the minimum. Normally all three stresses are positive (compression); cr~ may be negative (tension) in some cases. Where only lithostatic pressure is d e v e l o p e d , (~1 -~ (~2 = (~3, and the stress ellipsoid becomes a sphere. Adapted from Aber (1988a, fig. 9-1).
Figure 3-3. Fractures produced experimentally in a block of Solenhofen Limestone, where cr1 > cr2 > cr3. Fracture sets A and B are conjugate shear fractures; C are extension fractures; and D are release fractures. Adapted from Aber (1988a, fig. 9-4); based on Hobbs, Means and Vi611iams (1976, fig. 7.31).
MOHR
2~
.-.....,lw
t-
STRESS
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1
u
Kb i
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SHEAR STRESS .J,t .
.
.
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.
.
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.
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05
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Figure 3-2. Hard augen (gray) within a ductile matrix subjected to shear stress parallel to foliation. The sense of shear on oblique microfractures o f augen is opposite to overall sense of shear within the rock mass. Adapted from Aber (1988a, fig. 9-2); based on Simpson and Schmid (1983,
Figure 3-4. Basic Mohr diagram showing normal stress (horizontal axis) and shear stress (vertical axis) developed on a plane at 0 = 30 ° (2 0 = 60 °) with cr1 = 7 kb and cr3 = 2 kb. Solving the stress equations for this situation gives n = 3.25 kb and "c = 2.17 kb. Adapted from Aber (1988a, fig. 9-
fig. 9).
6).
These equations are often presented as a Mohr diagram or stress circle, which is a graphical means to visualize .and solve the stress values (fig. 3-4). A Mohr diagram represents the nature of rock failure under a given set of physical conditions. A semicircle with its center on the normal stress axis at (c~ + (y3)/2 connects cyl and o 3, and the radius of the semicircle equals (°l - (]3)/2" The angle 20 is measured from 0"3 in clockwise direction. The point on the semicircle for a given 20 angle indicates a normal stress value and a shear stress value for a plane oriented at angle 0 within a rock. Note: the Mohr circle has no physical reality; it is simply a graphical solution to the stress equations.
over a range of conditions. As confining pressure increases, the ~1 stress and 0 angle at which failure occurs also increase. The points of rock failure on each circle define a line that is more-or-less straight at lower confining pressures and flattens out toward higher confining pressures (fig. 3-5). This line is the failure envelope; its intercept with the shear-stress axis gives cohesive strength (Xo), and its slope is the angle of internal friction (¢). Any point on or above the failure envelope represents stress conditions that will cause the rock to fracture. Construction of a Mohr failure envelope represents a powerful means for evaluating rock strength.
A given rock or sediment type can be subjected to testing under different physical conditions, for example increasing confining pressure (o 3) in order to evaluate rock strength
The position of folds likewise can indicate the stress regime at the time of deformation. The fold axis is assumed to be perpendicular to o 1 and o: stresses and parallel to cy3 stress.
Kinematic analysis
500
35
1000
~500
2000
2500
Figure 3-5. Mohrfailure envelope for Carrara Marble. T° represents cohesive strength, and 0 is the angle of internal friction. Adapted from Aber (1988a, fig. 9-7). These conditions lead to some basic rules of thumb for interpreting the direction of ice movement from glaciotectonic structures. • Thrust f a u l t - strike of thrust is normal to direction of ice movement. In most cases, dip direction is upglacier, opposite to ice advance. In a few cases, back (or antithetic) thrusts may dip in the downglacier direction. •
•
Normal f a u l t - strike of fault plane is normal to direction of ice movement. Dip may be in either direction, up- or downglacier, but downglacier dip is more common. Fold axis of fold is approximately horizontal and normal to direction of ice movement. Asymmetrical folds have vergence (sense of overturning) in the downglacier direction.
Croot was among the first to apply this method in glaciotectonics (Croot 1987, 1988b). He analyzed structural displacement of small composite ridges in front of Eyjabakkaj6kull, an outlet glacier of Vatnaj6kull Ice Cap in southeastern Iceland (fig. 3-6). Several sets of small composite ridges were formed as the glacier surged to its maximum Little Ice Age extent in 1890. These composite ridges contain dislocated sandur sediments comprised of interbedded sand, silt, clay, tephra, and basal gravel approximately 2 m in total thickness. The tephra beds provide stratigraphic markers for correlating disturbed and undisturbed sections. The dimensions, and thus volume, of each ridge set vary, but an upvalley/upglacier depression is associated with each. Each ridge set is visually proportional in area to the basin that lies upglacier from it. However, on closer inspection it becomes clear that only a portion of the material comprising the composite ridges was transported from the basins. A simplified section, parallel to ice movement through one of the ridge sets, reveals that only part of the material was transported any distance from its source (fig. 3-7). The remainder was thrust forward proglacially, and the transport distances for these thrust sheets are < 5 m each. The material removed from the upglacier basin is more deformed and represents about 50% of the cross-sectional area (shaded portion, fig. 3-7). Assuming this is a roughly consistent proportion, then 50% of the ridge volume was derived from the basin, amounting to 50,000 m 3. The basin immediately upglacier from the ridge set has a surface area of approximately 25,000 m 2. This corresponds to removal of 2 m of material from the area of the lake basin. This value is
As with all such rules of thumb, there are exceptions. A single structure may not be conclusive, but several kinds of structures taken together usually yield a reliable interpretation for direction of ice advance.
Balanced cross sections A useful graphical technique to evaluate strain is the balanced cross section. This method was introduced by Dahlstr6m (1969) and applied to the Canadian Rocky Mountains. The basic idea is to reconstruct undisturbed strata as they existed prior to dislocation by folding or faulting. A fixed marker, called a pin point, is the basis for placing strata into original, horizontal position. A comparison of the reconstructed section to the deformed section, then, demonstrates the strain as a change in length, by either shortening or extension of the section. The method depends on the assumption of conservation of mass (volume), in other words a given mass of rock or sediment displaced from a source retrains its volume within the displaced sequence. In a balanced cross section, volume is represented in two dimensions as area on the section.
i
Bc® s d
LANGJOKULL ........
.....
. ..-.. ,..: :::::i:ii .i::'f.~,aB:A~,A~OKUL[
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, ........................
,
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Figure 3-6. Sketch map showing locations of larger ice caps and selected outlet glaciers with composite ridges. EyjabakkajOkull is a small outlet glacier on the northeastern margin of VatnajOkull Ice Cap, southeastern Iceland. Adapted from Aber, Croot and Fenton (1989, fig. 4-11).
Aber and Ber
36 <................
0
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5
F O R M E R ................ ..
E X T EN S I0 N
ICE M A R G I N
~,~,
.......
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:.COMPRESSION
Figure 3-7. Section through a composite-ridge set in front of EyjabakkajOkull showing that portion (shaded) of the ridge transported more than 10 m from its original (predeformation) site. Adapted from Aber, Croot and Fenton (1989, fig. 2-14). Plasma :: :: : : ::
: : :::
:: : ;
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:[
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Figure 3-8. Sequential development of a composite-ridge set at the snout of Eyjabakkaj6kull, Iceland, as interpreted from post-deformation data. Adapted from Croot (1988b, fig. 11). consistent not only with depth of the lake (source) basin, but also with the thickness of individual thrust slices in the composite ridge. Croot reconstructed a series of balanced sections to illustrate the progressive displacement and deformation of the small
::i::~i~!~!::~;~;~+/:7,~.o ~ ~
.... ~,~ ~ ?
Figure 3-9. Taxonomy of microfabrics and microstructures in deformed glaciogenic sediments. Plasma refers to the matrix (particles <30 ~tm in size I, and skeletal grains are larger particles. Adapted from M~nzies (2000, fig. 2); used with permission, Geological Society. composite ridge during the 1890 advance of Eyjabakkaj6kull (fig. 3-8). The source basin is marked by low-angle normal faults that dip in the downvalley direction. The arrangement of these faults implies convergence of faults along a d6collement at a depth of 5-6 m beneath the former ice margin. Total amount of extension in this portion is estimated at 385%. The small composite ridges are characterized by
Kinematic analysis
37
fe:ss~rs sh~:dows
Figure 3-11. Rotation of a hard clast (gray) within a ductile plasma and resulting pressure shadows. The pressure shadows appear like tails pointing in the direction of overall shearing. Based on Hart and Boulton (1991, fig. 5); reprinted with permission from Elsevier. Faster > > > > >
.
....
@~.~
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.
.
.
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Figure 3-12. Model for rotational strain within a bed undergoing differential movement (faster at top, slower at bottom). The rotations may involve individual grains or particle aggregates. Adapted from van der Meer (1997, fig. 2); reprinted with permission from Elsevier.
Micromorphology
Figure 3-10. Photomicrographs of highly strained diamicton in cross-polarized light. Above - sigmoidal grain (s) suffered internal strain. Arrows indicate sense of shear. Field of view ~2½ mm across. Middle - galaxy structure. Note spiral fabric (concentric layers) of grains around the core clast (c). Field of view ~4 mm across. Below- reworked diamicton pebbles (d) embedded within plasma. Field of view ~4 mm across. Images courtsey of W.R. Jacobson, Jr.
Microscopic structures within deformed glaciogenic materials have proven valuable for discerning the manner of strain imposed by glacial stress, and a large body of research has been conducted on this subject (e.g. Menzies and Woodward 1993; van der Meer 1993, 1997; Menzies 2000). The method involves collecting oriented samples, which then are impregnated with epoxy and cut into thin sections. The thin sections are examined with a petrographic microscope to describe and measure the physical and structural properties of the deformed material (Khatwa and Tulaczyk 2001). On this basis, a large number of characteristic brittle and ductile microstructures are recognized (fig. 3-9).
stacked thrust faults dipping in the upvalley direction. Total shortening is 54%. This example demonstrates the power of balanced sections to evaluate kinematics in glaciotectonic settings.
These studies have demonstrated that much strain takes place by rotational movements within the deformed material. Rotations may be internal to a single grain, take place between adjacent particles, or involve particle aggregates. Typical
Aber and Ber
38 microstructures include sigmoidal grains, galaxy fabric, and reworked clasts (fig. 3-10). Shearing may produce pressure shadows (fig. 3-11). Differential movement within a deforming layer may be likened to a series of turning wheels or ball beatings (fig. 3-12). The styles and orientations of such microstructures may be utilized to determine how the material has strained, and this could lead to reconstruction of the stress regime.
Superimposed deformation One of the great challenges in structural geology is working out multiple phases of deformation, each of which may have a different character and orientation. In shields and mountains each phase of deformation imparts a new set of structures that overprint or even obilterate older structures. Glaciotectonic structures display similar possibilities. Pedersen (2000) elaborated four settings for superimposed glaciotectonic structures (fig. 3-13). • Glaciotectonic structures superimposed on preglacial tectonic structures. • Glaciotectonic structures representing two distinct regional ice advances.
• Glaciotectonic structures for different phases of deformation by the same ice advance, • Glaciotectonic structures superimposed by postglacial neotectonic structures. In addition to these, glaciotectonic and neotectonic structures could develop simultaneously in some situations. In all these settings, multiple phases of deformation lead to complicated structural patterns that can be quite difficult to tease apart from field observations based on limited exposures, test drilling, GPR, or other techniques. In fact, excellent exposures are often necessary to unravel such details, for example in long cliff sections or large quarry faces. Among the most remarkable structures are refolded folds, that is folds affected by two (or more) phases of deformation from different directions. A dramatic example of this superimposed structure is the "arrowhead" fold pattern (fig. 3-14). Crossing fold sets create interference patterns within a rock or sediment body, just like crossing wave sets on a water body. Erosion of the folds may reveal such interference as arrowhead or crescent patterns (fig. 3-15). Refolded structures may, in rare cases, be evident in the landscape, where fold crests coincide with hill crests, as en 6chelon patterns (Pedersen 2000).
C
A
~:te~sio~I f~ult
Axial p~ane ~\ G2
DI
\
B
\
thrust ~ult
\
D2
\
D \~ ..............
:
!:i]
Figure 3-13. Four settings for superimposed glaciotectonic structures. A- glaciotectonic structures superimposed on preglacial tectonic structures, B - glaciotectonic structures representing two distinct ice advances, C- glaciotectonic structures representing different phases of deformation within the same ice advance, D - glaciotectonic structures superimposed by postglacial neotectonic structures. Adapted from Pedersen (2000, fig. 1).
Kinematic analysis
39
8 2
Figure 3-14. Remarkable exposure of an "arrowhead" fold pattern in Eocene strata, island of Mors, Denmark. Oblique photograph showing a horizontally cleaned surface in a claypit. See next figure for geometric explanation. Image courtesy of S.A.S. Pedersen.
Z
I
p-
0
~-t
it i
LITHOL~Y
STR~RES
~
~
AXPLi
F1
Figure 3-15. Model for refolded structure. Fold phase 1 consists of tight, overturned folds with gently dipping axial planes, which are warped by fold phase 2, an upright, open anticline. Erosion across the top of the resulting folds reveals the interference pattern as arrowhead and crescent outcrop patterns. Compare with photograph above. Taken from Pedersen (2000, fig. 7). Pedersen (1993, 1996) demonstrated the challenge and reward of deciphering superimposed deformation at Feggeklit, island of Mors, Denmark. A sequence of deformed Paleogene and glaciogenic strata are well-exposed in a coastal cliff section, nearly 1 km long (see fig. 2-23). Dislocated bedrock consists of the Fur Formation comprised of distinctively interlayered clayey diatomite, tephra beds, and cemented zones. Numerous tephra beds provide stratigraphic markers throughout the
c,aYey,,,
¢',:t .,.=p.t:e, ~¢¢~
Figure 3-16. Stratigraphy of Fur Formation and glaciogenic strata exposed at Feggeklit, northwestern Denmark. 1 southern portion of cliff section, 2 - middle section, 3 northern section. Individual tephra beds are numbered for stratigraphic correlation. Adapted from Pedersen (1993, fig. 3).
formation. Olst and Holmehus Formations underlie the section. The glaciogenic sequence includes three tills" T1 Saalian, dislocated; T2 - probably Saalian, dislocated; T3 Weichselian, discordant (fig. 3-16). The cliff section trends NNE-SSW parallel to the direction of glacier movement, and thus normal to strike/trend of most structures. Several phases of deformation took place as Weichselian ice advanced from the NNE and eventually overrode the vicinity. Initial phases of deformation generated various fractures, folds and faults, and took place in a proglacial environment. The final phase was subglacial and produced breccia (glaciotectonite). On this basis, Pedersen
Aber and Ber
40
ds s ~ a r •
•
i
::stress
Figure 3-17. Photograph of deformed Fur Formation in southern portion of Feggeklit section, zone 1. Dark tephra layers are interbedded with light-colored diatomite. Note the low amplitude, long wavelength folds and small faults. Height of cliff exposure ~10 m. Photograph by J.S. Aber (in Pedersen 1993, plate 13).
ANASTOMOSING J~NT!I~
C~ATE
(1996) recognized three structural zones, from least disturbed in the distal (southern) part of the section (zone 1), to moderate deformation in the central portion (zone 2), to greatest dislocation at the proximal (northern) end of the section (zone 3). Each section is represented by characteristic structures.
FAULTING: •._................._.l~ ~L
R A ~ THRUST~G %
• Zone 1 - - long wavelength, low amplitude folds, anastomosing joints, and conjugate faults (fig. 3-17). • Zone 2 - thrust fault ramp, conjugate faults and box folds, listric splay thrust faults. • Zone 3 - - low-angle thrusts and splay thrust faults, overturned anticlines.
Figure 3-18. Fracture and fault structures at Feggeklit (below) showing typical dip angles (0). Mohr circles (above) showing relationship between angle of failure (20) and increasing ice pressure (Cru). Adapted from Pedersen (1993, fig. 7).
The dip angles (0) of fractures and faults may be related to increasing ice pressure during advance toward and over the vicinity (fig. 3-18). Balanced cross sections for Feggeklit are the basis for constructing a basal d6collement, which is projected to lie in bentonite of the upper Holmehus Formation, ~100 m deep (fig. 3-19). The balanced sections indicate approximately 100 m of horizontal shortening, which represents about 10% overall strain for the Feggeklit structural complex. Based on varves and other chronologic control, Pedersen (1996) estimated that deformation took place within a 50-year time span and that regional ice advance averaged 10 m per year, a reasonable value for steady-state (non-surging) ice movement.
stratigraphy in a routine manner. Individual till sheets cannot be traced far. Morphologic features may be relicts of older glaciations. A glacial advance that created much deformation may leave few deposits of its own. In many ways, the problems of stratigraphy in glaciotectonic terrain are analogous to stratigraphy of complexly deformed crystalline shields. The stratigraphy of shields is often considered as a sequence of deformational events. Each event is marked by a distinctive style, type and orientation of structures and is characterized by certain metamorphic or igneous rocks.
Kineto-stratigraphy In many regions with abundant glaciotectonic phenomena, it may be difficult or even impossible to conduct glacial
Berthelsen (1973, 1978) developed a similar stratigraphic approach, called k i n e t o - s t r a t i g r a p h y , for unraveling complex Quaternary sequences in glaciotectonic terrain. He defined a kineto-stratigraphic drift unit as, the sedimentary
unit deposited by an ice sheet or stream possessing a characteristic pattern and direction of movement (Berthelsen 1973, p. 23).
Kinematic analysis
41
~ W
NNE
~n ~int
Q
pin pG nt
:,./. ............ .......... : ~ ....... :
:~.
: .....
. . . . . . . . . . .
'
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.......
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./~
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.............
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Figure 3-19. Schematic balanced cross sections for Feggeklit, northwestern Denmark. The pin point is placed at the southern end of the exposed cliff sequence. 1 - simplified stratigraphy and structure of the exposed section. 2 - balanced cross section with thrust faults projected to a basal ddcollement. Also shown are antithetic (backthrust) faults. 3 -final deformed section. The final section is approximately 100 m shorter than the balanced section. Adapted from Pedersen (1996, fig. 17).
A kineto-stratigraphic drift unit includes all primary sediments (till + stratified drift) that were deposited during all stages of a particular glacier advance and retreat (fig. 320). The primary sediments may be deformed by the same ice advance subsequent to deposition. Such deformations are termed domainal; deformation of subjacent and older strata is extra-domainal. The lower boundary of domainal sediments is the limit of penetrative deformation, which also marks the bottom of the kineto-stratigraphic drift unit (compare with fig. 1-9). The unifying factor in kineto-stratigraphy is the directional character of the deposits that can be related to ice movements
during a particular glaciation. The directional character is revealed by many kinds of features: striations and grooves, till fabrics, indicator erratics, and most importantly glaciotectonic structures. Domainal deformations are most useful in this regard, as there is no doubt about the glacier advance to which they correspond, but extra-domainal structures may be used also for kineto-stratigraphy. The kineto-stratigraphic principle of Berthelsen (1978, p. 29) is that: Deposits laid down by successive glaciations can be distinguished by means of the kinetic patterns deduced from the domainal and extra-domainal deformation
\ S
A, 3
Az '
Af
-
~:Tz
Figure 3-20. Schematic cross section of deposits classified in a kineto-stratigraphic drift unit. Black = till; stipled = stratified drift. A -- advance phase; S = maximum (sandur) phase; R = recessional phases; K = kame. Ice advance from right to left. Taken from Berthelsen (1978, fig. 1).
Aber and Ber
42
related to each glaciation, and that the glacial deposits should be correlated according to their directional elements.
opposite flanks may differby as much as 120 °. Near the lobe axis, ice flow direction will not change much during the course of lobe advance and retreat. However, flow directions will shift significantly on lobe flanks as the lobe expands and shrinks. These local patterns of ice movement and glaciotectonic deformation are consistent in light of the lobate model for glaciation (see chap. 11).
In order for kineto-stratigraphy to be applied successfully, three prerequisites must be met. First, consecutive ice advances must come from different directions. Multiple advances from the same direction would all be classified in a single kineto-stratigraphic drift unit. Second, each advance must display a consistent and recognizable regional pattern of movement. And third, it must be possible to relate various directional features to the proper ice advance. This is a matter for careful field observations and structural analysis.
Since its inception, kineto-stratigraphy has been applied for working out glacial advances in diverse situations. Hicock and Dreimanis (1985) demonstrated the advantages of this approach in British Columbia and Ontario, Canada. In fact, they found glaciotectonic structures are more reliable than till fabrics for establishing the direction of ice movement. Glaciotectonic structures may be measured quickly in the field and serve as a check on other directional indicators. Elsewhere in North America, middle Pleistocene ice-lobe advances were deciphered with kineto-stratigraphy in the central Great Plains of the United States (Aber 1991).
The second p r e r e q u i s i t e ~ c o n s i s t e n t pattern of ice movement~may be the most difficult to deal with, because kineto-stratigraphy is an empirical method. It is tested regionally, but developed by fitting together local results. The local results may at first seem inconsistent or even haphazard when considered in isolation. The direction of ice movement may vary markedly across a region and may even vary significantly at individual sites for different phases of the same glaciation.
Recognition of glaciotectonic structures proved invaluable for unravelling glacial stratigraphy and glacier dynamics in the Kap Ekholm sections, Spitsbergen, Svalbard, where a sequence of marine and glacial strata are exposed in cliff sections up to 30 m high and more than one km in total length. This key section had been studied before in considerable detail, but the presence of glaciotectonic structures was not recognized by earlier investigators (Boulton
Consider an idealized ice lobe advancing from the north. A northerly flow direction is developed only along the central axis of the lobe. In the lobe flanks, ice flow diverges toward the southeast and southwest. Thus, flow directions on
IE
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F
15
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.
"-~o-.~o._-~o_.......:.._..~
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I
0
Figure 3-21. Stratigraphy and structures exposed in a portion of the Kap Ekholm sections, Spitsbergen, Svalbard. Low-angle thrusts and small folds are located within Formation F, especially in the 20-22 m level. The sandy diamicton (shaded) at the top of the formation is interpreted as a deformation till. Formation F consists of mid-Weichselian marine strata that were deformed by late Weichselian glacier overriding. Adapted from Mangerud and Svendsen (1992, fig. 15); reprinted with permission from Elsevier.
43
Kinematic analysis
1979), which led to some erroneous stratigraphic interpretations. Mangerud and Svendsen (1992) identified low-angle thrusts and normal faults as well as various folds and vertically tilted strata. Many of these structures had escaped previous recognition, because they strike/trend N-S, parallel to the main direction of cliff exposures (fig. 3-21). The orientation of these structures demonstrated late Weichselian ice movement from east to west, crosswise to the adjacent fjord, which suggests a relatively thick ice cap that moved independently of local topography. Kineto-stratigraphy has proven most successful in the country of its origin--Denmark (Berthelsen 1978; Houmark-Nielsen 1981, 1987). Five phases of Weichselian glaciation are recognized (fig. 3-22). Old Baltic ice lobes moved into the southern islands from the southeast. Two old Baltic advances are now recognized and considered to be mid-Weichselian in age (Houmark-Nielsen 1994, 1999). Late Weichselian glaciation took place in several stages beginning with the Norwegian advance coming from the north and covering northern Sj~elland, Sams¢, and Jylland (Pedersen 2005).
with at least one or two readvances and a shift to more easterly movement during the course of ice coverage. Following a brief interstade, the Young Baltic ice lobe overspread the islands from the southeast and reached eastern Jylland. The final Ba31thav readvance took place in the form of ice tongues moving from the south along channels between larger islands and the mainland. All these late Weichselian advances happened during a relatively short time interval between 20,000 and 14,000 years BE The kineto-stratigraphy of Denmark is based on the traditional model for Weichselian glaciation (Holmstr6m 1904; Andersen 1966), in which ice lobes were fed by ice streams following topographic depressions and emanating from the interior of the ice sheet. This model is best illustrated by the Young Baltic phase (fig. 3-23). The ice lobe advanced from east to west along the southwestern Baltic, and divergent ice flow turned toward the northwest and north in southern Denmark. Local variations in ice m o v e m e n t and glaciotectonic deformation could be explained by development of small ice tongues on the margin of the main lobe.
The Main Weichselian advance next came from the northeast and reached into central Jylland. This advance was complex
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Figure 3-22. Limits or areas of Weichselian glacial advances in Denmark with some key glaciotectonic sites indicated (dots). 1 - Old Baltic advances, 2 - Norwegian advance, 3 Main northeast advance, 4 - Young Baltic (East Jylland) advance, and 5 - Bcelthav advance. Adapted from Berthelsen (1978, fig. 11).
Figure 3-23. Reconstruction of Weichselian Young Baltic glaciation in Denmark and adjacent areas. E = East Jylland advance; B = Bcelthav readvance. Taken from HoumarkNielsen (1987, fig. 138).
Aber and Ber
44 This ice-lobe model for Weichselian glaciation was challenged by Lagerlund (1980, 1987), who postulated the existence of marginal ice domes. These domes presumably grew during episodes of colder climate on older stagnant ice or inactive surge lobes. The domes developed their own dynamic flow in a radial pattern that was independent of the main Weichselian ice sheet (fig. 3-24). The question of ice lobes versus marginal ice domes is controversial and is still the subject of continuing debate after more than 25 years of research (e.g. Ringberg 1988; Houmark-Nielsen 1999). Kineto-stratigraphy, through its emphasis on ice movement patterns, has proven to be a valuable approach in this situation.
Glaciodynamic sequence and event Pedersen (1993, 1996) integrated the concept of kinetostratigraphy within a comprehensive framework for all aspects of glacial deformation, erosion and deposition. He defined several terms as follows (Pedersen 1993, p. 68-69). • Glaciodynamic event m one regional glacial advance characterized by a unidirectional dynamic impact. • Glaciodynamic sequence ~ the result of a glaciodynamic event. The glaciodynamic sequence may include both Quaternary and pre-Quaternary sediments affected by one glaciotectonic deformation. • Glaciotectonic deformation ~ deformation caused by a prograding glacier which forms at least two phases, 1)
foreland folding and thrusting and 2) subglacial brecciation and cataclasis. • Glaciotectonic phases m deformation phases related to a glaciotectonic deformation event. The glaciotectonic phases are characterized by specific structures and their relationship marks the succession of phases. • Glaciotectonic u n i t - a rock unit deformed by one glaciotectonic deformation. The rocks are mostly sediments, but bedrock is not excluded. • Kineto-stratigraphic u n i t - a till or composite till unit including a glaciotectonite and the glaciofluvial sediments related to the advancing and retreating ice margin. The relationship of the glaciodynamic event and the glaciodynamic sequence is demonstrated in Figure 3-25. The glaciodynamic sequence includes all sediments deposited during an ice advance as well as pre-existing strata impacted by glaciotectonic deformation. In general, the glaciotectonic unit is found in the lower portion of the sequence, and the kineto-stratigraphic unit comprises the upper portion. The boundary between the two is determined by the lower limit of penetrative deformation, marked by brecciation or cataclasis, at the base of the kineto-stratigraphic unit. This boundary represents a glaciotectonic unconformity, which may serve in itself as an important stratigraphic marker (Houmark-Nielsen 1988, 1999).
YOUNGER
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! Figure 3-24. Reconstruction of a marginal ice dome during the Young Baltic glaciation of Denmark and adjacent areas. Gray shading indicates center of the elongate dome across southernmost Denmark, northernmost Germany and northwestern Poland. Adpated from Lagerlund et al. (1983, fig. 5).
OLDER
DOWN
Figure 3-25. Schematic illustration showing the relationship of glaciodynamic sequence, glaciodynamic event, kinetostratigraphic unit, and other impacts resulting from the regional advance of an ice lobe or cap. Glaciotectonic structures dominate the lower portion of the sequence, and the kineto-stratigraphic unit, including at least one till, forms the upper portion of the sequence. Adapted from Pedersen (1993, fig. 1).
Chapter 4 Hill-hole Pair
Introduction The hill-hole pair is perhaps the simplest and most instructive type of glaciotectonic landform. It represents a basic combination of ice-scooped basin and ice-shoved hill. Other types of glaciotectonic landforms are variations of the hillhole theme. The association of individual ice-shoved hills with discrete source depressions was first described by Jessen (1931) from northern Jylland, Denmark. The pairing of anomalous hills and upglacier depressions in central North Dakota was noted by Bluemle (1970) and Clayton and Moran (1974), who correctly recognized the glaciotectonic origin of the hill-hole pairs. In the past, these hills were often misidentified as kames or in-place outliers of bedrock, depending on their internal composition. Hillhole pairs are now widely recognized. Bluemle and Clayton (1984, p. 284) described the hill-hole pair as, a discrete hill of ice-thrust material, often slightly
crumpled, situated a short distance downglacier from a depression of similar size and shape. The hill and associated depression are usually next to each other, but may be separated in some instances by as much as 5 km. Both pre-existing drift or bedrock may be involved in the dislocated hills. The depression represents the source of material now in the hill. Depressions today are often the sites of bogs, lakes, estuaries, bays, or simply low spots in the surrounding topography. The volume of the depression ideally should be equal to or slightly less than the volume of the hill. Where this can be shown, a direct genetic link is demonstrated between the hill and hole. However, as depressions are often partly filled with younger sediment, it may be difficult to determine a depression's original volume. In some cases, the source depression for a hill cannot be identified. The depression may actually exist but is hidden by younger sediment cover or is under a large lake or sea. It can be demonstrated in a few situations that the dislocated mass was a pre-existing hill that was simply moved to a new location (Moran et al. 1980). Hills without associated depressions are known in northeastern Poland and Belarus (Levkov 1980). On the other hand, anomalous depressions without associated hills are also known in several regions, for example in central Poland (Ruszczyfiska-Szenajch 1976, 1978) and eastern Alberta (Aber, Croot and Fenton 1989).
The basic morphology of typical hill-hole pairs is shown clearly by Antelope Hills, central North Dakota (fig. 4-1) and at Hundborg, western Denmark (fig. 4-2). Antelope Hills is composed mainly of deformed, soft Cretaceous or Paleocene sedimentary bedrock upthrust in multiple ridges some 90100 m above the surrounding landscape (Carlson and Freers 1975; Clayton, Moran and Bluemle 1980). A depression to the west, partly occupied by a small lake, is the source for Antelope Hills, and spillway channels trend around the northern and southern ends of the ice-shoved hill. The hill at Hundborg is situated along the Eocene outcrop belt of northern Jylland, wherein spectacular ice-pushed structures were created in the soft clayey bedrock (Gry 1940). Sjcrring SO is the source depression for the hill at Hundborg. Characteristic morphologic features of simple hill-hole pairs include: • Arcuate or crescentic outline of hill; concave on the upglacier (proximal) side, convex in the downglacier (distal) direction. • Multiple, subparallel, narrow ridges separated by equally narrow valleys following the overall arcuate trend of the hill. • Asymmetrical cross profile of hill; higher with steeper slopes on convex (distal) side, lower with gentler slopes on concave (proximal) side. • Topographic depression on concave (upglacier) side of hill; area and shape of depression approximately equal to that of hill. • Meltwater spillway channel or esker often extends from the source basin around or across the hill (Bluemle and Clayton 1984). Another small hill-hole pair example is Devils Lake Mountain in northeastern North Dakota. The ice-shoved hill is a slightly arcuate, continuous, single ridge that parallels a narrow source basin on its northwestern side (fig. 4-3). The iceshoved ridge is approximately 4 km long, 1 km wide, and stands up to 55 m above the source basin (fig. 4-4). The southwestern ends of the source basin and the ridge are connected by a linear escarpment that bounds the hill-hole pair (fig. 4-5). This lineament is interpreted as a tear fault, which confirms the geometric link between the source basin and hill. Where all these morphologic traits are present, a hill-hole pair may be identified with confidence, even with little knowledge of subsurface conditions. However, original hill-
46
Aber and Ber Figure 4-1. Topographic map of Antelope Hills vicinity, central North Dakota. The source basin is centered in the west half of section 19. Locations of road cuts revealing deformed bedrock shown by solid dots. Elevations in feet; lO-foot contour interval (~3 m). Adapted from U.S. Geological Survey, 7.5-minute quadrangle, Antelope Lakes, North Dakota.
. . . . . . . . . . . .':::
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Figure 4-2. Topographic map of Hundborg vicinity, western Denmark. SjOrring SO represents the source depression for material shoved into the arcuate ridge immediately to the south. Elevations in m; contour interval = 5 m. Adapted from map sheet 1116 Thisted, Geodcetisk Institut (Denmark).
hole morphology was commonly modified by later glacial and nonglacial processes, so the ideal set of characteristics is not always present. Source basins are often reduced in apparent size by sediment infilling. Ice-shoved hills, likewise, are commonly reduced in size by later erosion, or they may be partly concealed by deposition of glacial sediment. The morphologic trends of one hill-hole pair may be truncated by those of a younger hill-hole pair formed by ice advance from a different direction (Pedersen 2005).
Hill-hole pairs
47
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Figure 4-3. Low-height blimp aerial photograph of Devils Lake Mountain seen from the northwestern side. Superwideangle image; nearly all of the ice-shoved hill and source basin are visible. Takenfrom Aber (2003-04, fig. 3).
Figure 4-4. Perspective block diagram of Devils Lake Mountain as seen from the southwest. The source basin lies to left side of the ice-shoved ridge. Digital orthophoto image draped over digital elevation model. Large vertical exaggeration; image courtesy of by W.R. Jacobson, Jr. .),..... ..... ~
.............................................................
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Figure 4-5. Topographic map of Devils Lake Mountain vicinity, northeastern North Dakota. The source depression for Devils Lake Mountain is the narrow lake~marsh basin that parallels the northwestern side of the ice-pushed ridge. Note the linear escarpment connecting the southwestern ends of the source depression and ridge (A-A'). Elevations in feet; 5-foot contour interval (~1.5 m). Adapted from U.S. Geological Survey, 7.5-minute quadrangle, Devils Lake Mountain, North Dakota.
Aber and Ber
48 The sizes of hills and related depressions usually vary from 1 km 2 to 100 km 2, and often many hill-hole pairs of different sizes and shapes are found in close proximity to each other. Topographic and structural relief of hill-hole pairs generally range from 30 m to 200 m, although exceptions to these size and relief ranges are known.
surface is completely mantled by drift. Wolf Lake and Wolf Hill in combination, thus, represent a large hill-hole pair in a complex setting (Fenton and Andriashek 1983). The distinct morphologic expression alone makes this an outstanding example, even though almost nothing is known concerning its internal structures or materials.
The following case examples demonstrate hill-hole pairs in various settings. Wolf Lake, Alberta and Herschel Island, Yukon are both large Pleistocene hill-hole pairs from western Canada. Herschel Island is constructed of permafrozen, preglacial Quaternary strata, whereas the internal nature of the hill at Wolf Lake is unknown. Devils Lake comprises a complex of hill-hole pairs associated with a major aquifer in North Dakota, and submarine hill-hole pairs are described from the Norwegian continental shelf.
Wolf Lake and Wolf Hill are both shaped approximately as similar, aligned parallelograms, 7.5 km long and 3 km wide. The alignment of the lake and hill is emphasized by a conspicuous, southwest-trending lineament formed by the straight eastern edge of Wolf Lake and the straight eastern flank of Wolf Hill. The Wolf Lake lineament is more than 11 km long and displays 135 m of total topographic relief. Any lineament of similar prominence in a nonglacial setting would be recognized as a major vertical fracture, such as a strike-slip fault. This is the interpretation for the eastern as well as western boundaries of Wolf Lake (fig. 4-7). The lateral boundaries of the ice-scooped depression are essentially tear faults along which material from Wolf Lake basin was shoved into Wolf Hill.
Wolf Lake, Alberta Wolf Lake is situated in east-central Alberta near the Saskatchewan border (fig. 4-6). A large hill of ice-shoved material, here called Wolf Hill, is located immediately south of the lake. This hill reaches a maximum elevation of 759 m (2476 feet), some 158 m above the normal level of Wolf Lake. To the west of Wolf Lake, smaller, streamlined, ice-shoved hills, drumlins, and flutes are well developed. The land
Wolf Hill consists of a single, steep-sided, east-west trending, asymmetric mound with a rounded crest, the southern flank being steeper than the northern one and (fig. 4-6). A belt of
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Figure 4-6. Topographic map of Wolf Lake vicinity, Alberta. Position of ice-shoved ridges based on analysis of aerial photographs. Elevations in feet; contour interval = 50 feet (~15 m); perennial lakes shown by diagonal lining; ephemeral lakes shown by dashed outlines; F = fire watch tower Adapted from Aber, Croot and Fenton (1989, fig. 2-4).
Hill-hole pairs
49 :Ice Flow IL
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Figure 4-7. Schematic diagram showing fault-bounded depression of Wolf Lake from which material was displaced into the ice-shoved hill (gray). Ice flow from northeast toward southwest. Diagram adapted from Aber, Croot and Fenton (1989, fig. 2-6). narrow, parallel ridges occupies the hill's northern flank and aerial photographs show traces of these ridges extending down the eastern and western sides into the subsurface. Such ridges are the most typical morphologic trait of ice-shoved hills. The ridges indicate that at least the northern portion of the hill probably consists of imbricately stacked thrust blocks. Wolf Lake partially fills the depression excavated during deformation. Scarps, about 10 m high, form the eastern, western and northern margins of the lake (Fenton and Andriashek 1983). Bathymetric maps show the lake occupies three depressions corresponding to the eastern, central and western arms of the lake. Lane Lake is only a minor shallow marsh. The volume of the entire depression is only about 80% of the volume of Wolf Hill. This is believed to be the result of later infilling of the western and central arms of the lake. Comer Lake is a steep-sided depression lacking a hill on the downglacier side and is a "hill-less hole." The undisturbed terrain surrounding the hill-hole pair is essentially fiat. The model for glaciotectonic pushing by block movement of the glacier front, that was developed in northern Germany (Stephan 1985), may be applicable to Wolf Lake (fig. 4-8). Pushing of large masses occurred in front of individual ice blocks, whereas smaller ridges were pushed parallel to block side margins, which also could be the sites of strike-slip faulting. An en echelon pattern of hill-hole pairs connected at right angles by tear faults or elongated drumlins is the ideal result. In the Wolf Lake vicinity, a series of ice blocks may have been responsible for scooping the following depressions: 1)
Figure 4-8. Schematic map of block-movement model for glacier pushing. Long arrows show direction of hill-hole shoving; short arrows show direction of pushing for parallel ridges. Takenfrom Stephan (1985, fig. 3). Corner Lake, 2) eastern arm of Wolf Lake, 3) central Wolf Lake, 4) western arm of Wolf Lake, 5) and 6) smaller and larger finger lakes west of Wolf Lake. These depressions are bounded by tear faults, low ridges, or elongated drumlins. Wolf Hill was likely constructed by three of these ice blocks (2, 3 and 4). Aerial photographs of Wolf Hill do not, however, show that it is divided into three segments. Perhaps the contacts between segments are covered by sediment of englacial or supraglacial origin. In any case, the morphologic expression of Wolf Lake and Wolf Hill demonstrates a clear genetic pairing of the two. Morphologic and stratigraphic data have shown that all the features in the Wolf Lake area were formed by the southwestflowing Primrose Lobe during the last, local phase of the Late Wisconsin Cold Lake glaciation. The small lakes and hills west of Wolf Lake show evidence of smoothing by glacial flow; however, Wolf Hill, the major feature, has a fresh and unmodified appearance. The nearest ice margin, for which there is evidence, lies about 4 km west and 10 km south of Wolf Lake. Moran et al. (1980) concluded that Wolf Hill and other iceshoved hills on the Great Plains of central North America were created in a narrow (3-4 km wide) frozen-bed zone at the margin of active ice. Meanwhile, streamlined morphology was molded under a thawed bed farther upglacier. In the Wolf Lake area, the thrusting may have taken place during a stillstand or minor readvance to the previously mentioned ice margin with local molding of the thrust sediment as the frozen-bed zone moved through the area. Clearly the thrusting is a late-phase phenomenon, because Wolf Hill shows no evidence of prolonged sculpting by overriding ice.
Aber and Ber
50
Herschel Island, Yukon Herschel Island is located at the western end of Mackenzie Bay of the Beaufort Sea, northern Yukon, Canada (fig. 4-9). It is part of the Yukon Coastal Plain, a generally low, nearly flat area mostly <60 m in elevation. Herschel Island is most striking; it reaches a maximum elevation of 181 m (596 feet), covers > 100 km 2, and is bounded by steep sea cliffs >60 m high on its northern side. It is composed almost entirely of Pleistocene sediments derived from Herschel Basin and deformed by ice pushing from the southeast (Mackay 1959; Rampton 1982). Herschel Island plus Herschel Basin represent one of the largest hill-hole pairs in the world. A long ridge of similar ice-shoved Pleistocene sediments extends on the mainland from Kay Point to King Point reaching maximum elevations
~.
locally >75 m. The Yukon Coastal Plain is covered by tundra and underlain by continuous permafrost with a thickness generally >300 m (Rampton 1982). The deformed Pleistocene sediments of Herschel Island and the Kay-King Points ridge are consolidated by this permafrost. The ice-pushed structures of Herschel Island and the KayKing Points ridge are composed largely of preglacial Pleistocene sediments. No pre-Quaternary bedrock is exposed in any of these areas. The deformed strata consist of interbedded clay, silt, sand, gravel, and peat that were deposited in shallow marine, lacustrine, delta, lagoon, or flood-plain environments prior to the Buckland Glaciation. Regional correlation of these strata is hampered by discontinuous exposures, frequent facies changes, and glaciotectonic deformations. Nonetheless, a general stratigraphic framework is now available (fig. 4-10).
Mo!cotm River
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Figure 4-9. Map of northern Yukon, Canada showing geographic and glacial features. Contours in feet (lOOfeet, ~30 m); H.B. = Herschel Basin. Based on Mackay (1959, fig. 1) and Rampton (1982, figs. 1 and 18).
Hill-hole pairs
51
A~.a
Her~che|
S:tratiqra~ic Descriptions
to ~'b,b~qe
Is|and
River
=.Marina clays;
Marir~e clays
p.o~sibly contain sequence of freshwater sediments, Gel~er8 !
Malcolm Lake
Erosion ~rfaee
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ar~:
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poorly exposed,
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i:ce-wedge ca~ts. Vegetation ....
forest nearby(?),
tnterbedded
l ~ formed ~a~s: and grave 1~..
~|ay~ .,ti!i,
V e q e t a i i o n ......
V e g e t a t i ~ ..... boreal foresL forest4undra~ and shrub tundra,
•fore ~t:-tundra ahd tundra,
L.
Interbedded
~i:lt, ~and,,
and gravel wi[h peal; bed:~,
Vegetation--b~)reat forest::,
Mar, h~e clays
F r e s h w a t e r silt~ overlain and
u~derlain by marina sediments;
Vegetation ...... s~ub tundra,
s,edim en:ts f tom
~bir~e Point and rx)rthwest
sated~ 9rav~l~ arid peat,
..
Figure 4-10. Stratigraphic correlation of pre-Buckland Pleistocene sediments along the Yukon Coastal Plain between Herschel Island and Sabine Point (from Rampton 1982, Tab. 14, Erratum page). Published with permission of the Minister of Supply and Service Canada. Bouchard (1974) and Rampton (1982) divided the preBuckland strata on Herschel Island into three units: 1) upper unit consisting mostly of marine sediments, 2) a middle unit of mixed sediments, and 3) a lower marine-clay unit. Although variable in thickness, the three units may comprise up to 50 m of strata. The Buckland Glaciation is represented by pebbly till, kame deltas and terraces, various moraines, melt-water channels, and other glacial features, all subdued in morphologic expression. The Buckland ice limit descends northwestward from the Richardson Mountains, to the Buckland Hills, to just beyond the Firth River delta west of Herschel Island (fig. 4-9). Glacial erratics near the crest of Herschel Island show that the island was covered during the maximum Buckland advance.
emphasized by the presence of small, elongated lakes within intervening valleys. Elsewhere, the land slopes fairly uniformly toward the coast. Deep stream valleys and ravines form a distinctive trellis drainage pattern in the northern and central portions of the island east of the crest (Mackay 1959). Smaller ravines parallel ice-shoved ridges, but larger valleys cut across the ridges at nearly right angles. The larger ravines display a radial pattern with an apex near the concave side of the island on Thetis Bay. The valleys and ravines are the results of post-Buckland erosion of softer strata and fracture zones.
The Buckland Glaciation was presumed to be early Wisconsin in age by Rampton (1982), although Dyke and Prest (1987b) indicated a late Wisconsin age of 25,000 years BE The upper dislocated marine unit on Herschel Island was deposited during a somewhat warmer, high sea-stand before the Buckland Glaciation. This corresponds to the Pelukian transgression of Alaska, thought to be Sangamon and early Wisconsin in age. The middle and lower sedimentary units were deposited during still earlier glaciations and interglaciations. Herschel Island is overall an arcuate, asymmetrical dome, concave toward the southeast with the crest offset west of center (fig. 4-11). The northeastern and north-central margins of the island display conspicuous ridges, which form an arcuate pattern, concave southward, following the northern edge of the island (fig. 4-12). The pattern of ridges is
Figure 4-11. Topographic map of Herschel Island showing ice-shoved ridges and trellis drainage pattern. Elevations in feet; contour interval = lOOfeet (~30 m); diagonal lining shows larger lakes. Based on interpretation of aerial photographs.
52
Aber and Ber
Figure 4-12. Aerial photograph of northern and eastern Herschel Island, northern Yukon. Note ice-shoved ridges along northern margin and distinctive trellis drainage pattern. Mosaic of airphotos A24123-140 and 149 (1975). Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada. Thus, the present landscape of Herschel Island is in part an erosional morphology adjusted to underlying structures. The most common structures revealed in sea cliffs are lowangle thrust faults and open synclines and anticlines (Mackay 1959). Tilted beds usually show apparent dips of 5-20 °, but are nearly vertical in a few places. Overturned folds, repetition of beds, and inverted strata are also present in some locations. Shear planes with slickensides are especially common in clay beds. The available observations demonstrate a close agreement between internal structure and external morphology of Herschel Island. The island represents a single episode of ice pushing by an ice lobe moving from the southeast. The structures of Herschel Island include segregated ice layers and sheets, up to several m thick, which parallel the bedding of deformed strata. Crystallographic evidence proves that these ice layers developed in horizontal positions and were subsequently deformed along with the enclosing sediments by glacier pushing (Mackay and Stager 1966; Mackay,
Rampton and Fyles 1972). The ground ice probably formed when permafrost developed immediately before the Buckland Glaciation (Rampton 1982), and permafrost has persisted until the present. Based on elevation of the Buckland ice limit, Rampton (1982) constructed paleocontours on the maximum Buckland ice surface (fig. 4-11). The 500-foot contour lies just beyond the western edge of Herschel Island, so maximum ice elevation at Herschel Island is estimated to have been 600 feet (180 m). This happens to coincide with the highest elevation on Herschel Island itself, 596 feet. Ice thickness over the island may have been some greater, however, due to crustal depression beneath the ice lobe. The pre-Buckland sediments were undoubtedly permafrozen at the time of ice shoving. However, the age and thickness of the permafrost is uncertain. Climatic indicators in the pre-Buckland sediments suggest that conditions slightly warmer than today existed at certain times prior to Buckland Glaciation. Mackay (1959) and Rampton (1982) both agreed
Hill-hole p
a
i
r
that ice loading caused compression and increase in hydrostatic pressure in unfrozen sediments below the permafrost. Thrust blocks may have been detached at the base of the permafrost and pushed over a cushion of high-pressure fluid. This implies that permafrost at the time of thrusting was only 50-60 m deep, the maximum thickness of disturbed sediment. Mathews and Mackay (1960) pointed out, however, that clay may remain plastic at sub-zero temperatures. Thus, thrusting could have taken place along clay layers within still thicker permafrost.
s
5
3
Devils Lake, North Dakota Devils Lake, a large natural lake located in northeastern North Dakota, has attracted much scientific interest since the early nineteenth century. The landscape of northeastern North Dakota was heavily modified by glacier thrusting (Bluemle 1966), of which Devils Lake is a prime example. The lake occupies several connected depressions that were formed by glacier pushing of Cretaceous bedrock (shale) and glacial sediment (fig. 4-13). Sullys Hill, Crow Hill, and other hills immediately to the south are built of material thrust up from the Devils Lake depressions during late Wisconsin glacier advances. The ice-shoved hills are composed of a jumbled, brecciated mixture of deformed shale and glacial sediment.
Figure 4-13. Landsat TM image of Devils Lake vicinity showing major lake basins, ice-shoved hills (Sullys Hill Crow Hill), probable tunnel valleys (T), and spillway channels (S). DHB = Devils Heart Butte, HL = Horseshoe Lake, EDL = East Devils Lake. Based on Landsat TM band 5 (mid-infrared), 8 Sept. 2000.
54
Aber and Ber
N
OEViLS LAKE
.% \
J
Figure 4-14. Devils Lake region, northeastern North Dakota. The basins of Devils Lake represent source depressions f o r material thrust into the ice-shoved hills (Sullys Hill, Crow Hill). Geomorphic features interpretated from Landsat images, topographic maps, and ground observations. Position of lake shore based on 1988 images; ice margin positions based on Clayton, Moran and Bluemle (1980, fig. 32). T.L. = Twin Lakes; D.L.M. = Devils Lake Mountain. Adapted from Aber et al. (1997, fig. 2).
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Relief from the bottom of central Devils Lake to the top of Sullys Hill exceeds 200 m. The ice-scooped depressions and ice-shoved hills are among the largest and best-developed glacial landforms of the hill-hole type in the United States (Bluemle and Clayton 1984).
series, and the sizes of hill-hole pairs diminish away from Sullys Hill.
A major aquifer, known as the Spiritwood aquifer, trends under the Devils Lake vicinity from east-southeast to westnorthwest. This aquifer is developed partly in bedrock and partly in loose sediment filling the preglacial Cannonball valley. The major depressions of Devils Lake are located above this aquifer. High ground-water pressure in the Spiritwood aquifer likely facilitated major glaciotectonic thrusting (Bluemle and Clayton 1984; Hobbs and Bluemle 1987).
These two trends of ice-shoved hills may be interpreted as features formed by thrusting at the margins of two converging ice lobes (Aber, Spellman and Webster 1993). One lobe advanced locally from the north or northwest in the West Bay/Crow Hill vicinity, while the other moved locally from the northeast in the East Bay/Devils Heart Butte region, with Sullys Hill at the junction of the two. The two ice lobes probably advanced together, or at least at about the same time, to the North Viking and Cooperstown ice margins. However, some parts of the Crow-Sullys Hill complex may have formed in association with the Heimdal or Pekin ice margins.
An interpretation of glacial geomorphology is based primarily on visual examination of Landsat images, along with colorinfrared air photographs and conventional topographic maps, in combination with ground observations (Aber et al. 1997). Sullys Hill forms a focal point for the glacial landforms in the vicinity (fig. 4-14). Two series of ice-shoved ridges and associated source depressions (lake basins) converge at Sullys Hill. One series extends southwest to Crow Hill and on to the west. The other trends southeast, past Devils Heart Butte, to near Horseshoe Lake. These two series are associated respectively with the North Viking and Cooperstown ice margins. Ice-shoved hills are discontinuous along these two
A well-defined melt-water drainage system can be seen south of the ice-shoved hills (fig. 4-14). Five routes are: Big Coulee, Black Slough, the channel that begins at Twin Lakes, the Long Lake channel, and a series of unnamed lakes east of Big Coulee. All five channels carried melt water away from the glacier margin at the ice-pushed hills into the Sheyenne valley a short distance to the south. To the north, Sixmile Bay may have been eroded by a subglacial melt-water stream flowing toward the Devils Lake basin. Creel Bay and other linear channels may have a similar origin. These bays/valleys are here interpreted as tunnel valleys formed by melt-water erosion under the ice (see fig. 4-13).
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55
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Figure 4-15. Topographic map of Sullys Hill-Devils Heart Butte vicinity. T = tunnel valleys across Sullys Hill, S = spillway channels south of Sullys Hill. Adapted from Tokio, North Dakota 15minute quadrangle, U.S. Geological Survey (1951). Forested portions appear in gray tone. Elevations in feet; contour interval = 20 feet (~6 m). Each numbered square is one square mile (~2.5 km2).
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Narrow tunnel valleys cross Sullys Hill itself and lead directly to open spillway channels to the south (fig. 4-15). Good preservation of this assemblage of glacial landforms suggests that subsequent glaciation either did not advance over this vicinity or did little to modify the hill-hole pairs and associated melt-water landforms. The association of large ice-shoved hills, melt-water spillways, deposits of hydrodynamic blowouts~Devils Heart Butte, and tunnel valleys indicates that a substantial volume of water was released during creation of these landforms. This water was derived from both the glacier and the Spiritwood aquifer (Bluemle 1993).
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Norwegian continental shelf Hill-hole pairs are widely distributed in regions of former glaciation underlain by unconsolidated or poorly consolidated sedimentary strata. The previous examples, drawn from land and shallow-marine localities all formed in terrestrial environments, when glacial sea level was well below and beyond the sites of glaciotectonism. S~ettem (1990) demonstrated hill-hole pairs near the edge of the Norwegian continental shelf in water depths at 200-300 meters. In this setting glaciation was surely marine-based. Two sites are elaborated on the basis of bathymetric maps, shallow seismic profiles, and limited sediment sampling~Tr~enabanken and FuglCybanken (fig. 4-16).
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Figure 4-17. The "Horseshoe" at Trcenabanken. Sea floor is approximately 300 m deep in vicinicty of the ice-shoved hill which has a strongly curved outline. The presumed source depression (hole) is located on the partly enclosed, concave side of the hill. Taken from Scettem (1990, fig. 2). The "Horseshoe" is a hill with strongly curved outline on the eastern portion of Tr~enabanken (fig. 4-17). The middle, northwestern portion of the hill stands more than 100 m above the surrounding sea floor, which is about 300 m below sea level. The northern margin of the hill is a steep escarpment that extends eastward into a low ridge; the southern margin curves to the southeast also into a lower ridge. This arrangement is strongly arcuate in outline, and the concave side partly encloses a topographic basin. The hill covers an area of approximately 20 km 2, has a volume of ~1 km 3, and rests on a substratum composed of Quaternary sediment. A shallow sediment core from the hill crest revealed till, which had moderate undrained shear strength (10-90 kN/m 2) The basin associated with the Horseshoe occupies an area about 35 km 2 and has an estimated volume of 0.7 km 3. Glaciomarine or marine clay is sparse on the hill and within the basin. The Horseshoe and the adjacent sea-floor depression represent a well-preserved hill-hole pair of nearly ideal form. The steep distal (northwestern) margin of the hill and its abrupt boundary with undisturbed sea floor imply the hill was not overridden by the glacier margin subsequent to deformation.
The main portion of the hill is, thus, interpreted as a result of pushing at the front of the last glacier advance to reach the vicinity. This advance took the form of an ice tongue. The lower limbs of the hill may have formed by lateral squeezing along the flanks of an ice tongue or as material was scraped aside during advance of the ice-tongue front (see fig. 4-8). The moderate undrained shear strength of hill sediment is compatible with fast ice flow, perhaps a surge, and rapid sediment deformation. On FuglCybanken, Steinbitryggen is an irregular ridge that rises about 90 m above the adjacent sea floor (fig. 4-18). The ridge extends approximately 30 km east-west, covers an area of about 58 km 2, and has a volume of about 1.6 km 3. At the eastern end of Steinbitryggen is a depression, known as Sopphola, that covers 56 km 2 in area and has a volume of about 1.7 km 3. The main mass of Steinbitryggen has a triangular or arrowhead shape, includes several distinct ridges, and is located adjacent to Sopphola. Sopphola comprises several discrete basins. The arrangement of Sopphola and the proximal mass of Steinbitryggen suggest ice pushing from southeast to northwest. However, the point of the arrowhead is elongated into a low ridge that extends
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These two examples retain fresh topographic expression, little modified by postglacial erosion or sediment cover. The nearly equal volumes of hills and related holes corroborate the lack of postglacial alteration in these landforms. On this basis, S~ettem (1990) concluded that glaciotectonism was young, likely Late Weichselian in age. Given the site locations, this interpretation supports the reconstruction that maximum Late Weichselian glaciation reached the edge of the continental shelf. This interpretation contradicts the less-extensive Late Weichselian ice limit proposed earlier by Vorren and Kristoffersen (1986), but is compatible with later reconstructions by Mangerud et al. (1998), who believed Late Weichselian ice advanced to the continental shelf edge west of Svalbard and the Barents Sea.
This Page Intentionally Left Blank
Chapter 5 Composite ridges Introduction The most typical and distinctive glaciotectonic landforms are ice-shoved ridges found in many glaciated plains. Prest (1983, p. 45) aptly described such ridges as, "a composite of
great slices of up-thrust and commonly contorted sedimentary bedrock that is generally interlayered with and overlain by much glacial drift." The term composite ridges (= transverse-ridges of Clayton, Moran and B luemle 1980) is used here for such ice-shoved ridges. Composite ridges that include a substantial volume of deformed pre-Quaternary bedrock should not be called end moraines. Composite ridges span a considerable size range. Large composite ridges may be up to 200 m high, 5 km wide and 50 km long. In map view, composite ridges are often arcuate and concave upglacier with a radius of curvature of 2 to 10 km (Clayton, Moran and Bluemle 1980). Individual ridges typically display several 10s of m of topographic relief and are a few 100 m in width. Ridges and intervening valleys, which often contain small elongated lakes, form a subparallel pattern that follows the general curved outline of the iceshoved hill. This arcuate pattern marks the margin of the ice lobe or tongue that shoved up the ridges. The ridges are developed on the crests of folds or the upturned ends of thrust blocks. A close correspondence typically exists between structural features and topography. Large composite ridges usually involve considerable disruption of preQuaternary bedrock, which may comprise a major volume of the ridges. However, the depth of structural disturbance is generally not greater than about 200 m (Kupsch 1962). Small composite ridges (<50 m high) are perhaps the most common glaciotectonic landform. They display the same morphologic traits and structural features as do large composite ridges. Small composite ridges are also found in similar topographic settings, such as escarpments, islands, or valley sides. A source depression is located a short distance upglacier in some cases. Small composite ridges may or may not include consolidated bedrock; in fact, many are composed mostly of unconsolidated Quaternary strata. Being more susceptible to both glacial and nonglacial erosion, such ridges cannot maintain a high topographic relief. Finally, both large and small composite ridges are usually associated with ice margins marking glacier stillstands or readvances. The folds and thrust blocks that form composite ridges have usually been detached, transported some distance, and stacked into an imbricated structure (fig. 5-1). Composite ridges
are, thus, allochthonous in a glaciotectonic sense, and it may be possible to recognize the upglacier depression from whence material in the ridges was derived. However, in many cases a discrete source basin cannot be identified specifically, as with the hill-hole pair. The typical morphology and structure of large composite ridges are displayed at Prophets Mountains, a premier example, located on the Missouri Coteau in central North Dakota. Prophets Mountains reach a maximum elevation of 690 m and covers roughly 20 km 2 (fig. 5-2). Prophets Mountains consist of multiple, parallel ridges trending generally north-south with a slight arcuate tendency, concave toward the east. The overall hill is asymmetric with higher elevations and steeper slopes toward the western side. Total relief exceeds 100 m, and individual ridges are typically 1025 m above adjacent valleys. Till is the predominate glacial sediment over much of the Prophets Mountains vicinity, and a series of esker ridges is preserved along the northwestern flank of Prophets Mountains. Sub-drift bedrock in the region consists of upper Cretaceous Hell Creek Formation and lower Tertiary Cannonball Formation (fig. 5-3). Folded and contorted bedrock at the southeastern edge of Prophets Mountains rests at least 100 m above its normal stratigraphic position (Bluemle 1981). The simple arcuate pattern of ridges within Prophets Mountains corresponds to ice pushing by a single ice tongue advancing directly from the east. An adjacent source depression is not visible at the surface; however, a likely
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Figure 5-1. Schematic block diagram of typical thrust structure of composite-ridges. Arcuate pattern reflects movement of an ice lobe from the right.
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source depression is buried in the western portion of township 147, range 76 (fig. 5-3). Test drilling has revealed a closed bedrock depression of about 75 km 2 as much as 50 m below the presumed preglacial valley floor. Prophets Mountains are located above a large, complex aquifer developed partly in Cretaceous sandstone and partly in a buffed valley tributary to the preglacial Knife River (fig. 5-3). At present, ground water drains toward the east through this aquifer, but eastward drainage was blocked by glacier ice during thrusting of Prophets Mountains. Ground-water flow was reversed, and pore-water pressure built up enough to facilitate glacier thrusting, according to Bluemle and Clayton (1984).
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The term push moraine is commonly and loosely used in reference to ridges created by ice shoving. As applied here, push moraines are a restricted subset of usually small composite ridges that consists largely or wholly of glaciogenic strata. Composite ridges that contain appreciable non-glacial material should not be called push moraines. One common situation for push moraines is thrusting of contemporaneous proglacial or ice-contact drift during continued glacier advance. Many push moraines of this type were created in outwash deposits by glacier advances on Arctic islands of northeastern Canada during the Little Ice Age (K~ilin 1971). The push moraine of glacier C79 on southwestern Bylot Island is an example of a recent small composite ridge (fig. 5-4).
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Figure 5-3. Maps of Sheridan County: right - physiographic features, left - generalized bedrock topography (dashed lines) and possible preglacial drainage. Presumed buried source depression (left) and Prophets Mountains (right) marked in gray. Kh = Hell Creek Formation, Tc = Cannonball Formation. Marginal numbers indicate townships and ranges; each township is 6 miles (~10 km) on a side. Adapted from Bluemle (1981, figs. 1 and 20).
The distal portion consists of imbricately stacked scales of outwash sand and gravel. Each scale forms a low bench or terrace that slopes upglacier at a shallow angle (2-10°). Wellpreserved bedding planes within the scales dip at similar low angles, and the original stratigraphy of outwash beds is repeated in each thrust block. These distal ridges carry no trace of till and so must have been thrust in front of the ice margin. Wood twigs from the top of the innermost scale yielded a corrected C-14 date of 120+80 years BE thus establishing the maximum age for ice pushing (Klassen 1982). The higher proximal ridge stands 30-35 m above the modem sandur, is till covered, and marks the ice margin at the time of thrusting. The proximal ridge is presumably also cored by thrust blocks of outwash strata. The scales of outwash were likely thrust while permafrozen. Low ridges forming a zigzag pattern in front of the push moraine are probably patterned-ground fractures. The presence of these frost cracks may have controlled the size and initial development of the scales. Small composite ridges may be difficult to identify as glaciotectonic landforms and often have been mapped as end
moraines. Many so-called end moraines are now recognized to consist partly or wholly of ice-shoved material (Moran et al. 1980). Where disturbed bedrock is present, the glaciotectonic origin of such moraines is obvious. However, the absence of deformed bedrock does not preclude ice pushing as the primary means for constructing certain end moraines. Deformation of Quaternary strata by ice pushing is perhaps even more common than disruption of bedrock. Recognition of this situation can prove troublesome, however, as Moran ( 1971) pointed out. This is particularly true of push moraines, where the disturbed glaciogenic strata may be similar in lithology to the enclosing drift. A complete transition is possible between conventional end moraines built by primary deposition and push moraines created by secondary deformation. Only by careful examination of internal structures and stratigraphy can push moraines be identified properly, as the following Icelandic example demonstrates. Krtiger (1993) recognized a modem end moraine with stacked imbricated structure at the edge of S16ttj6kull, northern margin of M3~rdalsj6kull (fig. 5-5). The moraine consists of overlapping, frozen slabs of till derived from below the ice
62
Aber and Ber Figure 5-4. Oblique aerial photograph of push moraine in front of glacier C79, Bylot Island, Canada. Ice advance came from right to left. Proximal ridge to right is till covered (bouldery surface); distal ridges to left are tilted scales of outwash sand and gravel (smooth surface). From Klassen (1982, fig. 55.4); copyright Geological Survey of Canada (GSC 203099-P); published with permission of the Minister of Supply and Services Canada.
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63
Composite ridges
margin and dipping upglacier. The S16ttj6kull moraine was constructed in the 1980s by annual fluctuation of the stationary ice margin. Each winter the glacier margin froze onto the underlying till and transported a till slab forward and upward during a seasonal advance. This process was repeated several times in essential the same position, so that a small moraine built up (fig. 5-6). The morphology and intemal structure of this moraine closely resembles a push moraine; however, it was constructed in a wholly different manner. Kriger (1996) noted several potential criteria for distinguishing the S16ttj6kull-type moraine from typical push moraines. • Till slabs in the S16ttj6kull-type moraine represent progressively deeper stratigraphic intervals excavated from a narrow marginal zone. Whereas, push-moraine slabs are from the same stratigraphic interval which was stripped from a broad upglacier source. • Thickness of till slabs in the S16ttj6kull-type moraine corresponds to depth of annual freezing beneath the glacier
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margin. Thickness of thrust blocks in push moraines relates to a discrete d6collement, often a distinct stratum or horizon. • The source basin for till slabs in the S16ttj6kull-type moraine is several times deeper than individual till slabs. In contrast, depth of the source basin for a push moraine approximates the thickness of each thrust block. • The last-deposited portion of the S16ttj6kull-type moraine is the most proximal till slab; youngest-deformed portion of a push moraine is the most distal thrust block. Composite ridges are topographically and structurally similar to such thrust and folded mountain belts as the Canadian Rockies or Swiss Alps that were formed by thin-skinned tectonics during plate collisions. The only real difference is size, ice-shoved ridges being one or two orders of magnitude smaller than true mountains. The following examples of composite ridges are, in fact, miniature mountains produced by ice shoving of soft sedimentary strata. MCns Klint, southeastern Denmark is undoubtedly the most famous and spectacular of all glaciotectonic sites. Large scales of chalk and drift are beautifully exposed in a scenic cliff and form a rugged landscape inland from the cliff. The combination of cliff exposures and composite-ridge morphology provides a 3-dimensional display of MCns Klint's glaciotectonic structure. The Dirt Hills and Cactus Hills, southern Saskatchewan, are still larger examples of composite ridges built mainly of upper Cretaceous bedrock. Flade Klit in Denmark is another famous site that includes a substantial volume of deformed Eocene strata. In contrast, Utrecht Ridge in the Netherlands and Brandon Hills, Manitoba are composed wholly of unconsolidated Pleistocene strata. Additional Polish examples of classic composite ridges are Ostrzesz6w Hills (chap. 9) and Trzebnica Hills (chap. 11).
Mons Klint, Denmark
~.~ ,.,--.; ;.~ ~.~ ~;=.q .~,
~'i4
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Figure 5-6. Model f o r sequential development of small end moraine at Sl(ttjOkull. 1 - glacier, 2 - lodgement till, 3 mass movement deposit, 4 - outwash, 5 - thrust plane, 6 annual moraine. Taken from Kriiger (1994, fig. 87).
The large chalk cliffs of eastern Men, southeastern Denmark are justifiably famous for their scenic beauty and distinctive geological structure (fig. 5-7). MCns Klint has been called the cradle ofglaciotectonics (Pedersen 1994). The cliffs are perhaps the finest example of glaciotectonic features in the world, and they have a long history of geological study beginning with Agricola (1546). The earliest known illustration appeared in the 1700s (see fig. 1-2), and the copper engraving of MCns Klint by Puggaard (1851) remains a classic section of the cliff (fig. 5-8). Johnstrup (1874) first recognized the glaciotectonic genesis of MCns Klint, and the first map showing ice-shoved ridges on Men was published by Haarsted (1956). The high, rugged terrain, known as HOje Men, located at the eastern end of the island is composed of several dozen chalk scales that were piled up during late Weichselian ice advances
Aber and Ber
64
Figure 5-7. Photographs of southern MCns Klint from the beach. A - chalk masses of Sommerspiret (center) and Ncelderendenakke (left) standing >100 m high. The pinancle of Sommerspiret fell down during a storm in the early 1990s. B - Grcederen chalk mass with pinancle 82 m high. Photos by J.S. Aber (1986). 20,000 to 13,000 years BP (SjCrring 1981). Hcje Men generally exceeds 75 m elevation reaching a high at 143 m (fig. 5-9). Dronningestolen (the Queen's throne), the largest chalk cliff in the center of MCns Klint, is 128 m high (fig. 510). Presumably undisturbed chalk bedrock is situated 20 to 40 m below sea level (Haarsted 1956; Jensen 1993), so >160 m of structural relief is indicated.
~H
, '
.
The individual scales exposed in MCns Klint consist of upper Cretaceous (Maastrichtian) white "writing chalk" that was deformed along with drift. Chalk now forms ridges and cliffs because of its greater resistance to erosion, whereas intervening drift has been eroded into valleys that form the falls along the coast (fig. 5-11). The chalk is quite uniform in lithology, aside from occasional layers of flint nodules, and thus, stratigraphic correlation between chalk masses is
H
H
~
........
~ T H
Figure 5-8. MCns Klint section as viewed from the east; Jcettebrink at southern end (upper left), Slotsgavle at northern end (lower right). Black lines within blank chalk masses show deformed flint layers. Reproduction of copper engraving by Puggaard (1851); adapted from International Geological Congress XXI Session, Norden, Guidebook I (1960).
65
Composite ridges
difficult at best. Surlyk (1971) divided the Danish Maastrichtian into ten brachiopod biozones: zones 1-7 are lower Maastrichtian and 8-10 are upper Maastrichtian. These biozones are the best method for establishing correlation between chalk masses.
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Figure 5-9. Topographic map of MOns Klint and HOje MOn vicinity, southeastern Denmark. Individual chalk cliffs and inter-chalk falls are indicated with heights of selected cliffs in parentheses. Contour interval = 25 m; Ht. = hotel. Adapted from Abet (1985, fig. 1).
Figure 5-10. Photograph of central MOns Klint from the beach showing chalk mass of Dronningestolen. Chalk cliff is >125 m tall. Photo by J.S. Aber (1986).
The drift strata are well exposed in only a few places, notably at Hundevmngsfald and Sandfald. The composite sequence includes from the top down (Hintze 1937; Surlyk 1971): 1) discordant drift, 2) upper dislocated till, 3) stone-poor clay, 4) cross-bedded sand, and 5) lower dislocated till. Konradi (1973) concluded that the lower dislocated till and all the overlying drift must be Weichselian in age, because the lower till contains abundant reworked Eemian foraminifera. The stone counts carried out by Hintze (1937) showed the lower dislocated till had a Baltic source, whereas the upper dislocated till was deposited from the northeast by an ice advance which caused considerable local erosion. Both of these fills were deposited before the southern portion of MCns Klint was thrust. MCns Klint and HCje Men can be divided into three morphostructural regions on the basis of ridge morphology, cliff structures, and chalk stratigraphy (Aber, Croot and Fenton 1989). The southern region includes Jmttebrink through Sommerspiret, with chalk biozones 3, 4 and 5. The chalk scales form a series of imbricately thrust anticlines that dip southward and are increasingly deformed toward the north. These chalk masses continue inland as long, straight to arcuate ridges, concave toward the south. This region was thrust by ice movement directly from the south. Pedersen (2000) has designated this section the Grhryg imbricated zone (fig. 5-12). Dronningestolen along with Gr~ederen and Maglevandspynten make up the central portion of the cliff. Dronningestolen is a huge composite of many lesser chalk floes (biozones 3 through 8) folded and stacked on top of each other in the overall form of a broad anticline. Dronningestolen continues inland as a massive ridge, beyond which the central region is marked by many short, offset ridges. This zone includes the Dronningestolen antiformal stack above the Maglevandsfald ramp (fig. 5-12). It was deformed primarily by ice pushing from the south and perhaps also from the northeast. The northern region of Vitmunds Nakke through Slotsgavlene includes biozones 5 through 8 in scales oriented oblique to the coast. A regular shift is displayed in structural strike along the cliff from southeast in the south, to south in the center, to southwest in the north. This corresponds to the arcuate pattern of ridges inland from the cliff, which were deformed by ice pushing from the east or ENE. Pedersen (2000) identified this portion as the Slotsgavle foreland thrust zone (fig. 5-12).
66
Aber and Ber
Figure 5-11. Aerial photograph of MCns Klint and HCje Men, southeastern Denmark. Chalk masses form high cliffs and continue inland as sharp-crested ridges. More rugged portion of ridges covered by beech forest; presence of chalk masses in fields shown by light tone. Compare with topographic map (fig. 5-9). Reprocessed from Geodcetisk Institut, Denmark, 1974 (a. 29/88). The probable source for chalk masses now found in MCns Klint and Hcje Men is offshore in the shallow Baltic Sea. Jensen (1993) recognized an arcuate, horseshoe-shaped basin that encloses the southeastern end of the island, which he identified as the likely source depression (figs. 5-13 and 14). The depression is located circa 5 km offshore, which corresponds to the distance Pedersen (2000) calculated for balancing the displacement of chalk floes in MCns Klint. The island of Men is positioned above a basement high, known as the Men Block, which is an eastern extension of the Ringkcbing-Fyn-Falster High (Baartman and Christensen 1974). Below the eastem end of the Men Block is a faultbounded horst that corresponds exactly with the presumed source depression (fig. 5-14). On this basis, Jensen (1993)
proposed that thrusting of chalk was facilitated by a buried basement obstacle in the path of ice advancing from the southeast. Various models have been proposed over the years to account for two directions of glaciotectonic thrusting at MCns Klint. The earliest idea involved simultaneous deformation from two directions between two ice lobes (Gripp 1947). Jensen (1993) revived this interlobate interpretaton. He envisioned glaciotectonic compression from the northeast and southeast between two ice lobes during a final Weichselian readvance. However, few have accepted this model in recent years. Most geologists ascribe deformation at MCns Klint to multiple phases of ice advance based on regional patterns of glaciation.
Composite ridges
67
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Figure 5-12. Schematic diagram of structural zones within MCns Klint glaciotectonic complex. Imbricate thrusting dominates the Gr&ryg and Slotsgavle zones, which are separated by the Dronningestolen antiformal stack and Maglevandsfald ramp. The main phase of thrusting came from the SSE, and was followed by superimposed thrusting from the ENE. Takenfrom Pedersen (2000, fig. 11). The Weichselian glaciation of southern Denmark occurred in distinct phases separated by brief ice retreats (Berthelsen 1978; Houmark-Nielsen 1987): 1) Old Baltic advance from the southeast that deposited the Ristinge Klint Till, but caused little glaciotectonic disturbance; 2) Main Weichselian phase with initial advance from the northeast (Mid Danish Till) and Storeba~lt readvance from the east (North Sja~lland Till), both of which resulted in major glaciotectonic disruptions; 3) Young Baltic advances coming first from the southeast (East Jylland Till) with late readvances by ice tongues along major straits from the south (Ba~lthav Till), both of which created much ice-push deformation. In the latest phase, distinct ice lobes existed in the Storeb~elt and Oresund (fig. 5-15). The upper and lower dislocated tills at MCns Klint are correlated respectively with the Mid Danish and Ristinge Klint Tills. The overlying discordant drift probably relates to the Young Baltic advances, which overrode Men following the thrusting of composite ridges. Thrusting of the southern (Gr~ryg) and central (Dronningestolen) portions of MCns
Figure 5-13. Contour map on glacial surface morphology in the Baltic region around the island of Men. Asterisks indicate position of probable source basin for material shoved into MCns Klint and HOje MOn. Land areas shaded gray. Adapted from Jensen (1993, fig. 15). Klint is connected to the Young Baltic advance from the southeast or south, probably by the Storeba~lt ice lobe. Berthelsen (1979) and Aber, Croot and Fenton (1989) suggested that deformation from the northeast (Slotsgavlene zone) was caused by the earlier Main Weichselian ice advance. However, Pedersen (2000) has collected structural data that support a younger age for what he termed the Jydelejet phase of superimposed thrusting from the east (fig. 5-12). This was presumably caused by a final readvance of the Oresund ice lobe. D i r t Hills a n d C a c t u s Hills, S a s k a t c h e w a n
Large composite ridges of the MCns Klint type are well developed in southern Saskatchewan and adjacent Alberta on the Canadian Plains. The greatest development of composite ridges is found in the Dirt Hills and the Cactus Hills (areas 2 and 3, fig. 5-16). The Dirt Hills reach a maximum elevation of 880 m (2887 feet), some 300 m above the Regina Lake Plain to the north and 150 m above the adjacent Missouri Coteau. The Cactus Hills are nearly as high in elevation. The Dirt Hills and Cactus Hills together encompass a region approximately 1000 klTl2 in extent (see fig. 2-21). The Missouri Coteau is a major northeast-facing escarpment that marks the boundary between the Saskatchewan and
Aber and Ber
68
!i:
SJ~LLARO
\ M~N
q
Q
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Figure 5-14. Structural contour map with faults and anticline on presumed basement (top of pre-Zechstein horizon) in vicinity of Men, western Baltic Sea. Eastern end of the Men block (horst) marked by asterisk, which corresponds exactly to location of probable source for chalk thrust into MCns Klint and HCje Men. Modified from Baartman and Christensen (1974, plate 4); copyright Geological Survey of Denmark and Greenland.
Alberta Plains. The Missouri Coteau as well as lower hills to the northeast are bedrock features resulting from preglacial erosion and subsequently modified by glaciation. A mantle of drift covers most of the region today.
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The disturbed bedrock structures of the Dirt Hills were apparently first noted by Bell (1874), and Fraser et al. (1935) described faulted bedrock in the northern Dirt Hills near Claybank. The true glaciotectonic origin of the Dirt Hills and Cactus Hills was first demonstrated by Byers (1959),
20 km
Terminal moraine
Higher elevation of the Alberta Plain is a reflection of more resistant terrestrial sandstone interbedded with mudstone and lignite. The Saskatchewan Plain, in contrast, is underlain by softer marine shale. The regional geologic structure consists of essentially flat-lying strata which dip gently to the east or northeast (Fraser et al. 1935). Steeply dipping, folded, and faulted bedrock structures are common, however, in iceshoved hills along the Missouri Coteau.
~Ib Ice 'flowdLreclion
Figure 5-15 (left). End moraines and ice-flow directions related to final phase of late Weichselian glaciation in southeastern Denmark. 1 = Storebcelt ice lobe, 2 = Oresund ice lobe. Asterisk indicates location of MCns Klint. Adapted from Jensen (1993, fig. 2).
69
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Figure 5-16. Ice-shoved hills o f southern S a s k a t c h e w a n and adjacent Alberta, Canada. 1 = Radville area, 2 = Dirt Hills, 3 = Cactus Hills, 4 = Chaplin Lake area, 5 = The Coteau, 6 = L a n c e r area, 7 = M o n i t o r area, 8 = Neutral Hills, 9 = Killarney Lake area. B a s e d on Kupsch (1962, fig. 1).
who believed the deformations had been produced by subglacial frictional drag. A general analysis of the genesis of these hills along with similar ice-shoved hills in western Canada was presented a short time later by Kupsch (1962). Christiansen (1961) and Parizek (1964) described the ice-pushed ridges in conjunction with regional resource mapping, and the great depth of structural disturbance was confirmed by test drilling on the crest of the Dirt Hills (Christiansen 1971 a; Christiansen and Whitaker 1976). Three upper Cretaceous bedrock units can be seen deformed in the Dirt Hills and Cactus Hills. These three in ascending order are (fig. 5-17): Eastend (Ke), Whitemud (Kw), and lower Ravenscrag (Klr). The deformed strata are predominately terrestrial, bentonitic or kaolinitic sandstone, plus mudstone and lignite, having a total thickness up to 90 m. They are underlain by presumably undisturbed marine shale of the Bearpaw Formation (Kb). Folded and thrust bedrock scales stacked in an imbricated pattern comprise the overall structure of the Dirt Hills and
Cactus Hills (fig. 5-18). At most sites, drift covers the bedrock but is not involved in the deformations. The drift is dominated by dolostone and crystalline erratics in the pebble fraction. However, at two sites (6 and 9) older blocks of drift rich in quartzite and chert are exposed as a result of glaciotectonic uplift. Not only have scales been displaced horizontally, but considerable vertical movement has also occurred. Maximum structural uplift is documented at site 9, where a block of Eastend Formation is standing vertically some 200 m above its normal stratigraphic position (Aber 1993c). Remarkable agreement exists between orientations of bedrock structures, trends of individual ridges, and overall outlines of the Dirt Hills and Cactus Hills. This amply confirms Kupsch's (1962) conclusion that ice-shoved ridges on the Missouri Coteau are direct or first-order morphologic expressions of bedrock structures produced by ice pushing. Bedrock structures at most sites are related to a single episode and direction of ice pushing, but some sites (fig. 5-18, site 13) show evidence for multiple phases of deformation. Variations in bedrock competence clearly influenced structural development. Thrust faults are usually located
Aber and Ber
70
Twm:
...........................
Figure 5-17 (left). Stratigraphic column for bedrock of the Dirt Hills and Cactus Hills vicinity. Deformed units shown by gray shading. Based on Fraser et al. (1935) and Parizek (1964).
. . . . . . .
within lignite or claystone beds; conversely, thicker sandstone layers comprise the larger folds and fault blocks.
Tu r i
. ........................................................
Given the fact that ridges are first-order morphologic expressions of bedrock structures, the primary regional structure of the Dirt Hills and Cactus Hills can be determined directly from topography. Ridge morphology is particularly evident on aerial photographs (fig. 5-19) due to the relatively dry climate and sparse vegetation.
iiliiii!iiii i!i;i!ii! !iiiiiii i!iiii:/!i iiiii iiiii!i!i!
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Three classes of ice-shoved ridges are identified on the basis of topographic expression and structural information. The three ridge classes include a subdued class I and two prominent classes, II and III (fig. 5-18). Ridges of classes I and II were overridden and smoothed by ice, whereas the highest class III ridges of the southern Dirt Hills were never overrun. Class I ridges were presumably created during an earlier glacier advance of unknown age. Class II and III ridges were thrust up by the last ice advance to push onto the Missouri Coteau.
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Figure 5-18. Map of Dirt Hills and Cactus Hills vicinity showing pattern of ice-shoved ridges, locations of sites, structural features (circles), and ice tongues: G = Galilee, S - Spring Valley, A = Avonlea. Modified from Aber (1993c, fig. 3).
Composite ridges
71
Figure 5-19. Aerial photograph of northern Dirt Hills, Saskatchewan. Numbered sites (1-4, fig. 5-19). Tiny elongated lakes occupy narrow valleys between ice-shoved ridges. Aerial photo A21657-20 (1970). Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada. The Dirt Hills and Cactus Hills form two large loop-shaped ranges which along with nearby hills define the outlines of three ice tongues that caused the thrusting of bedrock ridges. These three ice tongues were: 1) Galilee, west of Cactus Hills, 2) Spring Valley, between Cactus Hills and Dirt Hills, and 3) Avonlea, east of Dirt Hills. Most of ice-shoved ridges of the Missouri Coteau are presumed to be late Wisconsin in age, based on incorporation of interstadial strata of mid-Wisconsin age. For example, deformed ridges of the Lancer area (fig. 5-16, site 6) include organic soil and peat that yielded an uncorrected radiocarbon date of 31,300__.1400 years BP (Aber 2006). Thrusting of the Dirt Hills and Cactus Hills did not happen during initial advance of the late Wisconsin Lostwood ice sheet, which reached its maximum position in Montana and North Dakota about 20,000 radiocarbon years ago (~23.4 ka cal. age; Fullerton, Colton and Bush 2004). A subsequent major regional readvance to the Wood Mountain vicinity took
place about 14,000 radiocarbon years ago (~16.4 ka cal. age). At that time, the Dirt Hills and Cactus Hills did not yet exist. Thrusting of the hills occurred later, probably around 13,000 years BP (Fenton, Moran et al. 1983; Dyke and Prest 1987b), during a strong readvance of the Weyburn lobe (Christiansen 1956). This lobe generated lateral ice tongues that pushed into embayments of the Coteau. Thrusting of bedrock occurred around the margins of these ice tongues due to rapid loading and forward movement. All three ice tongues caused thrusting of class I bedrock ridges during an earlier advance, however the time and nature of this advance are uncertain. The main thrusting of class II and III ridges occurred during a readvance of the Galilee and Spring Valley ice tongues. These ice tongues overran class I ridges and thrust up new ridges to the south forming higher portions of the Dirt Hills and Cactus Hills. The Avonlea ice tongue also readvanced at this time, but without thrusting up any new ridges.
Aber and Ber
72 The Galilee and Avonlea ice tongues reached positions marked by the Ardill end moraine and Lake of the Rivers spillway during the readvance. The Spring Valley ice tongue, however, stopped on the inner (northern) side of the Dirt Hills, from where a spillway was cut across class III ridges toward Skyeta Lake (fig. 5-20). The class III ridges, thus, formed a nunatak between active ice to the north and older stagnant ice lying on the Coteau to the south. Building of the Ardill end-moraine system and cutting of associated spillways were related to the same ice advances that caused the main phase of thrusting in the Dirt Hills and Cactus Hills. Following thrusting of the hills, building of the end moraine, and cutting of the spillways, the ice tongues stagnated and downwasted leaving an irregular accumulation of hummocky moraine over much of the area north of the ice margins. Rapid loading of competent sandstone bedrock (lower Ravenscrag, Whitemud, and upper Eastend Formations) over saturated, incompetent mudstone strata (lower and middle Eastend Formation) caused thrusting around the margins of the ice tongues. The fact that bedrock was most likely thawed and saturated is confirmed by abundant evidence for meltwater spillways, large proglacial lakes, and wasting stagnant ice masses throughout southern Saskatchewan during
deglaciation. There is no evidence that permafrost existed during the late Wisconsin deglaciation. Ubiquitous meltwater means that thrust blocks could not be moved by freezing onto the undersides of the ice tongues, but were displaced by squeezing out from under the ice margin.
Flade Klit, Denmark During late Weichselian glaciation, Eocene bedrock and glacial strata were folded and thrust into composite ridges in several portions of the western Limfjord district, northwestern Denmark (fig. 5-21). This glaciotectonic region is well known, beginning with the work of Gry (1940). The bedrock, consisting of clayey diatomite interbedded with volcanic ash layers, was especially susceptible to ice-push deformation. Flade Klit on the island of Mors is a prime example of composite ridges in this setting, and the cliff section exposed at Hanklit is particularly noteworthy. Flade Klit is part of a composite-ridge complex that includes several morphologic arcs (fig. 5-22). The Flade Klit composite-ridge arc is approximately 3.5 km long (E-W) and 1.5 km wide. It forms a gentle crescent, concave toward the north (figs. 5-23 and 24). Maximum elevation within Flade Klit reaches 88 m at SalgerhCj, which is 100 m above the floor of Thisted Bredning of the Limfjord
N
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900
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800
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Figure 5-20. Schematic model for thrusting of southern Dirt Hills. North-south profile of present land surface through high ridges west of Skyeta Lake spillway (shown by dashed line; numbered sites same as fig. 5-18). Basal d~collement is located in claystone beds of the lower Eastend Formation. Profile of ice tongue based on assumed minimum thickness of 300 m at northern end of section. Adapted from Aber (1993c, fig. 17).
Composite ridges
73 Gullerup thrust sheet to the south. The Hanklit thrust sheet has the form of a broad anticline with a sharply overturned and stretched nose toward the south. The fold nose displays brittle fracturing of competent beds and plastic squeezing of ductile layers (see fig. 1-15). The Fur Formation is fiddled with fractures (joints) and small faults. Most are connected with glaciotectonic deformation. However, some normal faults were pre-existing structures, as they are truncated by the thrust fault at the base of the sheet. Thrust zones above and below the Hanklit sheet display intensive deformation with slickensides, strong shearing and intrusive structures.
estuary immediately to the north. The estuary basin is presumably the source of material now in Flade Klit, although a specific source depression has not been identified. Structure of the Hanklit thrust complex was investigated in detail by Klint and Pedersen (1995). They identified three thrust sheets within the Hanklit section. Each thrust sheet consists of Fur Formation overlain by glaciogenic strata with a total thickness of ca. 60 m (fig. 5-25). The lower Eocene Fur Formation consists of clayey diatomite layers, known as "mo-clay," composed of diatoms and montmorillonite clay. This sediment is highly porous and has a density of ca. 0.8 g/cm 3. Mo-clay is interbedded with basaltic tephra. Each tephra bed is numbered in a unique sequence that allows precise correlation from site to site. The Hanklit glaciogenic group has a base of coarse, glaciofluvial gravel that fines upward into a glaciolacustrine facies of clay, silt and sand beds. Till caps the glaciogenic sequence. It is massive, clayey and has a few scattered stones, mainly chalk and flint. The till is interpreted as a subglacial lodgement deposit derived from the north.
As constructed by Klint and Pedersen (1995), the SalgerhCj and Gullerup thrust sheets have similar overall structures, although the Gullerup sheet likely ends southward in a blind thrust (fig. 5-26). The basal d6collement is projected ca. 60 m below sea level, and overall horizontal shortening of the thrust section is approximately 40%. A geo-electric survey indicated these thrust sheets extend at least one kilometer inland to the east (fig. 5-27). JCrgensen et al. (2005) traced glaciotectonic disturbance as low-resistive structures in TEM maps to ca. 30 m below sealevel at Hanklit.
The Hanklit thrust sheet forms the major exposed portion of the Hanklit section (fig. 5-26). A remnant of the SalgerhCj thrust sheet overlaps its northern end, and it rests on the
Given the general geological setting, two phases of late Weichselian glaciation might have caused the thrusting displayed at Hanklit: 1) ice pushing directly from the north,
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km
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Figure 5-21. Notable composite-ridge systems in the western Limfjord district, northwestern Denmark. Locations of Hanklit and Flade are indicated on the island of Mors. Modified from Klint and Pedersen (1995, fig. 1); copyright Geological Survey of Denmark and Greenland.
Aber and Ber
74
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Figure 5-22. Composite ridges in the northwestern portion, island of Mors, including the Mosebjerg, StcerhCj, Bjergby, and SalgerhCj morphologic arcs. Position of the Hanklit profile is indicated (fig. 5-26). Modified from Klint and Pedersen (1995, fig. 4); copyright Geological Survey of Denmark and Greenland.
Figure 5-23. Aerial photograph of Flade Klit vicinity, island of Mors, northwestern Denmark. Composite ridges form an open crescent, concave northward. Location of cliff exposure at Hanklit noted by asterisk. Reprocessed from Geodcetisk Institut, Denmark, 1986 (A. 29/88).
Composite ridges
75
Sa!~erh~i
Figure 5-24. Kite aerial photograph of Flade Klit vicinity, island of Mors; view toward northeast. Hanklit is the cliff exposure on the western end of the ice-shoved ridge that has its crest at SalgjerhOj. Photo by S.W. Aber (2005). U L L.E~i~P T~ ~ L~~T
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The positions of SalgjerhCj and other glaciotectonic complexes on Mors are closely related to buried valleys (fig. 5-28). JCrgensen et al. (2005) identified four generations of buried valleys, the youngest of which was formed during the Saalian glaciation or earlier. These incised valleys are filled with various till, glaciolacustrine clay, and glaciofluvial sand and gravel sediments from Weichselian, Saalian, and Elsterian glaciations. The main glaciotectonic d6collement is situated within the Holmehus Formation, below the Fur Formation, throughout northern Mors (see fig. 3-19). Within buried valleys, however, the d6collement zone and overlying Paleogene strata have been removed by erosion. Thus, glaciotectonic complexes are found outside and between buried valleys.
Utrecht Ridge, Netherlands L,,+ ............. ,!...... {
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Figure 5-25. Stratigraphic profiles for the SalgerhOj, Hanklit and Gullerup thrust sheets at Hanklit. Each thrust sheet is comprised of Eocene Fur Formation below and Hanklit glaciogenic group above. Modified from Klint and Pedersen (1995, fig. 5); copyright Geological Survey of Denmark and Greenland. the so-called Norwegian advance, and 2) the main Weichselian advance from the northeast. Klint and Pedersen (1995) interpreted the dislocated Hanklit glaciogenic group as probably Saalian or early Weichselian in age. Thrusting of Fur Formation and glacial strata then took place during the late Weichselian Norwegian advance, which subsequently overrode the Flade Klit vicinity.
A major zone of Saalian glaciotectonism trends W-E across the central Netherlands (Laban and van der Meer 2004). The largest composite ridges stand >100 m above sea level. In relation to adjacent source basins, total relief exceeds 200 m in some cases (van der Meer 1988-89). Among these impressive glaciotectonic landforms is Utrecht Ridge, located on the southwestern edge of Gelderse Vallei in the central Netherlands (fig. 5-29). It is one of a series of ridges composed of imbricately thrust scales of unconsolidated Pleistocene strata that are situated around three sides of a glacial basin located in Gelderse Vallei (Ruegg 1991). Utrecht Ridge trends northwest-southeast, and it rises >50 m above the adjacent lowlands, forming a conspicuous sight in an otherwise flat landscape. Utrecht Ridge was created when an ice lobe excavated Gelderse Vallei basin. The ice-shoved hills are heavily eroded as a consequence of their Saalian age combined with unconsolidated character of Quaternary strata and are mere ruins of the original landforms. Kwintelooijen sand pit is located on the inner or northeastern side of Utrecht Ridge
Aber and Ber
76 N
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Figure 5-26. Geologic profile of the exposed portion of Hanklit (above) and constructed section of thrust sheets (below). The basal d#collement is about 60 m below sea level. Overall shortening is ~40%. Modified from Klint and Pedersen (1995, fig. 15); copyright Geological Survey of Denmark and Greenland. /'i Gk~.ciolacust~me and fluv a
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Composite ridges
77
Figure 5-28 (right). Buried valleys on the island of Mors, northwestern Denmark, identified on the basis of large-scale TEM mapping. H-S = Hanklit-SalgjerhOj glaciotectonic complex. Asterisk (*) shows location of young buried valley immediately west of H-S. Adapted from JOrgensen et al.
Valley generations Youngest generation Oldest generation
(2005, fig. 9).
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(fig. 5-30). The pit came under coordinated, interdisciplinary study during the 1970s and early 1980s on account of the many Paleolithic artifacts and fossils found there (Ruegg and Zandstra 1981).
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The sedimentary strata now forming Utrecht Ridge were originally deposited as alluvium of the ancestral Rhine/Meuse Rivers. Kwintelooijen sand pit contains three formations, in ascending order: Kedichem, Urk, and Drente (Ruegg 1981). The Kedichem Formation consists of very fine sand, clay, peat, and loam deposited in a flood-plain/back-swamp environment. Sediments of the Urk Formation are mainly fine to coarse sand, gravelly sand, and coarse gravel deposited by the ancestral Rhine River (Zandstra 1981). The glaciofluvial Drente Formation contains fine to coarse sand, gravelly sand, and gravel. These formations total >20 m in thickness.
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The Drente Formation, by definition, includes all drift related to the Saalian glaciation. At Kwintelooijen, it accumulated as ice-marginal and sandur deposits in front of glacier ice.
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Figure 5-29. Ice-pushed hills and glacial basins of the central Netherlands. K = location of Kwintelooijen sand pit within Utrecht Ridge on the southwestern side of Gelderse Vallei. Based on van den Berg and Beets (1987).
Aber and Ber
78
features demonstrate that Utrecht Ridge probably was not overridden by the Gelderse Vallei ice lobe (van der Wateren 1985), although post-Saalian erosion may have removed an original till cap. Utrecht Ridge consists of thrust blocks striking parallel to the ridge and dipping on average 35 ° to 40 ° NNE (fig. 531). Thickness of scales varies from about 25 m to only a few m, but each includes a basal portion of fine-grained sediment of the Kedichem Formation. Thrust blocks are imbricately stacked and gently folded. Thrusts at the base of each scale contain shear planes, small isoclinal folds, breccia and slickensides in a zone of intermingled sediments several dm thick. Many normal faults forming conjugate sets are also present.
Figure 5-30. Topographic map of Utrecht Ridge showing location of Kwintelooijen sand pit (K) and position of section (fig. 5-31). Elevations in m; contourinterval = 5 m. Modified from van der Wateren (1981, fig. 1).
Much of the Drente sediment was probably reworked from underlying Urk Formation deposits. In southern Gelderse Vallei, Saalian till is buried 15-30 m below sea level. A test boring made at the bottom of the sand pit penetrated thrust and contorted strata to a depth of about 24 m below sea level (Zandstra 1981). Utrecht Ridge near Kwintelooijen is about 2.5 km wide with a plateau top between 45 m and 60 m above sea level (fig. 530). The ridge is a s y m m e t r i c in cross profile. The northeastern side slopes 5-15 ° , whereas the southwestern flank slopes only 2-5 ° . This difference is partly explained by the presence of sandur deposits covering the southwestern portion of the ridge (van der Wateren 1981). Utrecht Ridge and other ridges surrounding Gelderse Vallei are cut by several dry valleys representing former spillways, and the ridge plateaus show no trace of till. The morphologic
The present elevation of Utrecht Ridge is less than its initial elevation, as a result of lowering by normal faulting and by later erosion. Van der Wateren (1985) estimated that scales were initially pushed up at least 100 m above the basal drcollement. Structural restoration results in a good balance between volumes of Utrecht Ridge and the excavated basin of southern Gelderse Vallei. Van der Wateren (1981) calculated that potential shear stress developed at the glacier sole would be far too small to upthrust blocks 100 m above the drcollement. Instead of shear stress, the lateral pressure gradient caused by differential loading of the substratum during ice advance provided the driving force for glacier thrusting. A series of deep glacial basins behind ice-shoved hills stretches across the central Netherlands (fig. 5-29). The basins and associated hills increase in size from west to east. Owing to greater subsidence, ice-pushed ridges of the western basins are now mostly buried beneath younger sediments. These basins exceed 100 m depth and are floored with Saalian till in many places. The basins also contain tunnel valleys that appear to lead toward spillways breaching the marginal ice-shoved hills.
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Composite ridges
79
Prior to the Saalian glaciation, the Netherlands was an alluvial plain with slight topographic relief, having no major valleys or hills. The northern Netherlands was underlain by mostly fine-grained sediments over which the Saalian ice sheet advanced easily. As the ice sheet reached the central Netherlands, it overran coarser, gravelly sediment (Urk Formation) of the ancestral Rhine/Meuse Rivers. This caused an increase in basal friction, thickening of the ice sheet, and melt-water erosion of subglacial basins and tunnel valleys. The marginal effect of this modified subglacial topography was to generate ice lobes in a region where the landscape was previously flat. The thicker ice sheet in combination with melt-water erosion of subglacial basins and development of ice lobes was an ideal situation for thrusting of composite ridges above a basal d6collement in the Kedichem Formation. Two mechanisms contributed to thrusting: 1) lateral pressure gradient due to differential loading (van der Wateren 1985), and 2) direct glacier pushing against the sides of the basins (van den Berg and Beets 1987). Thus, the glacial basins were created by joint subglacial melt-water erosion and glaciotectonic thrusting. Certain factors remain unresolved, for example the nature of permafrost. Most Dutch geologists have assumed that deformation took place in permafrozen sediments and that the thickness of scales is an indication of the depth of permafrost (de Jong 1967). Van der Wateren (1985), however, challenged the assumption of permafrost as a prerequisite for ice thrusting.
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Figure 5-32. Map of the Netherlands showing locations of ice-shoved hills and directions of ice movement. Five phases of ice-shoving are noted from south to north (a-e). Adapted from Aber (1985, fig. 2).
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g
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Jelgersma and Breeuwer (1975) explained the origin of glacial basins as deep erosion caused by glacier surging, and this is hardly compatible with ice advance over permafrost. Another possibility is that the ice sheet was cold-based during the advancing stage. Later, when the ice warmed and became more mobile, glacier lobes cut deep basins and pushed up ridges (ter Wee 1983). Five phases of Saalian ice pushing were recognized previously in the Netherlands (fig. 5-32), three in the central portion and two in the northeastern region (Maarleveld 1953; ter Wee 1962). According to the traditional interpretation, the first (oldest) phase (a) marks the maximum Saalian ice coverage, when Utrecht Ridge was built (Ruegg 1991). Each younger phase represents a readvance of uncertain magnitude during Saalian deglaciation. During each of these phases, the ice margin was highly irregular, with ice lobes extending beyond the main inland ice sheet.
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Figure 5-33. Physiographic features of southwestern Manitoba. Location of Brandon Hills shown by solid dot at northern end of Tiger Hills upland. Adapted from Klassen (1979, fig. 1).
Aber and Ber
80
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Figure 5-34. Topographic map of Brandon Hills vicinity showing glacial features. Based on interpretation of aerial photographs. Provincial highways shown by numbered circles; H.R. = Hydraulic Road. Elevations in feet; contour interval = 50 feet (~15 m). Locations of exposures shown by solid dots. Adapted from Aber (1988b, fig. 4). and van der Meer (2004) concluded that none of these models can be accepted uncritically at present.
Brandon Hills, Manitoba Brandon Hills comprise a small group of subparallel ridges, located at the northern end of Tiger Hills upland, about 11 km south of the city of Brandon (fig. 5-33). The Tiger Hills upland is part of the Manitoba Escarpment, located near the eastern edge of the Saskatchewan Plain. This escarpment rises abruptly as much as 300 m above the Manitoba Plain to the east, and is marked by a series of uplands--Duck Mountain, Riding Mountain, and Tiger Hills. These uplands are cored by upper Cretaceous shale of the Riding Mountain Formation. The bedrock of southwestern Manitoba is almost completely mantled by thick drift consisting mainly of till. Stratified drift is abundant on the uplands, however, where ice stagnated during deglaciation. The glacial geomorphology and stratigraphy of southwestern Manitoba were described in several investigations (Klassen 1975, 1979; Fenton, Moran et al. 1983; Fenton 1984). Brandon Hills have usually been mapped as part of the end moraine atop Tiger Hills upland (Prest, Grant and Rampton
1967). Welsted and Young (1980), however, questioned designation of Brandon Hills as an end moraine, because they found that much of the hills consists of stratified drift. They also considered, but rejected, the possibility of glacier thrusting apparently because deformed bedrock (shale) is not present. However, further investigation of Brandon Hills has shown that they are, in fact, ice-shoved ridges consisting entirely of drift (Aber 1988b). Brandon Hills occupy a rectangular area roughly 10 km eastwest and 4 km north-south (fig. 5-34). The hills exceed 590 m (1600 feet) elevation, up to 100 m above the Little Souris River valley immediately to the north. Brandon Hills include three distinct morphologic types: composite ridges, esker ridges, and kame-and-kettle moraine (fig. 5-35). Closely spaced, subparallel, composite ridges make up the northwestern, central and eastern portions of Brandon Hills. The ridges resemble a giant fishhook in overall plan view. The composite ridges are covered by a thin veneer of till. A high esker ridge meanders over the eastern end of Brandon Hills. This ridge extends with a couple of interruptions to the southwest and then northwest to form a large semicircular loop around the eastern half of Brandon Hills. The esker
Composite ridges
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Figure 5-35. Aerial photograph of Brandon Hills vicinity, Manitoba. Composite-ridges are covered by forest; esker ridges are grass covered. Location of site 6 is indicated (fig. 5-36). Aerial photograph A24519-I 79 (1976). Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada.
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Figure 5-36. Measured section from pit wall transverse to ridge axis at site 6, as it appeared in 1986. Scale in m; position of major fault zone shown by arrow (<).
Aber and Ber
82 ridges are composed predominately of sand and gravel and rest on a continuous substratum of till (Welsted and Young 1980). The southwestern flank of Brandon Hills consists of low kames and lake-filled kettles, such as Lake Clementi (fig. 5-34). The internal structure of composite ridges in Brandon Hills is fairly simple at most sites, consisting of 2-3 m of stratified sand and gravel overlain by 1-2 m of sandy till. The till is banded and is composed mainly of material eroded from subjacent stratified drift. Deformed stratified drift is visible in the largest and deepest exposure (fig. 5-34, site 6). Site 6 is a large gravel pit cut into a small southeast-trending ridge. Pit walls reveal two dislocated stratified drift blocks, each 15 m thick, which are faulted together, tilted southwestward, truncated and capped by 1-2 m of discordant till (fig. 5-36). Both stratified blocks have a similar sequence; the lower portions of each consist of sandy pebble gravel with scattered cobbles, passing conformably upward into trough cross-bedded, medium to coarse pebbly sand. Along the fault separating the upper and lower stratified drift blocks, trough-bedded sand of the lower block is either cleanly truncated or else sheared into a zone of foliated sand up to 0.5 m thick. This sand forms small apophyses which extend into the gravel above, and small thrusts and kink folds are present within this zone. Near the southwestem end of the section, a complex of high-angle to near-vertical faults cuts through sand of the upper unit, and the till/sand contact is offset here by small normal faults. The structural data are in complete agreement; faults all strike subparallel with the eastsoutheast trending ridge.
tongue pushed against the eastern end of Brandon Hills, causing ridges there to curve southward. The ice-shoved ridges were then overridden, and a veneer of till deposited on the ridges. When ice advance ceased, a subglacial meltwater tunnel system developed over the eastern end of the hill, in which an esker system was deposited, and kames accumulated among stagnant ice blocks on the southwestern side of the hill. Ages given for the Marchand advance range from >15,000 years BP (Klassen 1975) to about 11,200 years BP (Fenton, Moran et al. 1983). The discrepancy in ages is based on acceptance of selected radiocarbon dates. Teller, Moran and Clayton (1980) reviewed the complications of using radiocarbon dates from the Canadian Plains. The compilation of Laurentide ice-margin positions by Dyke and Prest (1987b) indicates Brandon Hills are between 13,000 and 12,000 years old. In any case, Brandon Hills are an excellent example of the typical morphology and intemal structures of composite ridges developed wholly within unconsolidated glaciogenic strata.
IDAUPHJN ACTIVE GLACIER: MARGIN D!RECTION :OF
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The Lostwood (late Wisconsin) Glaciation completely covered Manitoba and reached far southward into the United States (Fenton 1984). Deglaciation was accomplished by stagnation of large areas of marginal ice alternating with minor readvances or stillstands of active ice. The Assiniboine Sublobe was responsible for the final readvance--Marchand phase--into Tiger Hills (fig. 5-37). This sublobe consisted of two ice tongues; one advanced directly over what is now the city of Brandon and the other moved into Tiger Hills. A re-entrant between the two ice tongues was developed near the eastern end of Brandon Hills. The creation of Brandon Hills is most easily explained by a single ice advance, which pushed up ridges consisting of displaced blocks of older stratified drift. This drift may have been deposited when ice stagnated over the Tiger Hills upland prior to the Marchand phase. Western and central ridges of Brandon Hills were shoved up along the southern flank of the Brandon ice tongue. Meanwhile, the Tiger Hills ice
NORTHERN ICE
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Figure 5-37. Position of active glacier margin during last ice advance into Tiger Hills upland. Note two ice tongues comprising Assiniboine Sublobe; position of Brandon Hills shown by asterisk. Adapted from McGinn and Giles (1987).
Chapter 6 Cupola Hills and Drumlins Introduction Many conspicuous hills, both small and large, have the general characteristics of ice-thrust masses, but lack the hillhole relationship or the typical ridged morphology. Bluemle and Clayton (1984) placed such landforms from North Dakota into the category of irregular hills, that Clayton, Moran and Bluemle (1980, p. 46) defined as, an irregular jumble ofhills with no obvious transverse ridges and no obvious source depression. The glaciotectonic origin of such hills can be proven only with evidence of subsurface deformation of bedrock or drift, although such hills may be suspected from other evidence. Ice-shoved hills of irregular shape are probably much more common than heretofore recognized. Perhaps the most typical form of these irregular hills is the type that Smed (1962) first called cupola hill (kuppelbakke in Danish). Cupola hills have an internal structure similar to composite ridges; however, unlike composite ridges their morphology was modified substantially by the action of overriding ice. Cupola hills possess three basic attributes (Smed 1962): ° Internal structure - deformed glacial and interglacial deposits, plus detached floes of older strata or bedrock. • External form- long, even hill slopes with overall domelike morphology, unlike marginal moraines; varying from near circular to elongated ovals in form; 1-15 km maximum length, most 7-10 km long, from about 20 m to > 100 m high. • Discordant till - overridden by ice which truncated deformed structures and laid down a basal till layer. The internal structure of cupola hills is due to ice shoving of pre-existing strata during glacier advance; whereas, the external form reflects the smoothing effects of later subglacial erosion and deposition. Where subglacial modification is slight, a subdued ridge morphology may be preserved. With increasing subglacial erosion and deposition, all traces of the ridges disappear and are replaced by a smoothed cupola form. The long axis of a cupola hill may be either parallel or perpendicular to ice movement. Still further subglacial molding may create streamlined, drumlin-shaped hills (fig. 6-1). Eventually, the morphologic expression of earlier, ice-
shoved hills may be completely altered or removed by prolonged glaciation. Thus, a complete transition from composite ridges, to cupola hills, to streamlined hills and drumlins, and finally to plains is possible. Cupola hills are common in regions having soft substratum that was subjected to multiple glacial advances from different directions.
Drumlins are remarkable glacial landforms that have attracted much geological attention and generated diverse theories for their genesis. Explanations for the creation of drumlins fall into the three general categories of glacial geomorphology: deposition, erosion, and deformation (see fig. 1-5). Among the various theories for drumlin genesis, Hart (1995a, 1997) emphasized the glaciotectonic nature of many drumlins and noted a continuum of depositional, deformational, and erosional processes. Many drumlins, in fact, have cores of disturbed glacial and preglacial strata, for example in Poland (Piotrowski and Wysota 2001) and Latvia (Zelcs 2001). The following case examples include substantial dislocations of Cretaceous bedrock in cupola hills on Martha's Vineyard, Massachusetts and Hvideklint, island of Men, Denmark. Elblgg Upland, Poland is a huge glaciotectonic massif comprised entirely of unconsolidated Pleistocene sediment. Drumlins of glaciotectonic origin are described from Ristinge Klint, Denmark and the Saadj~irve drumlin field, Estonia.
Aquinnah, Martha's Vineyard, Massachusetts Martha's Vineyard is one of several offshore New England islands built mostly of two prominent late Wisconsin endmoraine systems (fig. 6-2). These moraines are largely glaciotectonic features formed by advancing ice fronts (Oldale and O'Hara 1984), which piled up glacial, interglacial, and preglacial strata along with thrust and folded masses of Tertiary and Cretaceous bedrock. Aquinnah Cliff, at the western tip of Martha's Vineyard, has long been famous for its exposures of deformed bedrock and drift. It was known formerly as Gay Head for its multi-colored display of upper Cretaceous strata, including red, orange, yellow, brown, white, gray, and black layers. The cliff reaches 45 m in height and extends ~1 km north-south. It cuts through the Aquinnah cupola hill of about 10 km 2 (fig. 6-3). Close structural and scenic similarities between Aquinnah Cliff and M0ns Klint have been pointed out often (fig. 6-4). Upthrust Cretaceous and Tertiary floes of Aquinnah Cliff represent the glacially modified edge of the Atlantic Coastal
84
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Figure 6-1. Topographic map of cupola-hills and drumlins in central North Dakota. Ice movement from NNE. Elevations in feet; contour interval = lO feet (~3 m). Adapted from Jessie, North Dakota, 7.5-minute quadrangle, U.S. Geological Survey.
Plain. Cape Cod and all the offshore islands are underlain by similar sedimentary strata of the Coastal Plain. This bedrock normally lies some depth below sea level and dips gently seaward. All surface exposures of Cretaceous and Tertiary materials on Martha's Vineyard and other islands consist of deformed floes or thrust scales. Topographic troughs in the crystalline basement surface of the mainland channeled ice flow resulting in more-or-less parallel ice lobes at the glacial margin. The late Wisconsin ice margin of southern New England was grounded on land, and the lobes apparently oscillated at times, but not simultaneously (Sirkin 1976, Oldale and O'Hara 1984). Martha's Vineyard occupies the interlobate position between the Narragansett-Buzzards Bay and Cape Cod Bay lobes (fig. 6-2). The Vineyard moraine is part of the moraine system that extends eastward to Nantucket Island and westward to
Block Island and Montauk Point on Long Island, and an extensive outwash plain lies immediately south of the moraine on Martha's Vineyard. This terminal moraine marks the outermost advance by late Wisconsin glaciation on the Atlantic continental shelf (Ridge 2004). Thrusting of the Vineyard moraine took place near the end of glaciation, as considerable material is involved in the deformations, but only a thin discordant till was deposited over the top of the truncated structures. The traditional sequence for glacial and interglacial stratigraphy on Martha's Vineyard was developed in line with four major Pleistocene glaciations of the mid-continent region. Among the several glacial formations, the Montauk Till is the most prominent and is involved in deformations on Martha's Vineyard. Believing it to be early Illinoian in age, Kaye (1964a) assigned the main ice thrusting to a late
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Figure 6-2. Map of southern New England showing late Wisconsin end moraines. Box on western Martha's Vineyard indicates location of Fig. 6-3. Inset map shows edge of Coastal Plain, approximate limit of late Wisconsin glaciation, and positions of ice lobes: N = Narragansett-Buzzards Bay lobe and C = Cape Cod Bay lobe. Moraines from Schafer and Hartshorn (1965) and Sirkin (1980). Illinoian glaciation. Ages of the Montauk Till and correlative tills in southern New England are much contested. The Montauk Till is now generally regarded as early Wisconsin (e.g. Schafer and Hartshorn 1965; Stone and Borns 1986, chart 1). Another well-supported point of view, however, is Illinoian for age of the Moutauk Till (Ridge 2004). In any case, it is safe to say that thrusting of the terminal moraine system, including the Aquinnah cupola hill, is late Wisconsin (Cadwell and Muller 2004), as no glaciation (till) intervenes between that of the Montauk Till and the late Wisconsin till. Glacial advance to the terminal moraine is dated in the range 24,000 to 20,000 radiocarbon years ago (~28-24 ka cal. age; Ridge 2004). Different interpretations of the structure and stratigraphy of deformed bedrock and drift at Aquinnah Cliff have generated many geological controversies over the years. In the 18th and 19th centuries, tectonic, landslide or volcanic activities were favored explanations for Aquinnah Cliff, and some geologists promoted the concept of a post-Miocene "Vineyard orogeny" (Shaler 1888, 1898). By the late 1800s, however, most geologists had accepted a glaciotectonic origin for the structure (Merrill 1886b; Hollick 1894; Woodworth 1897; Upham 1899), in which limbs of overturned folds and thrust faults dip northeastward (fig. 6-5).
Upper Cretaceous, Tertiary, and Pleistocene formations make up the disturbed preglacial strata of Martha's Vineyard. Upper Cretaceous strata comprise the largest volume of material exposed in Aquinnah Cliff. There strata are dominantly clastic sediments weakly indurated or unconsolidated in character, including clay, silt, sand, and gravel. Especially prominent are white kaolinitic intervals, black lignite layers, and red clay beds. These strata have been used since prehistoric time for pigments and supported a ceramic brick industry in the 19th century. In addition, two other preglacial formations outcrop in the cliffmMiocene glauconitic sand and lower Pleistocene Aquinnah conglomerate. Where undisturbed, Tertiary strata are normally encountered about 20-25 m below sea level. The Tertiary/Cretaceous boundary lies about 70 m below sea level (Hall, Poppe and Ferrebee 1980). The most detailed sketch of Aquinnah Cliff, made by Kaye in 1959 (fig. 6-6), shows a series of imbricated scales, each 20-30 m thick, dipping toward the northeast. This general scheme is interrupted in places by complex folds or downdropped blocks. Although many structural details have changed as a consequence of rapid cliff erosion, the basic elements described by Kaye are still visible today. Beginning at the northern end, half a dozen large scales culminate
Aber and Ber
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Figure 6-3. Topographic map of western Martha's Vineyard, Massachusetts. Asterisk indicates location of Aquinnah Cliff at western end of the island. Central portion of Aquinnah is >150 feet (45 m) elevation. High ridges of Chilmark and West Tisbury region generally exceed 200feet (60 m) with highest points above 300feet (90 m) elevation. Elevations and soundings in feet; map adapted from Gamble, Gamble and Allaire (1999).
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Figure 6-4. Photograph of cliff section, Aquinnah, Martha's Vineyard, Massuchusetts. View northwestward. Lighthouse slide is the low portion in foreground; Red Spur, Lignite Valley and the Great White Wall appear in the cliff beyond (see fig. 6-6). Photo by J.S. Aber (2005).
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southward with the Great White Wall. These scales are comprised mainly of Cretaceous strata in their lower portions, and strongly folded Tertiary and Pleistocene materials are present toward the tops of some. The Great White Wall is faulted against a downdropped block of nearly horizontal drift resting on Tertiary bedrock• at the base of the cliff. Immediately south of the down-dropped block, there is a large mass of steeply tilted and thrust Cretaceous strata, in which lignite is conspicuous (fig. 6-7). Continuing southward, nearly horizontal overthrusts are visible at Lighthouse Slide, followed by a complexly deformed zone in which the Quadrilateral Fold is developed. The southernmost portion of the cliff (south of Devil's Den) consists of low-angle overthrusts involving Cretaceous strata and drift. Overall, the consistency in structural development suggests that a single northeasterly ice advance thrust up scales and subsequently overrode the Aquinnah cupola hill.
The cupola hill of Aquinnah is oval in shape, roughly 5 km long and 2 km wide, extending northwest-southeast (fig. 63). The cupola hill is not ridged, but has instead uniformly sloping sides, on which there are many small, irregular knobs and depressions. Since abandonment of agriculture, however, most of the landscape has overgrown in scrubby oak-pine forest during the past half century (Foster and Motzkin 1999), so little of its geomorphology can be seen on the ground nowadays. Maximum elevation reaches nearly 60 m at the center of the cupola hill in the town of Aquinnah, some 85 m above the seafloor of Vineyard Sound to the north. To the northeast of Aquinnah, elevations on the main portion of Martha's Vineyard exceed 250 feet (75 m) in the Chilmark and West Tisbury vicinity (fig. 6-3). This area is marked by northeast-trending ridges composed of Montauk Till and displaced bedrock. Cretaceous strata striking northeast are uplifted at least 160 m above their normal undisturbed
Aber and Ber
88
Figure 6-7. Closeup photographs of deformed strata in Aquinnah Cliff. A-faulted lignite (black) and kaolinitic strata (white) in the Lignite Valley section of the cliff. Field of view is ~10 feet (3 m) across. B - red and gray clay breccia (below) and contorted lignite (above) in the Red Spur section of the cliff. Scale card is 8.5 x 15 cm. Photos by J.S. Aber (2005). position. These scales were apparently thrust up from the northwest resulting in conspicuous composite ridges (fig. 68). The Chilmark-West Tisbury ridges, along with Aquinnah cupola hill and other smaller cupola hills at Squibnocket and Nashaquitsa, outline an ice tongue, which advanced from the Menemsha Bight southward through the depression of Menemsha and Squibnocket Ponds. This same advance eventually overran the cupola hills, but may not have covered the highest Chilmark-West Tisbury ridges, which stood above the ice as nunataks (Kaye 1964b). The Menemsha ice tongue was an offshoot of the Narragansett-Buzzards Bay ice lobe. Comparison of the Vineyard moraine with other coastal moraines of Massachusetts, Rhode Island, and New York indicates that most of these so-called end moraines had a similar glaciotectonic origin (Oldale and O'Hara 1984). The moraine systems actually consist of various composite ridges and cupola hills created by ice pushing. The New England coastal end moraines were formed by ice advance against the Coastal Plain cuesta and against ice-contact stratified drift. Floes of bedrock and drift were sheared off and transported forward and upward beyond the ice front. Structural uplift of at least 160+ m occurred on Martha's Vineyard. This
thrusting was facilitated by impermeable silt and clay beds, within both bedrock and glacial strata, where high pore-water pressure could develop during ice advance. Upper Cretaceous kaolinitic and red-clay zones seem to have been particularly susceptible to glaciotectonism. Gustavson and Boothroyd (1987) compared the late Wisconsin glaciation of coastal New England with the modern Malaspina Glacier of Alaska. They concluded that the Laurentide Ice Sheet of southern New England was a temperate ice mass, in which a large volume of melt water discharged from fountains or subglacial streams at the ice margin. Under such conditions, excess ground-water pressure could have developed in aquifers beneath and beyond the ice margin and led to thrusting of largely unfrozen sediment and bedrock.
Ristinge Klint, Denmark A classic example is found at Ristinge Klint, where a small cupola hill forms the main part of Ristinge peninsula on the island of Langeland, southern Denmark (fig. 6-9). Ristinge was mentioned by Smed (1962) along with cupola hills on
Figure 6-8. Kite aerial photograph looking eastward over Menemsha Bight toward composite ridges of Chilmark and West Tisbury on the horizon. The highest ridges at scene center stand more than 300 feet (90 m) above sea level. Compare with Fig. 6-3. Photo by S.W. Aber (2005).
89
Cupola hills and drumlins
from the southeast (SjCrring et al. 1982). This famous sequence is the basis for late Quaternary stratigraphy in southern Denmark (Madsen 1916; Berthelsen 1973; Ehlers 1978; SjCrring 1983b; Houmark-Nielsen 1987, 2004). It is clear the Young Baltic ice lobe moving from the southeast was responsible for thrusting at Ristinge Klint as well as subsequent overriding and smoothing of the drumlin-shaped hill. SjCrring et al. (1982) assigned the thrusting of Ristinge Klint to the initial East-Jylland phase of the Young Baltic glaciation; however, Houmark-Nielsen (1981, fig. 12) showed thrusting of Ristinge Klint as a result of the later Badthav readvance. In either case, it seems unlikely that permafrost could exist in this situation (Berthelsen 1975). Thrusting within the clay d6collement zone was likely facilitated by increased hydrostatic pressure as a result of glacier-ice loading.
Figure 6-9. Topographic map of Ristinge peninsula with Ristinge Klint cutting across the southern side of the drumlin. Elevations in meters; contour interval = 5 m. Adapted from map sheet 1311 I RudkObing, Geodcetisk Institut, Denmark.
Hvideklint, Mon, Denmark The hills north of the village of Hjelm, including Hvideklint, on the southern coast of Men island (fig. 6-11) are an ideal example of a cupola hill built of displaced bedrock and drift. Hvideklint has been examined by several geologists (Haarsted 1956; Berthelsen, Konradi and Petersen 1977; Aber 1979; Berthelsen 1979). At Hvideklint, several upthrust floes of upper Cretaceous white chalk, covered and separated by deformed glacial strata, form a scenic cliff nearly 1 km long and up to 20 m high (fig. 6-12). Unlike the higher cliffs of eastern Men, the chalk masses at Hvideklint do not form distinct ridges that can be traced inland.
nearby ~ r ¢ and Fyn. Many of these hills have drumlin forms. High points of each are toward the southeastern ends with the hills tapering northwestward. These hills are the largest drumlins in Denmark, but their internal structure is hardly typical of drumlins. Smed (1962) concluded that ice pressure from the southeast had dislocated and streamlined these cupola hills. Ristinge Klint displays a repetition of more than 30 scales, each up to 20 m thick and each containing essentially the same Quaternary sequence, imbricately stacked in a remarkably consistent structure (fig. 6-10). Saalian "shiny clay" or Eemian marine clay forms the base of each scale, indicating that these clay formations were the zones of failure in which thrust faulting was initiated. The internal stratification of each scale is little disturbed, although some foliation and drag folding is developed adjacent to thrust faults. Structural measurements along the cliff show the average direction of upthrusting toward 322 °, corresponding to ice movement
Undisturbed chalk bedrock underlies southern Men at elevations of-25 m to -50 m (Ter-Borch and Tychsen 1987). At the surface, exposures display deformed Quaternary strata and detached chalk floes. Although the chalk at Hvideklint is lithologically similar to that of MCns Klint, the Hvideklint chalk is older. The western portion of Hvideklint contains Campanian chalk, and chalk in the eastern portion is from brachiopod zone 2 of the lower
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Discorda:nt Tit[ (D) Upper D i s l o c a t e d T i[ [ ( Ud )
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Figure 6-11. Map of Men, southeastern Denmark. HV = Hvideklint, H = village of Hjelm, S = city of Stege, MK = MCns Klint. Dashed line shows position of cupola hill. Modified from Aber (1979, fig. 1).
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Figure 6-13. Composite stratigraphic column for Hvideklint. The upper portions of chalk floes are often brecciated (V symbol) or sheared together with drift forming m~lange (X symbol). Adapted from Aber (1979, fig. 4). drift overlie detached floes of chalk, in which chalk breccia and chalk-till m61ange are locally developed.
Figure 6-12. Photograph of Hvideklint showing chalk masses exposed in the cliff face. Photo by J.S. Aber (1979). Much of the cliff face later collapsed. Maastrichtian (Surlyk 1971). The Campanian chalk is the oldest chalk exposed in Denmark. Although the chalk floes appear to rest in more-or-less horizontal positions, they are in fact cut by numerous faults and are brecciated in many places. Three tills along with associated stratified drift are exposed at Hvideklint (fig. 6-13). A discordant till overlies upper and lower dislocated tills. Large bodies of deformed stratified drift are found at or near the base of the upper dislocated till. The lower dislocated till also includes some extremely contorted lenses of stratified drift. The till and stratified
Pebble counts reveal that both the discordant and upper dislocated tills are generally characterized by a large amount of local material and about equal contents of crystalline and Paleozoic types typical of northeastern derivation. However, the lower dislocated till has noticeably less chalk and more Paleozoic limestone indicative of a Baltic source. The presence of reworked Eemian foraminifera in the discordant and upper dislocated tills indicates Weichselian age for those tills (Aber 1979). The cupola hill north of Hjelm has an oval outline and contains many irregular small hillocks and closed depressions (fig. 6-14). It rises to a maximum elevation of 44 m at BavnehCj. Only one subdued ridge, Glinsebanke, is present. Chalk is exposed at several places on Glinsebanke, so it is probably the upturned end of a chalk floe. There are otherwise no obvious linear trends in the cupola-hill morphology; land slopes are gentle to moderate throughout the hill (fig. 6-15). Hvideklint displays a progressive increase in deformation from a minimum at the northeastern end to a maximum toward the southwest. Thus, the cliff is conveniently divided into four portions (fig. 6-16)"
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Figure 6-15. Aerial photograph of Hjelm (lower left) and Hvideklint vicinity, southern MOn. Note rolling topography, which is largely cultivated, and lack of ridge morphology. Chalk is present at the surface in the northwestern corner of scene. Reprocessed from Geodcetisk Institut, Denmark, 1974 (A. 29/88).
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Figure 6-16. Measured profile of Hvideklint section as it appeared in 1979. Base of section is at beach level. Note 2X vertical exaggeration slightly distorts geometry. Scale in m; based on Aber (1979, fig. 3). • Central chalk floe (640-855 m) - The largest single chalk body at Hvideklint forms a broad anticline in which a conjugate thrust system is developed. The chalk is thrust over isoclinally folded and sheared stratified drift at its southwestern end.
• Western chalk floes (855-1080 m) - Highly disturbed chalk and drift display extreme stretching and thinning toward the southwest. A body of lower dislocated till is almost entirely enclosed by deformed chalk near the eastern end
Figure 6-17. Lower dislocated till (dark gray) contains bodies of contorted sand and gravel, and is surrounded by chalk floes. Hvideklint section meter 860-880 (see fig. 6-16). Scale pole marked in 20-cm intervals; photo by J.S. Aber (1979).
Figure 6-18. Extreme stretching, folding and intermingling of chalk, till, and stratified drift forming a chalk-till m61ange. Near western end of Hvideklint section, meter 1040-1060 (see fig. 6-16). Scale pole marked in 20-cm intervals; photo by J.S. Aber (1979).
93
Cupola hills and drumlins
Hvideklint was apparently little altered by subsequent Young Baltic advances. The lower dislocated till is problematic; it could be the Ristinge Klint Till or it could relate to a still older Saalian advance.
../.I-"............................:::L:,::::::::: ............... .......:.-............. x //
i
Elbl~g Upland, Poland
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]/
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2
Figure 6-19. Hvideklint structural data plotted on an equalarea stereonet. Solid dots = fold axes; arcs = faults. Note grouping of data in two clusters.
Elbl~g Upland is an exceptionally large glaciotectonic massif located near Gdansk Bay, a large and deep embayment of the Baltic Sea in northern Poland (fig. 6-20). The glaciotectonic origin of Elbl~g Upland was first suggested by Pazdro (1958). He noted complicated and variable hydrogeological conditions, as well as anomalous thicknesses and multiple repetitions of clay and till units, resulting from strong deformation. A great many rafts of Tertiary strata along with other glaciotectonic structures were mapped in the Gdansk region by Mojski and Sylwestrzak (1978) and Mojski (1979). On this basis, Ruszczyfiska-Szenajch (1985, 1993)concluded that glaciotectonism was an important process for the formation of Gdansk Bay. The geology and stratigraphy of Elbl~g Upland were investigated as part of regional geological mapping by
of this portion (fig. 6-17). Penetrative deformation has caused intermingling of chalk and drift in isoclinal folds and shear bands (fig. 6-18). The orientations of glaciotectonic structures are consistent along the cliff section. Fold axes are mostly near horizontal, and all trend either southeast or northwest. Thrust and normal faults also strike northwest-southeast (fig. 6-19). These structural data correspond to northeasterly ice movement at Hvideklint. The orientation data are clustered in two groupings: (1) 100to 140 ° and (2) 160to 165 °. This suggests actually two episodes of structural disturbance were associated with ice advances from about 30 ° and 70 ° . The chalk floes were presumably derived from somewhere in the Fakse Bugt vicinity north of Men. Southern Men was subjected to the same late Weichselian ice advances that affected eastern Men (chap. 5) and Ristinge Klint. The discordant and upper dislocated tills relate to the Main Weichselian advances which came initially from the northeast and then shifted to an east-northeasterly direction. The upper dislocated till probably was deposited by the initial advance from the northeast, and so correlates with the Mid Danish Till. This advance transported the chalk megablocks and constructed the ice-shoved hill. The discordant till was then deposited by the Storebaelt readvance from the east-northeast, and so corresponds to the North Sja~lland Till. This advance caused some additional structural disturbance and smoothing of the cupola hill.
X
RUSSIA
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.~
t:l k ~n
..............
MA×iMUM
E×TENT
MAZURI
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60
80
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Figure 6-20. Location map for glaciotectonic features of the Gdansk Bay (Zatoka Gdahska) region in northern Poland, showing bathymetric contours on the Baltic Sea floor. 1 = zones of numerous glacial rafts and other glaciotectonic disturbances (from Mojski and Sylwestrzak 1978; Mojski 1979); 2 = Elblqg Upland. Position of Eemian Sea limit f r o m Makowska (1979b). A d a p t e d from Aber and Ruszczytiska-Szenajch (1997, fig. 2).
94
Aber and Ber
Makowska (1978, 1979a, 1988, 1992). She demonstrated that Elbl~g Upland is comprised almost entirely of Pleistocene sediments, most of which are interstadial and/or interglacial in origin and were deposited originally in deltaic, lacustrine, fluvial, and/or marine environments. She also demonstrated the upland has a nearly continuous cover of the youngest till, from the Baltic phase of North Polish (recessional late Vistulian) Glaciation (Makowska 1995).
and southwest, till plain of the Bauda and Pasigka lowland to the east and southeast, and coastal lagoon (Zalew Wi~lany) to the north. The estimated volume of the upland above sea level is approximately 40 km s. These dimensions place Elbl~g Upland among the largest glaciotectonic landforms in the world (Aber and Ruszczyfiska-Szenajch 1997). Within Elbl~g Upland, two morphologic patterns are well defined: 1) eastern arcuate belt, concave toward the northeast, and 2) western arcuate belt, concave to the west (fig. 6-22). The two morphologic sets converge near the central, highest portion of the upland. Similar morphologic trends are not evident for the eastern and southern portions of the upland,
Elbl~g Upland is approximately equidimensional, covers an area of about 390 km 2, and rises to nearly 200 m above the surrounding lowlands (fig. 6-21). Adjacent lowlands include the delta plain of the Wisla/Nogat/Elbl~g rivers to the west
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Figure 6-21. Topographic map of Elblqg and surroundings, northern Poland. Maximum elevation reaches nearly 200 m at Milejewo near map center. Elevations in meters; contour interval = 20 m. Approximate boundary of upland marked by black outline. Adapted from map sheet N-34-XIV Elblqg, Sztab Generalny Wojska Polskiego, Warszawa, Poland.
95
Cupola hills and drumlins
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which have smoother surfaces. In addition, drumlins aligned NNW-SSE are situated in the vicinity of Pr6chnik. Surface exposures of internal structures are scarce. Among the best formerly was a large clay mine at Kadyny on the
0
5
!0
Figure 6-22. Surficial features of Elblqg Upland and surroundings. Two sets of arcuate trends are portrayed in surface morphology of the upland, one concave toward the northeast, the other concave to the west. These two sets merge near the center of the upland at Lake Stare. Locations of exposures indicated by solid dots; bore holes at Rakowo (R) and Bystra (B) shown by open circles; position of geological section shown by line AB-C (see fig. 6-25). Adapted from Aber and R uszczyhska-Szenajch (1997, fig. 4).
15
northern margin of the upland (fig. 6-23). The mine exploited a mass of Elbl~g Clay, which Makowska (1995) recognized as strongly deformed and displaced by ice pushing. Stratigraphic position of the marine Elbl~g Clay is problematic. It is considered to be post-Eemian; it probably represents the Krastudy interstade of early/mid Vistulian age (Makowska 1995; Mojski 1995). Elbl~g Clay strata are tilted toward the northeast; dip varies from less than 30 ° to nearly vertical. The clay exhibits much internal deformation, including fractures, faults, slickensides and brecciation. Till and stratified glacial sediment are deformed along with the Elbl~g Clay in repetitive sections. The lower portion of the exposure has consistent structural orientations that indicate ice pushing from the northeast (fig. 6-24). Upper portions of the clay mine display more complex deformations, and the upper structures have more variable orientations. These structural observations are consistent with those from other sites, which suggest multiple phases of deformation from different directions.
Figure 6-23. Eastern portion of Kadyny clay mine, as it appeared in 1993. Dark Elblqg Clay standing in vertical position. Top of exposure is approximately 70-75 m in elevation. See Fig. 2-2 for closeup view of brecciated clay. Photo by J.S. Aber (1993).
Test drillings throughout the Elbl~g Upland revealed several anomalous conditions (fig. 6-25). These features, considered together, point to strong glaciotectonic deformation of the sediments of Elbl~g Upland (Aber and Ruszczyfiska-Szenajch 1997).
Aber and Ber
96 N
• Quaternary sediments of Elbl~g Upland are almost 100 m thicker than equivalent Quaternary strata that build the adjacent Bauda-Past~ka lowland to the east. • Sediments of analogous originmglacial, fluvial, lacustrine or m a r i n e ~ o c c u r on quite different levels within comparatively short distances. • Distinctive sediments, such as the "red clay," are repeated in many places at quite different levels. • Rafts of Tertiary strata within the glacial sequence are found up to 130 m above sea level (at Bystra), and till is encountered more than 100 m below sea level. Makowska (1979a, 1988) earlier interpreted the origin of Elbl~g Upland in terms of repeated Pleistocene erosion and deposition, which resulted in great variations in thickness and position of strata, as well as many repetitions of sedimentary units. More recently she modified her interpretation to include glaciotectonic disturbances of the upper (but not lower) portions of Elbl~g Upland (Makowska 1995). She stated that displaced masses of Elbl~g Clay may have been partly overthrust from the Baltic Sea region. She further believed deformation took place by folding in a plastic state and, after freezing, pushing in the form of scales or blocks. Aber and Ruszczyriska-Szenajch (1997), in contrast, believed the entire mass of Elbl~g U p l a n d m u p p e r and lower portions--has a glaciotectonic genesis. They concluded that major deformation of Elbl~g Upland resulted from advances by two ice lobes, which pressed locally from the northeast and west, corresponding to the two main morphologic trends within the upland. An ice lobe in what is now the Wisia delta vicinity could account for ice squeezing of the western portion of Elbl~g Upland. Glaciotectonic deformation of the
PLANAR FEATURES
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Figure 6-24. Equal-area stereographic plot of structural measurements from the lower portion of Kadyny. Fold axis is transverse to ice movement; all other structures are parallel to direction of deformation. These data reflect ice pushing from the northeast. Adapted from Aber and RuszczyhskaSzenajch (1997, fig. 6).
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Figure 6-25. Geological section through the central part of Elblqg Upland and the lowland region to the east. Topographic profile is generalized. CrCretaceous, T- Tertiary (undisturbed), (T) - Tertiary rafts, r - red clay, P Pommeranian/Gardno glacial sediment. Arrows indicate assumed directions of glaciotectonic displacements. Adapted from Aber and Ruszczyhska-Szenajch
(1997, fig. 12).
Cupola hills and drumlins
97
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Figure 6-26. Saadjiirve drumlin field in east-central Estonia. Bedrock: 0 Ordovician limestone and dolostone, S - Silurian limestone and dolostone, and D - Devonian clastic strata. IM - position of section in lnglimiigi hill near southeastern end of drumlin field (see fig. 6-29). Lake Saadjiirv indicated by asterisk (*). Adapted from Rattas and Piotrowski (2003, fig. 1). Reproduced from Boreas by permssion of Taylor & Francis AS.
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upland from the northeast could have formed in like manner by an ice lobe in the Bauda-Past~ka lowland. These lobes may have branched from the main ice mass filling Gdansk Bay. It is uncertain whether the two ice lobes advanced at the same time or were separate events. However, available evidence suggests the upland formed as an interlobate feature that was squeezed nearly simultaneously from two directions, during the main phase of large-scale deformation. The process of ice squeezing resulted in the main constructional form of Elblgg Upland, and ice pressing must have created at the same time the presumed source depressions on both sides of the upland. Whether either of these ice advances
o
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overrode the upland is not clear. The final ice advance from the NNW did override the upland and caused formation of several dozen drumlins with disturbed cores in the Pr6chnik and Majewo vicinity (Wysota, Karasiewica and Sokotowski 2001). This final advance also deposited discordant diamicton and smoothed the upland surface. These events are clearly late Vistulian; however, the exact sequence of ice advances, sediment deposition and deformation is not established with certainty. The final glacier transgression is most likely connected with the Pommeranian phase, a major readvance of ice from the Baltic basin, now dated between approximately 15,500 and 14,600 years ago (Rinterknecht et al. 2005, 2006).
98 The position and composition of Elbl~g Upland are related to its geological situation on the margin of Gdansk Bay (fig. 6-20). Unconsolidated marine, lacustrine and fluvial sediments underlie the coastal estuary region. These sediments have highly variable textures (clay, silt, sand, gravel), porosities, and hydraulic conductivities. They are bounded to the south by Tertiary and Cretaceous strata with generally lower porosities and hydraulic conductivities. Major glaciotectonic deformation at Elbl~g likely took place when ice advanced into a lake located in Gdansk Bay. Such lakes most probably existed in the Gdansk embayment during several phases of North Polish glaciation.
Aber and Ber
Under these conditions, the ice sheet became much more mobile in the southern Baltic, and rapid ice-lobe advances-possibly surgesmprobably took place within thawed-bed zones leading toward proglacial lakes. During such advances, the unconsolidated and unfrozen sediments of the coastal zone developed high pore-water pressures within confined aquifers. The combination of rapid glacier loading and high ground-water pressure led to glaciotectonic dislocations at Elbl~g Upland and elsewhere around the margins of Gdansk Bay. Piotrowski (1993) interpreted a similar combination of factors for glaciotectonism in the Baltic region of western Germany.
Figure 6-27. Satellite image of the Saadjiirve drumlin field vicinity, east-central Estonia. Drumlins trend across the scene from northwest to southeast. Location of Lake Saadjiirv indicated by asterisk (*). Based on Landsat TM band 5 (midinfrared), 1999.
Cupola hills and drumlins
99
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Figure 6-28. Kite aerial photograph over southern portion of Saadjiirve drumlin field. View toward northwest along axis of drumlin that separates Lake Saadjiirv from Lake Soitsjiirv (upper right corner). See Fig. 6-26for location of Lake Saadjiirv. Photo by S.W. Aber (2000).
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Overall the drumlin field has a funnel shape, more than 25 km wide at the proximal end and converging to less than 5 km width toward the southeastern, distal end (fig. 6-27). Glacier movement was generally from northwest to southeast, as shown by drumlin orientation, which implies extending flow conditions as the ice passed from the Pandivere Upland toward the Peipsi Depression. The drumlins were created during an ice readvance between the Otepii and Pandivere phases of Late Weichselian glaciation, about 13,000 to 12,500 years ago (Raukas et al. 2004; Rinterknecht et al. 2006). The Saadjirve drumlin field displays two main portions: northwestern proximal half and southeastern distal part. Drumlins in the proximal portion are broadly spaced and relatively large, averaging 12 km in length. Smaller, closely spaced drumlins are typical of the distal part; their average length is only 1.75 km (fig. 6-28). The proximal portion of the drumlin field corresponds with underlying Ordovician and Silurian limestone and dolostone substratum (fig. 6-26). This bedrock is cavernous and has high permeability. In contrast, the distal part of the drumlin field is situated above
300' NW
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Saadj~irve drumlin field, Estonia
The Saadjirve drumlin field comprises approximately 120 individual drumlins distributed over an area of roughly 1200 kin 2 in east-central Estonia (fig. 6-26). The drumlins occupy a plateau between the Pandivere Upland to the northwest and the Peipsi Depression to the east. The Saadjirve drumlin field was investigated in detail by Rattas, who described various glaciotectonic structures and elaborated conditions under which drumlin formation took place. The following description is based primarily on her work (Rattas and Kalm 2001, 2004; Rattas and Piotrowski 2003).
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ero, Figure 6-29. Stratigraphic profile at Inglimiigi hill at southern end of Saadjiirve drumlin field. Massive sandy diamicton (Dmm) is the drumlin-forming upper till. It overlies stratified sediments that rest on a base of older till (Dmg). A slab of the older till is thrust into the stratified sediment (*), and stratified sediment is strongly sheared into a stratified diamicton (Dms) and reworked into the basal portion of the upper till. For location see Fig. 6-26; adapted from Rattas and Kalm (2001, fig. 2).
1O0
Aber and Ber
%\
Figure 6-30. Photomosaic of exposure at Inglimiigi hill at southern end of Saadjiirve drumlin field. Note numerous faults, dislocated blocks, and shear zones. For symbols see Fig. 6-29; image courtesy of M. Rattas. poorly consolidated Devonian sandstone, siltstone, clay and marl strata of much lower permeability.
of stratified sediment into the basal portion of the drumlinforming till (fig. 6-30).
Till is the primary constituent of most drumlins, and the drumlin morphology is a consequence of the youngest, surficial till cover. This drumlin-forming till varies in composition in relation to underlying bedrock. It is gray and clayey in the region of Ordovician and Silurian bedrock; whereas it becomes red and sandy in the area of Devonian substratum. Some drumlins include cores of outwash sediment resting on older till and covered by the drumlin-forming till. The latter commonly includes thrust, folded and strongly contorted bodies of outwash sand. These internal structures are demonstrated at Inglim~igi near the southern end of the drumlin field (fig. 6-29). Glaciotectonic features include a slab of older till dislocated into the overlying stratified sediment, penetative sheafing toward the top of the stratified sediment, and thrusting
Rattas and Piotrowski (2003) concluded that drumlins were created by non-uniform till deformation as the glacier advanced over a weak substratum. The contrasting morphology of drumlins in the proximal and distal portions of the field reflects differences in substratum lithology and permeability. Large, broadly spaced drumlins in the proximal portion developed where subglacial melt water could drain away effectively; small, closely spaced drumlins in the distal part formed where subglacial melt water was trapped. In the latter zone, subglacial melt water under high pressure facilitated till deformation, as shown at Inglim~igi. The funnel-shaped outline of the Saadj~irve drumlin field is consistent with accelerating glacier flow toward the southeast, which may represent the onset of local ice-stream flow conditions within the advancing ice sheet.
Chapter 7 Megablocks and Rafts Introduction
The common occurrence of flat-lying glacial rafts or megablocks throughout glaciated plains was not appreciated until fairly recently (Stalker 1976; Sauer 1978; RuszczyfiskaSzenajch 1987). These megablocks are more-or-less horizontal, slightly deformed, and are often buried under or within thick glacial strata giving little or no morphologic clue to their presence in the subsurface. They may, in fact, easily be mistaken for bedrock, if deep exposures or drilling logs are not available. In other cases, the megablocks form flat-topped buttes, small plateaus, or irregular hills, which have been mistaken also for bedrock outliers. Most megablocks are composed of poorly to moderately consolidated Mesozoic or Cenozoic sedimentary strata, but some consist of well-consolidated rocks or unconsolidated Quaternary strata. All megablocks exhibit some signs of deformationmshear zones, folds, faults, brecciation, etc.w as a result of ice pushing. This criterion distinguishes megablocks and rafts from large, undeformed erratic blocks (Aber 1985). The sources of many megablocks are unknown or poorly identified, but most were probably transported only a few km. Source depressions for megablocks often cannot be identified.
brecciated shale is intermixed with till in tongues and lenses between fault blocks. Shale forms the cores of anticlines; overlying limestone beds have broken sharply, and shale has pierced upward in the axial portions of folds (fig. 7-2). Striations on the upper surface of deformed limestone trend 155 ° , normal to fold axes and the strike of faults, all of which confirm ice movement toward the SSE. Lammerson and Dellwig (1957) interpreted deformation in a subglacial setting as a consequence of glacier pushing against a low escarpment and direct loading by the ice mass. Long-distance transportation of megablocks is documented in some cases. A group of famous rafts at Lukow in eastcentral Poland is composed of Jurassic clay derived from Lithuania >300 km to the northeast (Jahn 1950). Some of these rafts are >20 m thick, and they were probably transported in a frozen condition (Ruszczyfiska-Szenajch 1976). The Cooking Lake megablock, located near Edmonton in central Alberta, lies at the surface. It is about 10 m thick and covers an area of roughly 10 km 2. This megablock consists of Lower Cretaceous strata from the Grand Rapids Formation, which has its nearest outcrop in the Thickwood Hills >300 km to the northeast (Stalker 1976). In northeastern Germany at Bad Freienwalde, several megablocks of Rupelian (Tertiary) clay are the basis for a
A typical small m e g a b l o c k of upper P e n n s y l v a n i a n (Paleozoic) limestone is present at the surface near the glacial limit west of Topeka, Kansas (Dellwig and Baldwin 1965). The megablock consists of the Tarkio Limestone Member of the Zeandale Formation. Its exposed dimensions are 50 m north-south by 150 m east-west with a thickness of only 1 to 2 m. The internal structures include fracturing, thrusting and rotation of limestone blocks (fig. 7-1). The underlying Willard Shale is normally about 12 m thick in the vicinity. However, below the megablock only thin (30 cm) remnants of brecciated shale and glacial debris are present, resting on a striated surface of the undeformed Elmont Limestone. Striations agree with structural features indicating ice movement from the northwest. This megablock was probably derived from nearby (<1 km away) by horizontal sliding. In south-central Iowa, Lammerson and Dellwig (1957) described disruption of flat-lying, lithified Pennsylvanian (upper Carboniferous) limestone and shale over an area of several hectares. The disrupted sequence consists of Galesburg Shale and Dennis Limestone, totaling ca. 6.5 m in thickness, resting on undisturbed Bethany Falls Limestone. Limestone and shale are fractured, folded and faulted, and
Figure 7-1. Photograph of Paleozoic limestone raft near Topeka, Kansas. Upper limestone blocks are tilted and shoved over the undeformed limestone beds in lower part of exposure. Glacial sediment is present between and below the tilted blocks. The deformed blocks are part of a large, thin mass of glacially transported bedrock. Scale pole marked in feet (1 foot = 30 cm). Photo by J.S. Aber (1989).
102
Aber and Ber
Figure 7-2. Folds in deformed Pennsylvanian limestone and shale, south-central Iowa. A - Winterset Limestone Member and B - Stark Shale Member of the Dennis Limestone; C Galesburg Shale. Notice angular fold limbs and piercing of shale upward along fold axial planes. Photo courtesy of L.F. Dellwig.
ceramics industry (fig. 7-3A). The dislocated clay sequence is 60 to 100 m thick and contains septarian concretions. Often resting on horizontal beds of Pleistocene sand, the clay rafts have been uplifted at least 50 m by glacial transportation (Hannemann 1995). Also found in the same vicinity are rafts of Oligocene "glimmersand" (mica quartz sand) that likewise rest on Pleistocene strata (fig. 7-3B) All such megablocks have one trait in common; they are remarkably thin (typically <30 m) compared to their lateral dimensions (often > 1 km2), although a few are much larger. They are thin rock slices, which could have been transported only by freezing onto the underside of a glacier. Thus during transportation, the megablock was effectively the base of the ice sheet, and glacier movement was accomplished by sliding over a d6collement within the substratum. Deposition occurred when basal melting released the megablock from the ice. At that point, continued glacier movement could have further deformed the megablock, eroded it away, or covered it with till. The fact that such megablocks are scattered over broad regions behind ice-margin positions supports a subglacial origin for many megablocks (fig. 7-4). Some megablocks may also have been initially pushed in proglacial settings (Ruszczyfiska-Szenajch 1976). In certain cases, megablocks are aligned in chains of hills decreasing in size away from their sources. The following case examples span a great size range from the huge megablock (about 1000 km 2) at Esterhazy, Saskatchewan, to medium megablocks (1-10 km2) of southern Alberta, to a chain of small rafts (<1 km 2 each) at Kvarnby, Sweden. Although diverse in their other characteristics, these examples do share one feature in common; they all consist of
Figure 7-3. A - Quarry exposure in a megablock of Rupelian (Tertiary) clay at Bad Freienwalde, northeastern Germany. The clay thrust block contains numerous internal deformations. B - nearby exposure of Oligocene "glimmersand" (upper right) resting on Pleistocene strata. Photos by J.S. Aber (1995). iiiiiiiii~i iiiiil
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Figure 7-4. Possible means of detaching megablocks: a = beheading of butte, b - plucking from lee of hill, c = scooping from depression, d = thrusting at ice front. Vertical scale exaggerated; a, b and c may occur anywhere under the ice. Adapted from Aber (1985a, fig. 3).
weakly consolidated Cretaceous strata. A final example from northeastern Estonia is well-consolidated Paleozoic limestone.
Esterhazy, Saskatchewan During the course of geological and ground-water reconnaissance mapping in eastern Saskatchewan, a truly enormous megablock of Cretaceous shale was discovered at
Megablocks and rafts
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Figure 7-5. Bedrock contour map of Esterhazy vicinity, Saskatchewan. The 1600-foot contour between Esterhazy and Rocanville defines position of megablock. Small dots show test wells used in constructing cross section (fig. 7-7). Elevations of bedrock surface in test wells shown in feet; contour interval = lOO feet (~30 m). Adapted from Christiansen (1971b) and Sauer (1978, fig. 6). Esterhazy by Christiansen (197 lb). The megablock covers an area roughly 1000 klTl2 in extent (fig. 7-5). The region is part of the Saskatchewan Plain, a drift-mantled plain of relatively low relief underlain by soft Cretaceous sedimentary bedrock. The megablock has no morphologic expression at the surface. The monotonous plain is broken by a large spillway channel, the Qu'Appelle Valley, which cuts through the middle of the megablock (fig. 7-6). The Qu'Appelle Valley is > 100 m deep and about 2 km wide. Its bottom is drift filled, and bedrock outcrops are found along its walls. The position of the Qu'Appelle Valley is in no way related to the presence of the megablock. Cross cutting of the megablock by the valley was merely a fortuitous coincidence. Bedrock in eastern Saskatchewan is generally undeformed, dipping gently toward the south (fig. 7-7). The Cretaceous formations are mostly clastic strata consisting mainly of siltstone, claystone and shale. Hard, siliceous shale of the Odanah Member, Riding Mountain Formation is more consolidated than other strata and forms the main mass of the megablock. Highly folded and faulted bedrock of the Riding Mountain Formation is found west of Hazel Cliff, east of Tantallon, and south of the Qu'Appelle Valley both in surface exposures
and in drill holes. Breccia, slickensides, and mylonite are common microstructures. This deformation is confined to bedrock above about 1500 to 1600 feet elevation (Christiansen 1971b). The 1600-foot contour may be used as a minimum outline for the megablock. The general plan of the megablock is an egg-shaped oval, roughly 38 km long and 30 km wide. Bedrock in the westcentral portion reaches 1920 feet elevation, indicating a maximum thickness of about 100 m. The megablock is, thus, at least 300 times wider and 380 times longer than it is thick. Assuming an average thickness of 60 m and an area of 1000 km 2, the megablock's volume is estimated to be 60 km 3. A test well drilled southwest of Esterhazy near the western end of the megablock intersected 2 m of till after penetrating 80 m of brecciated and mylonitic bedrock (Christiansen 1971b). Similarly disturbed bedrock rests directly on undeformed bedrock south of the Qu'Appelle Valley. The actual d6collement zone is located in Riding Mountain Formation claystone below the Odanah Member, and some claystone above the Odanah Member is also part of the megablock. The megablock is situated on the northern edge of a major buried glacial valley, the Rocanville Valley. This valley is
Aber and Ber
104
Figure 7-6. Aerial photograph of Qu'Appelle Valley, eastern Saskatchewan. Bedrock outcropping along walls of valley is part of a huge megablock. Tantallon is small town near upper left corner. Aerial photograph A21748-94 (1970). Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada. filled with up to 60 m of sand capped by till. Melt-water erosion and filling of the valley presumably took place during an earlier (pre-late Wisconsin) glaciation, and the till cap was deposited during the late Wisconsin glaciation. It appears that the Rocanville Valley truncates the southern side of the megablock, and thus, emplacement of the megablock must predate cutting of the Rocanville Valley.
overriding ice sheet, in which case the megablock became the basal layer of the ice sheet. It is highly improbable that this megablock could have been pushed in front of an advancing glacier. Subglacial sliding of permafrozen material over a thawed, clay-rich substratum seems to be the most likely explanation for displacement of the Esterhazy megablock.
Ice flow in this part of Saskatchewan was generally from the northeast; however, neither the direction of emplacement nor the source of the megablock are known. The megablock may not have moved far, perhaps less lateral displacement than its own width, in order to produce the observed structures. The only conceivable means of displacing a megablock of such huge size was by freezing onto the bottom of an
Southern Alberta
Numerous, large megablocks comprised mainly of Cretaceous strata are scattered throughout the plains of southern Alberta (fig. 7-8). The total number of such megablocks is unknown, but is undoubtedly great. Most known megablocks are partly or wholly buried beneath thick drift, and some are known only
Megablocks and rafts
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from drilling records. Where megablocks make up a substantial volume of drift-plain deposits, as in many parts of southern Alberta, the surface landform may be called a megabloek plain. Stalker (1973, 1976) was apparently the first geologist to
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106
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extend for several km, but are always quite thin (<30 m). Selected megablocks are described in the following paragraphs. Many of these megablocks are exposed in bluffs where modem valleys cross buried valleys, but megablocks are also known in other situations. - Bluffs along the eastern side of the Oldman River, about 15 km west of Lethbridge, contain a 1.5-km-long megablock about 25 m above river level and 60 m below the bluff top. Tills are present both above and below the megablock. Preglacial gravel resting on Bearpaw Formation shale is present at the base of the bluff. The megablock includes a m61ange of sandstone with clay partings, coal seams and ironstone bands. The source of this megablock is unknown; likewise its direction and distance of transportation are unknown. Kipp
L a u n d r y Hill - This megablock is exposed in bluffs along the eastern side of the Oldman River and in the coulee (ravine) near Laundry Hill entirely within the city limits of Lethbridge. The megablock contains up to 14 m of dark, bentonitic shale and white, poorly indurated sandstone. It is overlain by multiple tills, gravel, silts, and varved lake beds. At least two tills are present below the megablock as well. Megablock thickness averages about 5 m, and it was originally >1 km wide. The megablock is nearly horizontal, and in many parts it displays remarkably little internal disturbance. Its source and distance of movement are uncertain. Bend - One of the largest megablocks yet discovered in Alberta outcrops in the bluff along the eastern side of the Oldman River about 15 km northeast of Taber. It consists of interbedded Cretaceous shale, coal, and sandstone strata. Intermittent exposures extend for >3 km along the bluff, and seismic shot-hole logs indicate the megablock extends a considerable distance behind the bluff. Maximum thickness is about 25 m with average thickness being about 10 m. Stalker (1976) estimated the Driftwood Bend megablock covers at least 10 km 2. Driftwood
The continuity and intact nature of this megablock have been stressed, but the megablock may actually include several discrete blocks lying adjacent to each other at about the same level. Some blocks are nearly fiat-lying, and others are strongly deformed (fig. 7-9). Deformation is minimal toward the southern end of the section and increases northward. At least one block near the northern end forms a large recumbent fold that is stretched southward into m61ange. The megablock is underlain by some 25 m of drift including till, and is overlain by 15 m of similar material. Immediately beneath the megablock, shear planes and slickensides are developed in the till. Once again, the source and distance of travel of this megablock are not known. Wolf Island - A megablock composed of Cretaceous shale,
Figure 7-9. Megablocks of Cretaceous strata (light tone) within thick glacial strata exposed along the bluff of the Oldman River, near Taber, south-central Alberta. A coherent megablock. The upland surface is a nearly fiat glacial plain. Note people standing at the base of the megablock, which is a shear zone in the underlying glacial sediment. B - recumbent fold with core of disrupted bedrock material resting on glacial strata. Photos by J.S. Aber (1984).
coal and bentonitic sandstone is exposed for a distance of nearly 1.5 km along the northern bluff of the Oldman River about 20 km east-northeast of Taber. It is both underlain and overlain by tills and other drift in the 75-m-high section. The megablock is more-or-less horizontal, up to 13 m thick, and is locally deformed particularly near its base, where it intermingles with underlying till. The megablock also includes up to 3 m of preglacial gravel resting on the Cretaceous strata. The gravel has a Bow Valley lithology, which is not normally found in the Oldman River system. The southern limit of Bow Valley gravel is more than 50 km to the north (fig. 7-8). Catchem - This megablock is known from only two drill holes about 20 km east of Pakowki Lake. The megablock is probably no more than 1 km 2 in extent, and its source and distance of transportation are not known.
1107
Megablocks and rafts Megablocks are seemingly ubiquitous in the Alberta Plain. The source and direction of movement for many are unknown, but some did travel considerable distance. The megablocks are not related to any single glaciation or particular topographic setting. Their emplacement appears to be widely scattered in time and space. Individual megablocks were probably transported by freezing onto the bottom of an overriding ice sheet. Thin or discontinuous permafrost would have facilitated the detachment of large, thin megablocks and their incorporation onto the base of the ice sheet. Basal melting would have later allowed separation of megablocks from the ice. At that point, megablocks could have been partly disrupted, pulled apart, folded or sheared together with drift.
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Kvarnby, Skhne, Sweden Several small upper Cretaceous chalk rafts, known since the beginning of the last century (Holst 1903, 1911), are situated immediately east of Malm6 in Sk~ne, southwestern Sweden (fig. 7-10). These chalk bodies were investigated by Ringberg (1980, 1983) in connection with geological mapping in the Malm6 area, and comprehensive analysis was carried out by Ringberg, Holland and Miller (1984). The chalk rafts are aligned in a northeast-trending chain stretching 4.5 km long in a zone 700 to 800 m wide (fig. 7-11). These chalk rafts are exposed in quarries, and several other chalk bodies are present near the surface or buried along this trend (fig. 7-12). The chain of chalk bodies is situated in a drift-mantled region along a morphologic boundary. A hummocky moraine of small hills is developed to the northwest at elevations 10-20 m above sea level. To the east and southeast, the land is 1560 m above sea level and more rugged. Kvarnby is located on the southwestern margin of a deeply buried bedrock valley, the Alnarp Graben, although the position of the chalk rafts is probably not related to the presence of the buried trough. It is known from numerous test wells that Danian (lower Paleocene) limestone, 50-100 m thick, forms the bedrock surface under this entire region. Soft, white Maastrichtian chalk (upper Cretaceous) underlies the Danian limestone. Microfossils in chalk rafts at Kvarnby indicate late, although not latest, Maastrichtian age. Outcrops of similar chalk bedrock are found to the northeast within the highly faulted Fennoscandian Border Zone at Romelefisen, to the south and southwest on Rt~gen, M¢n and at Stevns, and on the intervening Baltic sea floor (fig. 7-10). The two largest chalk rafts, which reach a maximum thickness of 25 to 30 m, are exposed in quarries at Kvarnby and ,gmgdala. The two exposures display close similarities (fig. 7-13). The Kvarnby Till underlies and is dislocated between the chalk rafts. The chalk bodies are included stratigraphically
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Aber and Ber
108
increases again, reflecting a shift in ice movement direction during till deposition. The Sunnanh Till (= Malm6 Till of Ringberg 1988) is separated from the S. Sallerup Till by a thin layer of deformed sand thought to be about 13,300 years BP. This youngest till is rich in Baltic rock types, much like the Kvarnby Till. Striations and till fabrics show ice m o v e m e n t from southwesterly to westerly directions (Ringberg 1988). Presumably it was deposited by a Baltic ice lobe that advanced into the Oresund strait from the south and then spread eastward into Skhne. However, the exact nature of this socalled Low Baltic advance in southern Sweden is the subject of much debate (Lagerlund 1987; Ringberg 1988).
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The chalk rafts were emplaced by the initial Baltic advance from the south, and were subsequently modified by the main advance from the northeast. The chalk rafts were carried more than 25 km from the Baltic sea floor to Kvarnby, probably while frozen to the base of the glacier. The chalk was likely transported as one or a few larger rafts, which were pulled apart during deposition and thrust together with the Kvarnby Till forming the northeast-trending chain of rafts.
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Figure 7-11. Bedrock contour map of Evarnby vicinity showing locations of chalk rafts and borings. Position of cross section (fig. 7-]2) indicated; contours on bedrock surface in m above or below sea level. From Ringberg, Holland and Miller (1984, fig. 3).
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109
Megablocks and rafts
SinimJied, Estonia The coastal cliff of northern Estonia, known as the Baltic Klint, is an escarpment 50-60 high held up by wellconsolidated Ordovician limestone and dolostone. The base of the cliff consists of Cambrian sandstone and unconsolidated clay just above sea level. Glaciotectonic disturbance of bedrock is common along the Baltic Klint (Rattas and Kalm 2004). At Sinim~ied in the northeastern comer of Estonia, three dislocated blocks of lower Ordovician limestone core three hills in a ridge that trends approximately 5 km eastwest and rises ,-40-50 m above the surrounding flat landscape. These blocks were transported at least 2 km south of the Baltic Klint (fig. 7-15). The blocks are tilted, folded, and locally thrust along with included lenses of till, which demonstrate the glaciotectonic origin of the structures.
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The megablocks at Sinim~ied are part of the end moraine marking the Pandivere phase of glaciation, about 13,000 to 12,500 years BP (Raukas et al. 2004; Rinterknecht et al. 2006). Immediately to the south are glaciofluvial deltas that merge southward into a glacial lake plain. The megablocks may have been pushed proglacially or dragged subglacially as ice advanced to the Pandivere limit (Rattas and Kalm 2004).
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The Baltic Klint is a tectonic fault zone along the southeastern margin of the F e n n o s c a n d i a n Shield. This zone is characterized by discrete fault blocks that were displaced by tectonic adjustments in the Paleozoic and more recently during the Neogene-Quaternary. Underlying Cambrian clay is deformed into diapirs and squeezed upward along preexisting faults. The presence of already-faulted bedrock may have facilitated glacial dislocation of megablocks at Sinim~ied, and glacial loading could have mobilized clay intrusions and diapirs. Numerous other glaciotectonic disturbances are documented in the well-consolidated Ordovician strata of northeastern Estonia (fig. 7-16). These mostly concealed structures are known mainly from boreholes. Glacially derived sediments are incorporated into faulted and brecciated bedrock or injected into karst cavities.
110
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Figure 7-15. Schematic illustration of the Baltic Klint and dislocated Ordovician limestone blocks of the Sinim~ied hills in northeastern Estonia. Taken from Rattas and Kalm (2004, fig. 2).
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Chapter 8 Intrusions, Diapirs and Wedges Introduction
Various kinds of intrusions, diapirs and wedges are widely and commonly reported from glaciogenic sequences. Intrusive structures range in lateral size from only a few cm to hundreds of meters (see fig. 1-12) and may have vertical dimensions exceeding 200 m (Morawski 2004b). Intrusive structures rarely give rise to distinctive or noticeable landforms, but they may dominate the subsurface character of a glacial sequence (fig. 8-1). Intrusions include all those structures in which one material was injected or squeezed in a mobile state into the body of another material. Some mixing between injected and host materials may take place along the intrusion boundaries, but the two materials still remain distinct. In other words, the injected and host materials do not become homogenized. The intrusive material is most usually a clay- or silt-rich sediment; whereas the host material may be almost any kind of unconsolidated sediment or soft bedrock. Creation of intrusive structures implies a significant difference in physical properties of the materials at the time of intrusion. The injected fine-grained sediment behaved as a fluid, for example quick-clay, that was mobilized by pore water trapped under high pressure. The host material, conversely, behaved in a less ductile or even brittle manner. Hydrofracturing and liquifaction are coupled processes for such intrusions (Boulton and Caban 1995). It is generally agreed that such intrusions took place in subglacial, water-saturated conditions with intergranular movement as the main mechanism of deformation (Brodzikowski and van Loon 1985). A distinction should be drawn at this point between those intrusions that originate from below and from above the host material. Diapirs, stocks, plugs, dikes and sills result when mobilized sediment intrudes from below the host material. Such structures resemble shallow intrusions of low-viscosity (basaltic) magma or diapirs formed by salt flowage. The force compelling upward intrusion is simply gravity acting on lowdensity material, such as clay- or silt-rich sediment, buried below higher-density, coarse-grained sediment or lithified strata. Diapirs may develop when fine-grained strata are compacted and become mobilized by high-pressure water. The sediment then behaves as a fluid and seeks to rise or escape toward zones of lower pressure. Upward intrusion continues as long as the sediment remains fluid, until a density equilibrium is achieved, or until the source of intrusive sediment is depleted.
Hydrodynamic blowouts with high water content have similar origins and may result in similar structures (e.g. Christiansen, Gendzwill and Meenely 1982; Rijkijk et al. 1999). The most important question concerns the nature of the compacting load that mobilized intrusion. Only those cases where glacier ice provided the diapir-inducing load should be considered glaciotectonic. Diapirs and other intrusions created by softsediment deformation induced solely by deposition of thick overburden are not truly glaciotectonic, even if they are found in glaciogenic sequences. The classic cliff exposures of chalk floes within contorted drift at Cromer, Norfolk, United Kingdom illustrate the importance of identifying the loading agent. This famous section was long thought comparable to chalk cliffs of Men, D e n m a r k (Slater 1926). However, Banham (1975) demonstrated that large, mushroom-shaped diapirs of clayrich Cromer Tills were created by loading beneath thick outwash sand and gravel, not by glacier advance (fig. 8-2). The chalk floes were incorporated in the First Cromer Till during an earlier and unrelated glaciation. Thus, the major diapiric structures of the contorted drift are not glaciotectonic. At LCnstrup Klint in Denmark, Pedersen (2005) demonstrated that extensive h y d r o d y n a m i c brecciation and mud mobilization took place in connection with thrusting of unconsolidated Quaternary strata. A series of large diapirs formed where overpressurized mud developed along hanging-
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Figure 8-1. Cartoon illustrating how large intrusive structures in the subsurface may have no morphologic expression at the surface. Adapted from Aber, Croot and Fenton (1989, fig. 7-1).
Aber and Ber
112
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Figure 8-2. Schematic section between Weybourne and Bacton, Norfolk showing large diapirs of clay-rich till containing chalk floes. 1 = Gimmingham Sands and Briton's Lanes Gravels, 2 = Cromer Tills containing chalk floes, 3 = Cromer Forest Series Bed, and 4 = Senonian chalk bedrock. Large vertical exaggeration; based on Banham (1975, fig. 14).
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L"::J 1 wall flats and propagated upward along hanging-wall ramps. Mushroom-shaped diapirs pierced through overlying thrust sheets. Glaciotectonic thrusting and diapirism happened simultaneously, as thrust sheets were carried forward on mobilized mud, so that diapirs were tilted as thrusting continued (see Rubjerg Knude, fig. 11-31). In Germany, Viete (1960) and Eissmann (1987) gave excellent examples of glacially induced diapirs. Diapiric intrusions of Miocene clays are known from the Ortowo region of Poland. The clays were squeezed upwards subglacially as a result of vertical pressure of the ice sheet during the Main Stadial of the Vistulian (Weichselian) Glaciation. The structures are recognized by high-resolution shallow seismic reflection studies (Morawski 2004b). These diapiric structures pierce more than 200 m of thick Pleistocene deposits. The
largest diapir reaches ground level forming an outcrop of Miocene deposits covering about 600 m 2, and geophysical investigations indicate the presence of several minor diapirs not reaching the ground surface. Diapiric processes were probably triggered by a deep incision formed by the subglacial Lyna River tunnel which ran in the marginal zone of the Vistulian ice sheet along the slope of an older glacial upland. The uplift of Miocene deposits caused an arcuate bend of the end part of the glacial channel (fig. 8-3). In Ukraine, according to Matoshko (1995) in the DnieperDonets region near the town of Pereyaslav-Khmelnitsky, the diapir field in Jurassic clays is approximately 300 km 2 in extent. It is associated with a zone of an approximately N-S striking regional fault. The injection structures differ in amplitude, reaching maximum heights of 140-160 m locally.
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Intrusions, diapirs and wedges
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meltwater deposits of the Dnieper Glaciation. In some rare cases, the till contains rock fragments of the same type as the injection structures. iI ~
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Wedges, veins, and fissures commonly penetrate from above along cracks opened in a brittle host material. Common wedge sediment varies from homogeneous till to laminated sand, silt or clay. The positions of wedges range from nearly vertical to nearly horizontal. Their strike orientations may be related to direction of ice movement or to direction of subglacial pressure gradient. Of course, wedge positions may also be controlled by pre-existing zones of weakness within the host material.
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Figure 8-4. Section through giant glaciotectonic diapir near Pereyaslav-Khmelnitsky, Ukraine. 1 - sand, 2 - sand and gravel, 3 - silt, 4 - diamicton, 5 - loess, 6- clay, 7- sandstone, 8 - dislocated sand and clay, 9 - boundary of diapir, 10 boreholes. Adapted from Matoshko (1995, fig. 163).
The diapirs have penetrated both Paleogene arenaceousargillaceous beds and Lower/Middle Pleistocene alluvial deposits, which have been preserved in some occasions as remnants, confined in Jurassic clays (fig. 8-4). On the basis of well drillings, the diapiric section is characterized by steep dips (40-60°), microfolding, microfaults, increased fissuring, slickensided surfaces, etc. In most cases, the dislocated Mesozoic and Cenozoic rocks are covered by till or by
A strong tendency exists among geologists to interpret wedge structures as fossil ice wedges, thermal crack fillings, or other permafrost features, perhaps because such periglacial patterned-ground phenomena are well known (Washburn 1980; Kessler and Werner 2003). At Voss, Norway for example, unsorted till wedges and laminated clay and silt veins were originally considered to be permafrost features (fig. 8-5). However, these wedges were later reinterpreted as glaciotectonic structures (Mangerud et al. 1981). Likewise, van der Meer (per. com.) reached the same conclusion for similar wedges in Switzerland (van der Meer 1980). Glaciotectonic wedges form in dynamic situations, whereas permafrost wedge fillings are created in a passive manner. Slickensides, foliation, grooves, mixing zones, drag folds, small thrusts, wedge apophyses and host xenoliths are all features that indicate a dynamic, intrusive origin for wedges. Lack of such structures is evidence for a passive wedge genesis, possibly related to permafrost conditions. Some wedges may have a dual genesis: originally formed in permafrost and later deformed glaciotectonically. Diapir and wedge structures that are glaciotectonic in origin are presented in the following case examples: Independence
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Figure 8-6. Northeastern Kansas showing Independence glacial limits, local ice movement directions, buried valleys (dotted), and locations of described sites: A = Atchison, W = Wathena. Adapted from Aber (1991, fig. 3). Reproduced from Boreas by permission of Taylor & Francis AS. Formation, Kansas; Herdla Moraine, Norway; and Systofte, Denmark. In these cases, the intrusive sediment is clay or silt rich and was injected as a result of glacier overriding. None of these structures has any morphologic expression. The Kronowo esker from Poland is a morphologic feature that was deformed in an ice-marginal setting.
Where thick Independence Formation fills preglacial valleys, as at Atchison and Wathena (fig. 8-6), its original stratification and structures can be seen in deep exposures. The Independence Formation stratotype at Atchison includes three informal members (fig. 8-7):
Atchison, Kansas The Independence glaciation was the most extensive ice coverage to take place on the Great Plains of central North America and is recognized as an important early glaciation of the Pleistocene. This glaciation is represented by the Independence Formation, a lithostratigraphic unit defined in northeastern Kansas (Aber 1991). The Independence Formation is quite old, roughly 600,000 to 700,000 years BP (Aber 1991), and is probably the oldest regionally preserved Pleistocene glacial sequence on land. Because of its great age, original glacial morphology is largely gone, and the upper portion of the Independence Formation is greatly altered by weathering and erosion. The Independence glaciation created many, scattered glaciotectonic deformations both in drift and in consolidated Paleozoic limestone and shale bedrock of the region (Dellwig and Baldwin 1965).
Figure 8-7. Photograph of Independence Formation stratotype at Atchison, Kansas. Upper and lower tills are separated by glaciolacustrine sand. I. Abdelsahed (above) and B. Nutter collect till samples from the diapir at the center of the section. Photo by J.S. Aber (1987).
Intrusions, diapirs and wedges
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• Upper till- brown, compact, stony, clay-rich till containing irregular blocks of stratified sediment, 3-15 m thick, northwesterly fabric and striations. • Middle sand- rippled, fine to very fine silty sand with scattered lenses of pebbly sand, basal portion mainly sandy pebble gravel, thickness 10 to 20 m, glaciolacustrine in origin. • \,
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• Lower till- gray, compact, stony, clay-rich till containing wood fragments, 1 to 12 m exposed, northeasterly fabric and striations. In the stratotype section, the lower till is dislocated in a pair of large diapirs which intrude up into the overlying sand (fig. 8-8). The larger diapir in the center of the section has been exposed for at least half a century. Dellwig and Baldwin (1965) interpreted the structure as a thrust anticline formed by frictional drag of ice readvancing over frozen sediments. The ice came from an easterly to northeasterly direction. Stream erosion during recent years has enlarged the section, revealing the second diapir in the lower part plus additional structures higher up. The two diapirs are strikingly similar in form and orientation. Both have enlarged, asymmetrical heads; both terminate upward at about the same level; and both dip northeastward. Till fabrics, striations on boulders, thrust faults, and the diapirs themselves show consistent orientations, measured over a period of many years, indicating ice pushing from about 60 ° (fig. 8-9). The intrusive character of the diapirs indicates deformation in a thawed and water-saturated state. The upper till caps the section and contains deformed masses of the subjacent sand. Ice movement, which laid down the
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Figure 8-10. Wathena section as it appeared in 1985. Scale in m; symbols same as Fig. 8-8. Adapted from Aber (1991, fig. 7) Reproduced from Boreas by permission of Taylor & Francis AS.. upper till, also produced a pair of recumbent folds in the underlying sand near the northeastern end of the exposure. The axes of these two folds plunge slightly toward the northeast. The recumbent folds reach down to the level of the diapir heads. The peculiar shapes of the diapir heads are explained by lowangle faults or overturned folds on their left sides. These secondary structural features are oriented toward the eastnortheast. Directional features within the upper till and deformations in the sand below, including diapir heads, all indicate ice movement from the northwest (fig. 8-9). The Wathena site (fig. 8-6) is located in gravel pits along the Missouri valley bluff. These pits have been worked intermittently for many years, resulting in a large exposure (fig. 8-10). Pre-Independence basal chert gravels are overlain by horizontally bedded fine sand and silt, above which gray lower till comprises most of the exposure (fig. 8-11). Bodies of very fine silty sand toward the top are overlain by upper brown till and younger loess. In several places, the lower till has intruded upward into the sand forming irregular plugs and sills. Likewise, certain bodies of the sand appear to have sunk into the lower till. These foundered sand bodies contain small blobs or lenses of gray till. Like the till diapirs at Atchison, the till intrusions here must have occurred in a mobilized state, when high fluid pressure had reduced the shear strength of the till's clay matrix and the intruded sand was not frozen.
Figure 8-11. Photograph showing central portion of Wathena exposure. Preglacial alluvium, which appears at the bottom (behind scale pole), is overlain by thick lower gray till. A till diapir and disturbed glaciolacustrine sand are visible toward the top. Scale pole marked in feet (1 foot -- 30 cm). Photo by J.S. Aber (1985).
Intrusions, diapirs and wedges
117
The evidence from Atchison and Wathena is consistent with ice movement from the northeast during deposition and subsequent displacement of the lower till and middle sand. Multiple northeasterly advances are probable. The initial advance laid down the lower till, and sand was then conformably deposited over the till during a brief ice recession. The sand accumulated in proglacial lakes within now-buried valleys that were blocked by ice to the northeast. Owing to its lake-bottom position, the glaciolacustrine sediment probably did not become permafrozen. Renewed northeasterly advances then dislocated the lake-floor till and sand with thrusting and diapiric intrusions. After retreat of the northeastern ice, a northwesterly advance deformed the upper portion of the sand, including diapir heads, and deposited the upper till.
southwestern Minnesota and easternmost South Dakota. At times the two lobes coalesced over the Coteau forming a broad ice fan to the south. The Independence ice lobes and fan were evidently quite dynamic.
Herdla Moraine, Norway The Herdla Moraine is a morphostratigraphic unit defined by Aarseth and Mangerud (1974). It forms one segment of the moraine system that was deposited around the periphery of Norway (fig. 8-13) during the Younger Dryas phase of glaciation (latest Weichselian). Moraine sediments are thickest and morphologically most prominent below the contemporaneous marine limit. Large ice-contact submarine fans and deltas, up to 100 m thick, are preserved in valleys and fjords both above and below present sea level. The Younger Dryas moraines of western Norway are disturbed in many places by small glaciotectonic structures, including diapirs, as a result of oscillations by the active ice margin (Sollid and Reite 1983).
Loading of lower till by either glaciolacustrine sand or ice readvance could be responsible for intrusion of diapirs. However, soft-sediment deformation seems unlikely in this case, because the maximum thickness of the sand is only 30 m. Assuming this maximum thickness plus an additional 30 m of lake water, an increased load of only about 9 kg/cm 2 was imposed on the underlying till. This is not a great load, equivalent to roughly 100 m of glacier ice.
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Figure 8-13. Location map for the Herdla Moraines western Norway: 1 = Herdla island (fig. 8-14); 2 = Herdlaflaket. Inset map shows Younger Dryas ice margin and reconstruction of ice domes and divides. Adapted from Aber and Aarseth (1988, fig. 1). The moraine on Herdla Island (fig. 8-13) is morphologically the most conspicuous in the Bergen region. The island lies across the western end of Herdlafjord and includes several bedrock knobs, between and around which the glaciomarine deposits are found (fig. 8-14). Till is present on the proximal (eastern) side of the island, and a boulder belt along the western edge marks the outer limit reached by the Younger Dryas ice sheet. Most striations and grooves on bedrock trend toward west or southwest between 240 ° and 275 °. A smaller, younger group trends toward the northwest between 290 ° and 300 ° . Clay- and silt-rich layers within the sediment are strongly consolidated. In connection with seismic stratigraphy over Herdlaflaket (fig. 8-13), the acoustic velocity of Herdla Moraines sediment was estimated to be 1800 m/s (Aber and Aarseth 1988). This compares with 1600 m/s for unconsolidated sediment. This degree of consolidation could be caused only by compaction beneath overriding ice, because the sediments have never been deeply buried otherwise.
Figure 8-14. Map of southern Herdla island showing geologic and glacial features. North to right; based on Aber and Aarseth (1988, fig. 2). Internal structure of the moraine is exposed in a road-cut section on the western side of the island above a bedrock knob. Moraine sediment there comprises several units (A-F, fig. 8-15) of interbedded clay, silt, sand and gravel deposits. This sediment was presumably laid down in water some 1015 m deep in front of the Younger Dryas ice sheet; a rapidly fluctuating sedimentary environment is indicated. The sediments were deposited as part of a submarine fan, when the ice margin was located along the eastern edge of Herdla Island. The sediments are disturbed by several secondary structures, the largest of which is a broad syncline, whose axis plunges gently toward the northeast (fig. 8-15). In the southern limb of the syncline, units B and C display structures typical of soft-sediment deformation. Coarse gravel of unit C appears to have settled or collapsed into unit B silt and sand. This deformation happened before deposition of unit D and folding of the syncline. Toward the center of the syncline, unit E forms a series of small, but elegant diapirs. The bigger diapirs, up to 1 m high, occupy the syncline trough, and progressively smaller diapirs and overturned folds are found on the syncline flanks. The smaller flank diapirs are overturned inward (toward the trough), whereas bigger trough diapirs bend outward. Sand of unit F above the diapirs is foliated and wraps smoothly around each diapir. Individual diapirs are thin, only 10-20 cm thick, and have curved shapes in the third dimension. The diapirs appear to be thin-walled partitions forming a bent, three-dimensional network or grid of intrusions (fig. 8-16). The partition network is bent according to its position within the broad syncline. Clayey silt sediment forming the diapirs thickens toward the center of the syncline, as if the
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Figure 8-1Z Summary equal-area stereonet plot of directional data for various features of the Herdla vicinity. Early phase of ice movement from east-northeast; late phase from southeast subparallel to Herdlafjord. Takenfrom Aber and Aarseth (1988, fig. 9).
Aber and Ber
120 The northern portion of the section is disrupted by thrust faults and associated folds. Unit C is thrown up along a northwest-dipping thrust that cuts through the entire sequence. This fault and related folds are subparallel with the broad syncline axis and presumably formed at the same time as the syncline. Together they produce overall structural shortening of the section in a SE-NW direction. Another thrust fault with overturned drag folds can be seen at the far northern end of the section entirely within unit D (see fig. 24). This fault strikes/dips 350/30 ° NE. These features indicate fault displacement in an ENE-WSW direction. The Younger Dryas glaciation was marked by a substantial expansion in southwestern Norway in which ice filled fjord valleys 800-1200 m deep that had been ice-free during the AllerCd. Glacier advance may have reached the Bergen district already by ~11,000 years BP, and following a brief hiatus advanced to Herdla by ~10,000 years ago (Mangerud 2004). At Herdla, two phases of moraine development took place corresponding to local changes in movement of the Younger Dryas ice sheet (fig. 8-17). During the early phase, the ice margin reached eastern Herdla, and a submarine fan was constructed. Ice movement at this time came from the east-northeast, and slight shifting of the ice margin created minor thrusting of moraine sediment. The direction of ice movement subsequently changed to southeasterly and the ice margin advanced to the western edge of Herdla. Most of the deformation in the section on western Herdla, including folding and thrusting of the broad syncline, occurred during this late phase. Consolidation of the sediment and intrusion of diapirs in a fluid state also took place at this time, due to increased loading by overriding ice. This direction of ice movement was essentially parallel to Herdlafjord. The overall position of Herdla Moraine and pattern of older striations are related to radial outflow from an ice dome to the northeast of Bergen during the early phase (fig. 8-13). The ice sheet must have been fairly thick, as ice flow was largely independent of local topography. The shift to fjordparallel ice movement in the Herdla vicinity probably occurred as the ice sheet became thinner, and ice flow was thus more responsive to local topography. The large Younger Dryas readvance in western Norway also caused crustal depression with resulting marine transgression (Anundsen 1985). Higher sea level may have increased the rate of calving and thus accelerated drawdown of the ice sheet. Another possible consequence of the marine transgression was the occurrence of glacier surges along the fjords. Surging is a common form of glacial advance for similar fjord glaciers in Spitsbergen today (Elverh¢i, LCnne and Seland 1983).
Rapid loading of saturated sediments during a glacier surge is an ideal setting for creation of diapirs and other glaciotectonic features. The glaciotectonic structures present on Herdla are, thus, related largely to activity of the local fjord glacier and may not correlate with glaciotectonic features in other portions of the Herdla Moraines.
Systofte, Falster, Denmark Humlum (1978) described a large till wedge exposed in a gravel pit near Systofte, on the island of Falster, southeastern Denmark. Various till intrusions have been noted in many parts of Denmark before (Hansen 1930; Berthelsen 1974) and must be rather common. Most of these appear to have originated from tills buried below the host. However, the wedge at Systofte was injected from above. The Systofte wedge is located in stratified sand and gravel of a kame that was overridden and is covered with about 2 m of till. The kame forms an elongated cupola-hill trending northsouth and standing a p p r o x i m a t e l y 10 m above the surrounding landscape. Cross bedding within the kame indicates southward or southwestward melt-water flow when the stratified drift was laid down. The wedge takes the form of a shallow trough, about 12 m across, that cuts at least 2.5 m deep into underlying sand and gravel (fig. 8-18). The wedge is about 1 m thick near its top, where it is continuous with the overlying till cap, but thins to only 10 cm thickness toward the bottom. The upper part of the wedge cuts across stratification in the sand and gravel, and numerous small drag folds are present adjacent to the wedge. Axes of these drag folds trend between 95 ° and 120 °, subparallel to the axis of the wedge trough. The deeper part of the wedge becomes parallel with bedding in adjacent stratified drift.
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Figure 8-18. Block diagram showing 3-dimensional geometry (from below) of till wedge at Systofie, Denmark. Small drag folds adjacent to wedge shown to left; large recumbent folds below till cap to right. Section about 12 m across; taken from Humlum (1978, fig. 6).
Intrusions, diapirs and wedges
121
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Figure 8-19. Equal-angular stereonet of till fabric (small dots) and folds at Systofte, Denmark. Solid fold axes are small drag folds; open-circle axes are large recumbent folds. Taken from Humlum (1978, fig. 7). The outer 1-2 cm of the wedge displays slight mixing of sand derived from adjacent walls, whereas the center of the wedge is enriched in fine silt and clay particles relative to the till cap. This textural zonation reflects the sorting and mixing action of fluid flow during intrusion of the wedge. The till cap has a well-developed fabric that indicates ice movement toward approximately 305 ° (fig. 8-19). The axes of large recumbent folds below the till cap are roughly at fight angles to this direction. Taken together, the structural and fabric information are consistent with ice movement toward the northwest during intrusion of the till wedge and deposition of the till cap. The last Weichselian ice to cover southeastern Denmark was the Young Baltic glaciation, consisting of East-Jylland and Ba~lthav phases. These advances came from the southeast on Falster (Houmark-Nielsen 1981, fig. 12). Humlum (1978) concluded that both till and stratified drift were unfrozen at the time of wedge intrusion. Sorting and mixing of sediment within the wedge and intensive drag folding of adjacent sand could hardly happen in frozen material. This implies that the Systofte wedge was intruded beneath a temperate glacier base and that the Young Baltic glaciation took place over unfrozen ground at this location.
Kronowo esker, Poland The Kronowo esker, located in western Mazury, northeastern Poland near Kronowo village (fig. 8-20), is >3 km long and approximately 200 m wide. Morawski (2003) first described this esker and interpreted it as deposits that accumulated in an ice crevasse. It runs from the NE towards the SW, with a slight ENE deviation in its eastern part. This landform is represented by a row of elongated hills from several m to a dozen m high. It runs across a hummocky glacial plateau
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Figure 8-20. Geomorphological sketch map of the Kronowo region: 1 - morainic plateau, 2 - dead-ice moraine, 3 landforms developed in ice crevasses (eskers), 4 - kames, 5 - kettle holes and valleys, 6 - glacial tunnel valleys, 7 outwash plains, 8 - ice-dammed areas. Adapted from Morawski (2003, fig. 1). covered with till, several m to >20 m thick. At culminations of individual hills, reaching 150-160 m elevation along the linear ridge of the Kronowo esker, sand-and-gravel glaciofluvial deposits are exposed on the surface, and flow till is present elsewhere on the ridge. The top of the esker, an erosional surface, is disconformably covered by a thin pavement layer of poorly sorted, massive deposits followed by a light brown sandy diamicton containing clearly visible structures indicating transport in mud flows. The thickness of this till ranges up to about 8 m in the northeastern part of the exposure. However, the thickness of the esker deposits varies from a dozen to >20 m. The deposits are represented largely by cross-bedded, poorly sorted sands and gravels. At the Kronowo exposure, the sandand-gravel esker succession is underlain by a compact till, 1.5 to 3 m thick. The till occurs only within the ice crevasse composing a basal layer of its fill. This layer lies at the same depth, or slightly lower, as the base of till composing the
Aber and Ber
122
glacial plateau surrounding the esker. The esker succession fills a tunnel cut into the Vistulian (Weichselian) till (Morawski 2003).
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Detailed analysis of the topographic trend of the Kronowo esker indicates that this is not a continuous ridge, and its individual parts are shifted relative to each other and arranged en echelon (fig. 8-21). Planes along which these shifts occur represent NW-SE oriented strike-slip faults. One of these faults shifts the quarried gravel-and-sand deposit, causing the NE quarrying works to be relocated 80 m northwards. These are glaciotectonic faults caused by NW-SE oriented horizontal stress exerted by the active ice body.
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Figure 8-21. Sketch map of the Kronowo esker: 1 - sand and gravel mantled with a thin flow till (esker sediments), 2 - till (Vistulian glaciation morainic plateau), 3 morphological axis, 4 - strike-slip fault zones, 5 - fan fold axis, 6 - pit boundaries, 7 - cross-section, 8 - borehole. Adapted from Morawski (2003, fig. 7).
Exposures in the gravel pit reveal strong compressional glaciotectonic deformations forming two fold limbs which are in contact and thrust over each other (figs. 8-22 and 23). The strike of the thrust surface is about 80 ° . The surface is inclined northwards at an angle of 60-70 °, indicating that the horizontal stress was directed from the north. This direction is also indicated by the fold asymmetry and by small, more-or-less horizontal faults observed in both fold limbs, particularly in the northern limb. The till under the esker series in lower parts of both fold limbs, north and south of its axis, lies horizontally in the original position. Towards the axis, the till layer is bent arcuately to reach a vertical position, and is inclined towards both sides. The arc diameter is about 20 m. The till, in both the exposures and boreholes, is dislocated by bending and squeezing upwards. In the southern fold limb, this break is about 10 m long; in the northern fold limb the till is truncated by a thrust surface.
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Figure 8-22. Geological cross-section of the pit walls with interpretation of glaciotectonic fold: 1 - f l o w till, 2 - till of the basal layer, 3 - esker sand and gravel, 4 - glaciofluvial sand and gravel of the esker basement, 5 - boreholes. Adapted from Morawski (2003, fig. 8).
123
Intrusions, diapirs and wedges
2
Figure 8-23. Photograph of the southern side of the fan fold with intrusive fold core to right side. Compare with left side of previous figure. Photo by Morawski (2003, fig. 11).
......
All the data collected by Morawski (2003) indicate that the original width of the crevasse in the Vistulian ice sheet was approximately 200 m. This crevasse was initially filled with diamicton, probably originating from mud flows and/or from the melted and collapsed tunnel roof. This sediment was probably subjected to short-lived erosion and subsequently deposited as till, up to several meters thick. This phase was followed by deposition of a sand-gravel succession, over 20 m in thickness, transported from NE to SW. After sedimentation and filling of the ice crevasse, probably already during the initial deglaciation, the ice sheet readvanced southwards for a short time. Only part of the ice sheet, located on the north of the crevasse, was presumably active at that time. The southern area was covered by stagnant ice. This readvance resulted in deformation and upward squeezing of the crevasse-fill (fig. 8-24, phases 1, 2). Squeezing of both the crevasse-fill and underlying deposits was initiated and took place with a contribution of vertical (loading) stress of the ice body, as was the case during formation of diapiric structures in ice-free areas (Brodzikowski 1980) and in squeezed or till-cored eskers (Rotnicki 1960). In the case of Kronowo, however, the vergence of glaciotectonic structures and the shift of individual parts of the esker by strike-slip faults, indicate a strong lateral push by the ice from the north. Such a push resulted in the formation of a large fan fold. The fold limbs were subsequently stretched and broken by expansion of the diapiric core. The limbs tilted out toward both sides and then were partly truncated by a thrust (fig. 8-24, phase 3 and 4). In the area exposed by quarrying, the distance over which the ice sheet advanced (magnitude of glaciotectonic narrowing of the ice crevasse) was determined as a sum of deformed and broken strata, and amounts to about 80 m.
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Figure 8-24. Development stages of the Kronowo crevasse infilling and glaciotectonic folding: 1 -Kronowo crevasse and reactivated ice sheet from the north (arrow indicates the direction of ice movement) initiated folding of crevasse fill; 2, 3 and 4 - successive phases of ice-sheet movement and crevasse-fill folding, crevasse width reduced to about 80 m, 5 - deposition of flow till after periodic erosion. Adapted from Morawski (2003, fig. 15).
Subsequent movement of the ice sheet resulted in en block and en echelon shifting of the crevasse-fill. The general direction of the shift was towards southeast. Individual horizontal shifts within the esker ridge can be estimated at 20-80 m. It can be approximated that local advance of the
124 ice sheet in the Kronowo region totaled a distance of ~150 m. The upper portion of the folded crevasse-fill was probably eroded by melt water. These deposits were subsequently covered by diamicton flowing down into the crevasse from melting dead ice blocks (fig. 8-24, phase 5). These observations indicate the Kronowo esker was formed by a crevasse-filling process in the continuously active ice sheet, or maybe during the initial phase of its stagnation and minor readvance (Morawski 2003). Meltwater erosion was followed by deposition, and the crevasse was infilled with flow tills and glaciofluvial sediments. The glaciotectonic fan fold was formed as a result of subsequent compression of the ice crevasse. The process of upward-squeezing of the crevassefill and partly also of the underlying glaciofluvial deposits (in a fold core) was probably caused by the same ice sheet movement which subsequently triggered strike-slip faulting across the esker ridge. The ability to define local directions of glacier movement from the spatial orientation of compressional glaciotectonic deformations is straightforward in most situations (chap. 2). However, the Kronowo esker was transected by strike-slip faults
Aber and Ber
oriented obliquely to the main folding stress. The ice mass presumably contained pre-existing crevasses (joint net), along which local movements of the ice body could occur. The ice sheet push was exerted on deposits filling a crevasse that trended obliquely to the direction of this push. Individual glacier blocks advanced not in the same direction as the main ice mass; they moved along crack surfaces, which resulted in strike-slip faults. On the basis of deformation of the Kronowo esker and comparison to regional glaciation, directional shifting in ice movement can be reconstructed. Ice advance took place from the north during the entire Main Stadial of the Vistulian (Weichselian) glaciation. Following ice retreat and stagnation, a minor readvance came from the NNW, which led to folding and compression of the esker ridge. Finally the esker was displaced along strike-slip faults trending NW-SE as a consequence of local movement of ice blocks along a crevasse network. The Kronowo exposures provide a unique opportunity for precise determination of the direction of icesheet movement during a short-lived period of its activity that occurred prior to the final deglaciation.
Chapter 9 Basement and Deep Crustal Structures Introduction An important outcome of the glaciotectonic mapping project in North America was recognition of basement and deep crustal structures as a significant category of glaciotectonism (Aber et al. 1995). Such features include basement faults and seismic zones wherein the crust responded differentially as discrete blocks to stress induced by glacial loading and unloading of the lithosphere. Such deformations are usually manifested at the time of deglaciation or shortly thereafter. Residual crustal strain may continue with decreasing amplitude and frequency long after the disappearance of ice loading. For example, an arcuate band of seismic activity is present along the Bell and Boothia arches of northernmost Hudson Bay and Foxe Channel. Basham, Forsyth and Wetmiller (1977) speculated that the Foxe-Baffin crustal block is responding to postglacial uplift independently from the rest of the Canadian Shield. The seismic band follows the edge of the Foxe sector of the Laurentide Ice Sheet. Seismic activity and block tilting toward the northeast were interpreted as the results of reactivation of preglacial structures caused by high differential stress during glacial unloading. Adams (1989) similarly suggested that much of the postglacial faulting in southeastern Canada and the adjacent United States resulted from stress release and flexural deformation of the upper crust as a result of glacial loading and unloading. The regional-scale location of faults coincides with the southern margin of the Laurentide Ice Sheet. Several zones of seismic activity are found at or near the southern limits of glaciation in stable continental crust of the United States and southeastern Canada. • Nemaha zone: northeastern Kansas and southeastern Nebraska; maximum Independence (pre-Illinoian) glaciation limit (Aber 1991). • New Madrid zone: southeastern Missouri, southern Illinois, and adjacent states; maximum Illinoian glaciation limit. • Anna zone: western Ohio; maximum late Wisconsin glaciation limit. • Atlantic coastal zone: from northern New Jersey to southern New Brunswick; maximum late Wisconsin glaciation limit. • Saint Lawrence zone: northern New York to southeastern Quebec; deglaciation ice limit of 12,000 years BP (Dyke and Prest 1987b).
These seismic zones lie at the intersections of northeasttrending (late Proterozoic) and n o r t h w e s t - t r e n d i n g (Mesozoic) basement structures, and the seismicity of these zones is related to continued opening of the North Atlantic Ocean (Barosh 1990). Aber et al. (1993) speculated that earthquakes in these and other ice-marginal seismic zones were accentuated by glacial loading and unloading. Other crustal disturbances one km or deeper in the subsurface also are included in this category. Incompetent strata, such as salt, were mobilized under the impact of glacial loading at considerable depth. As an example, White (1992) documented displacement of salt south of Lake Erie in northeastern Ohio (fig. 9-1). The combined thickness of Paleozoic salt beds exceeds 60 m in the region south of Cleveland near the limit of glaciation. The thickest salt (>90 m) is found in narrow zones within the cores of doubly plunging folds. The fold axes trend generally parallel to the south shore of Lake Erie and to the glacial limit. Near the limit of glaciation, the folds become somewhat arcuate and are spaced closer together. White concluded that stress imposed by successive glaciations in the Erie basin had mobilized the salt southward. The following examples include basement faults and seismic zones from the Canadian Shield and Appalachians in North America and the F e n n o s c a n d i a n Shield of northern Scandinavia in Europe. Deep salt displacement is examined in the Finger Lakes district of New York, and buried faults are described from Poland.
Canadian Shield and northern Appalachians This class of glaciotectonic structures is widely documented in the northern Appalachian region and from the southern Canadian Shield. The structures are of two types: small faults in basement rocks and zones of seismic activity. The small faults included are those considered to be the direct results of glacial loading and unloading, in contrast to faults resulting from regional tectonics, mine pop-ups, or frost heave. Glaciotectonic basement faults are found primarily in Paleozoic slates of the Appalachian region (Chalmers 1897; Goldthwait 1924) and in Proterozoic slates of the Canadian Shield of western Ontario (Lawson 1911; Oliver, Johnson and Dorman 1970). Small faults are also reported from other regions and from well-consolidated argillite, sandstone and limestone (Cushing et al. 1910; Broster and Burke 1990).
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Late glacial, postglacial, and historical earthquakes have shaken this region (e.g. Broster, Allen and Burke 1994). Shilts, Rappol and Blais (1992) presented multiple lines of evidence for paleoseismic events in the T6miscouataMadawaska Valley, southeastern Quebec and adjacent New Brunswick (fig. 9-5). This valley is a NW-SE lineament that was generally parallel to glacier movement. Older striations and erratics indicate ice flow was toward the southeast, but during later phases ice movement reversed toward the northwest as a consequence of glacier drawdown toward a calving bay in the lower St. Lawrence lowland. A large proglacial lake, Lake Madawaska, inundated the valley during and shortly after deglaciation. Through a series of drainage diversions related to postglacial erosion and crustal rebound, the modern Madawaska River and Lac T6miscouata came into being. Near Saint-Jacques, Shilts, Rappol and Blais (1992) described a small fault in slate that exhibited 7 cm of vertical offset of the striated rock surface. Also present in this vicinity are massive boulder deposits consisting of crudely layered rock debris resting on diamicton and glaciolacustrine strata containing numerous contortions and shear zones. They interpreted these as rock-avalanche or debris-flow deposits within glacial Lake Madawaska. The sediment fill beneath Lac T6miscouata was explored with subbottom acoustic profiling (SAP). Sediments with acoustically chaotic internal structure and hummocky surface morphology were found in many places within the lake and also nearby at Grand lac Squatec (fig. 9-6).
Basement and deep crustal structures
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fig. 1;. Individually these diverse features could be interpreted in several ways. Taken together, however, only earthquakes or severe seismic vibrations could explain these structural disturbances within the T6miscouata-Madawaska Valley (Shilts, Rappol and Blais 1992). Although the valley is historically aseismic, they concluded that it had been subjected to repeated, strong, late glacial and early postglacial seismic shocks. They further concluded that these seismic events were a response to vertical stress relief accompanying deglaciation and crustal isostatic adjustments. VERTICAL
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Figure 9-4. Schematic diagram of ice-load stress (above) and resulting small faults following deglaciation (below). Based on reverse faulting in slaty hills at Saint John, New Brunswick. Adapted from Broster and Bruce (1990, fig. 6).
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Northernmost Sweden, Norway, and adjacent Karelia, Russia have several large surficial faults in the Fennoscandian Shield, a region that is historically aseismic. These faults trend generally SW-NE and may be traced individually for 10s to 100s of km (fig. 9-7). The longest of these is the Piirve Fault, which extends as a nearly continuous fault-line scarp for about 150 km (fig. 9-8). It appears like a stationary wave running across the otherwise low-relief landscape. The fault is named after the Lapp word, "piirve," which refers to a breaking wave (Lundqvist and Lagerb~ick 1976). The Piirve Fault offsets several different Precambrian rock units. Throughout its whole length, downthrow is toward the west or northwest, and vertical displacement is up to 10 m. The fault plane is high angle with predominantly dip-
Aber and Ber
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Basement and deep crustal structures
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Figure 9-7. Map of recent earthquakes (1987-97) throughout Scandinavia and surroundings. Large circles >2.5 magnitude; small circles <2.5 magnitude. Large faults depicted in northern Sweden, Norway and Karelia (heavy black lines). Asterisk shows location of Piirve Fault. Diagonal lining indicates region with modern uplift >7 mm per year. Adapted from Gregersen (2002, fig. 3). slip movement. In places, steep eastward dip indicates reversed movement. Its trend is generally straight or gently curved, although a sharp double curve is present toward the northern end near Lake Kamasjaure, where it splits into multiple faults that form a small horst. Along most of its
length, the fault forms a marked step in the till-covered landscape. In places, bedrock is exposed in cliffs, and springs are common along the fault. The fresh appearance of the fault suggests it was offset after the last glaciation.
Aber and Ber
130
along the fault reflects greater ice loading toward the southeast, where highest uplift rates are centered today (fig. 9-7). Retreat of the ice mass was from northeast toward southwest, that is parallel to rather than across the fault trend. Thus, fault displacements may have been diachronous, migrating from northeast to southwest (Lundqvist and Lagerb~ick 1976).
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Since recognition of the P~irve Fault, several other roughly parallel faults have been documented in northernmost Sweden, Norway and Karelia (fig. 9-7). Lagerb~ick (1990) argued that such faults are evidence for large earthquakes. Further indicators of paleoseismicity include landslides and liquifaction of sediments. These faults are collectively about 9000 years old, an age which coincides closely with regional deglaciation. At this time, the lithosphere underwent fast adjustment to changing stress conditions as the ice mass was reduced. The stresses derived from two sources (Gregersen 2002).
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• Stored plate-tectonic stress: during glaciation the lithosphere is effectively locked, so that plate-tectonic stress gradually accumulates. Seismicity is extremely low in modern Greenland and Antarctica for this reason (Johnston and Kanter 1990). Presumably this mechanism applied to the Fennoscandian Ice Sheet as well. Removal of the ice mass allowed this stress to be released. • Isostatic stress: between 13,000 and 9000 years ago, about 500 m of regional uplift occurred, and since then another 300 m of rebound has taken place. Differential stress developed from transgressive removal of the ice mass. The combination of these two stresses led to widespread faulting and earthquakes in a pulse of activity about 9000 years ago. Since then, seismic activity has diminished, and northern Scandinavia now is subject to plate tectonic stress related to North Atlantic sea-floor spreading. Modern seismicity is relatively low in northern Scandinavia (fig. 97), and the probability for a large earthquake on these faults is considered to be practically zero (Gregersen 2002).
Salt displacement, Finger Lakes region, New York Postglacial movement of the fault is indicated by offset, truncation, gaps or irregularities of several morphologic features including: eskers, drumlins, roches moutonn6es, melt-water channels, outwash plains, and beaches of proglacial lakes. However, a few places have evidence for fault displacement older than glacial landforms. In total, Lundqvist and Lagerb~ick (1976) concluded that portions of the P~irve Fault experienced movements that indisputably postdate late glacial morphologic features. Rupture of the fault was a consequence of rapid deglaciation, and may have been facilitated by melt water lubrication of pre-exisiting fractures. Consistent northwesterly downthrow
Salt deformation, or halokinesis, is relatively common in the geological record because of two attributes of salt. First, it is lower in density than most other common sediments and rocks. Second, salt is subject to plastic deformation under relatively low stresses. Thus, salt buried beneath sedimentary strata has a strong tendency to flow into diapirs that pierce overlying strata and rise toward the surface. Another potential source of loading is glacier ice, but salt is more dense than ice, so salt would flow away from the load source toward zones of lower pressure. In the case of central New York, salt migrated southward away from glacial loading and downdip toward deeper parts of its sedimentary basin.
Basement and deep crustal structures
131
Figure 9-9. Near-vertical space-shuttle photograph of the Finger Lakes district of central New York. Eleven lakes are found in two groupsmeastern and western. These lakes occupy deep ice-carved valleys that resemble inland fjords. Selected lakes: 1 - Conesus, 2 - Canandaigua, 3 - Keuka, 4 - Seneca, 5 - Cayuga, 6- Owasco, and 7- Skaneateles. Image adapted from NASA Johnson Space Center, STS 51B-33-028, 4/85. Courtesy of K. Lulla.
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Figure 9-11. Panoramic image of Keuka Lake viewing toward the northeast. The lake is about 1.2 km wide to right and branches in two major arms toward left. Long, deep, ice-carved valleys are reminiscent of fjords. Photo by J.S. Aber (2005).
Central New York contains two major physiographic regions--Appalachian Plateau and Ontario Lowlands, which are separated by the Niagara Escarpment (fig. 9-9). Surface elevations range from 150-200 m in the Ontario Lowlands to 400-800 m in uplands of the Appalachian Plateau. The lowland between Lake Ontario and the Niagara Escarpment contains many thousands of drumlins. To the south, a zone of severe glacial erosion is marked by the Finger Lakes (figs. 9-10 and 11). These lakes occupy fjord-like valleys created by concentrated glacial and subglacial melt-water erosion (Mullins and Hinchey 1989). Bedrock floors beneath the larger valleys are 200-400 m below present lake levels. Encompassing the southern side of the Finger Lakes is the
Valley Heads moraine, which marks the southern limit of deep valley erosion and forms the modern drainage divide between the Susquehanna and Ontario basins (Cadwell and Muller 2004). Farther south, the Appalachian Plateau has a rugged upland topography, and fiver valleys delineate icemarginal drainage routes. The whole morphologic assemblage reflects a lobate style of glaciation in which ice flow radiated southward from the Lake Ontario basin. The glacial morphology of central New York represents a mature, end product of multiple glaciations (Coates 1974). The Ontario Lowland is underlain by lower Paleozoic sedimentary strata, including the salt-bearing Salina Group
OUTCROP OF SALINA GROUP
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Figure 9-12. Salt isopachs and structure contours on top of salt zone for salt in the Salina Group (upper Silurian), central and western New York. Structure contours based on sea level. Taken from White (1992, fig. 3).
Basement and deep crustal structures
133
(uppermost Silurian). These strata dip southward beneath the Appalachian Plateau. The Niagara Escarpment is formed by the Onondaga Limestone (lowermost middle Devonian), and thick middle and upper Devonian clastic strata comprise the upland region to the south (Rickard and Fisher 1970). Salt has long been mined in the Finger Lakes district, and its regional extent is known from much prospecting (fig. 9-12). From the Salina Group outcrop, salt dips southward reaching > 1 km depth near the New York-Pennsylvania border. White (1992) noted several anomalies in salt distribution. • Structural depression (basin) on the top of the salt zone in the northern Finger Lakes district. • Thinning of salt (isopachs) in the northern Finger Lakes district. • Thickening of salt (isopachs) in the southern Finger Lakes district. • Maximum salt thickness (>300 m) south of the Finger Lakes.
multiple Pleistocene glaciations, the region north of the Valley Heads moraine was covered many times, as evidenced by the magnitude of glacial landscape modification. However, only rarely during maximal phases did glaciation extend into the Appalachian Plateau beyond the Valley Heads moraine. The Ontario Lowland and northern margin of the Appalachian Plateau were, therefore, covered by relatively thick ice more frequently and for longer periods than was the region south of the Valley Heads moraine (Cadwell and Muller 2004). The cumulative effect of repeated differential loading by glaciation was migration of salt away from the northern Finger Lakes vicinity and toward the zone of lesser ice loading south of the Valley Heads moraine. Depth of salt displacement reached 1-2 km in south-central New York and northeastern Pennsylvania. A related structure is a salt-cored anticline that lies beneath and parallel to the N-S axis of Lake Seneca (fig. 9-13). Glacial erosion removed considerable overburden from this longest and deepest of the Finger Lake valleys. Upon deglaciation, salt flowed toward this zone of lower overburden pressure (L.E Dellwig, pers. com.).
The anomalous distribution of salt is restricted to the Finger Lakes region and does not extend eastward or westward. Farther south, in northeastern Pennsylvania, thickened salt cores are found in overturned and thrust folds of the Appalachian Plateau >2 km deep. The direction of overturning and thrusting is to the southeast, toward the Appalachian Mountains. White (1992) interpreted these anomalies as results of massive salt displacement southward in the downdip, downglacier direction under the impact of glacial loading over the Ontario Lowland region.
Poland contains several excellent examples of deeply buried basement structures that were reactivated during repeated glacial loading and unloading, and may have contributed to surficial glaciotectonic deformations as well. These structures have been documented through extensive test drilling and geophysical investigations.
Ice thickness in the Ontario Lowland region was on the order of 1500-2000 m, a mass load considered sufficient to induce halokinesis (White 1992). In contrast, ice thickness over the Appalachian Plateau south of the Valley Heads moraine was likely no more than a few 100 m. During the course of
Buried faults, Suwalki Lakeland m The Suwatki Lakeland is situated in northeastern Poland bordering with Russia (Kaliningrad district) and Lithuania (fig. 9-14). Analysis of geological data obtained from deep boreholes drilled in this area during past half century shows clear influence of the
Polish basement structures
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134
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Figure 9-14. Tectonic map of the Suwatki Lakeland district in northeastern Poland with glacial surficial features. 1 - tectonic discontinuities, 2 - main faults, 3 -Suwatki Anorthosite Massif (SAM), 4 - Wi~.ajny Elevation, 5- glacial features of phases and subphases of the Vistulian Glaciation, 6 - glacial tunnel valleys, 7- zones of glaciotectonic deformations. After Ber (2000, fig. 3; modified).
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gravity and seismic investigations, as well as maps of the sedimentary cover. Depth to the crystalline basement ranges from only 0.5 km in the northeast to more than 1.6 km in the northwest of the Suwatki Lakeland.
Geological data concerning the crystalline basement (Kubicki and Ryka 1982; Graniczny 1998) are based on results of
Since late Proterozoic time, the crystalline basement of this region has been subjected to continuing and repeated stress
135
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occurred several times in the Suwatki Lakeland. Such crustal movements were often focused along pre-existing fault zones and left an imprint in surficial glacial geomorphology (fig. 9-15). It can be assumed that loading and relaxation within the crystalline basement occurred several times, during each glacial and interglacial period, causing rhythmic glaciostatic movements. The amount of vertical depression during icesheet advance was up to 200 m, depending upon the ice thickness and duration. In interglacial periods, relaxation and complete rebound of the crust took place. The ice sheets from earlier, pre-Vistulian glaciations were thicker and more widespread. During their advances they crossed latitudinal fracture zones where ice-laid deposits were subjected to glaciotectonic deformations (figs. 9-14). It cannot be precluded that the process of ice sheet retreat could depend partly on the structural pattern of the deep basement, i.e.
conditions that led to formation of new and reactivation of pre-existing meridional (N-S) and latitudinal (E-W) tectonic fault zones (Znosko 1984). North-trending faults played a special role within the Suwalki anorthosite massif (SAM, fig. 9-14). They divided the area into low and high tectonic blocks already in the late Proterozoic. Tectonic movements of the crystalline basement were superimposed by glaciostatic loading and unloading, which took place during glacier advance and retreat (Weertman 1961; M6rner 1977, 1980; Liszkowski 1993). The former resulted in crustal depression and the latter in rebound, which P~TIC m
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variable activity along meridional and latitudinal faults separating individual tectonic blocks and differential shifting of the blocks. Glaciostasy, operating along dislocation zones and faults, exerted in turn an essential influence on the Vistulian glaciation in the Suwa~ki Lakeland and throughout northeastern Poland. Vertical movements along latitudinal and near-latitudinal faults conditioned development of parallel-trending river valleys, topographic lows and highs, maximum limits of ice sheets, and recession and oscillatory ice-marginal zones, where glaciotectonically deformed moraines are observed (Liszkowski 1993; Ber 2000). Zones of glaciotectonic deformations are associated with basement faults and vast stepwise tectonic fault zones; for example, the Szczytno-Okr~gte Lake dislocation zone (SJO-DZ, fig. 9-16) varies from a width of 10s of meters (northeastern Poland) up to >100 km in Lithuania (Sliaupa 1996). In general, north-, NNE- and NNW-trending (meridional) faults exerted an influence on the formation of similarity in orientation of subglacial and surface crevasses in the ice body, resulting in the development of eskers and subglacial tunnel valleys. A similar situation existed in Estonia (Tavast 1998); conversely in Belarus, eskers were controlled by latitudinal faults (Karabanov 1997). Pleistocene deposits of the Suwatki Lakeland were subjected to glaciotectonic deformations due to advance and retreat of
Figure 9-19. Main configuration of Neogene and Quaternary substrates and mid-Pleistocene horizons in western Mazury and Warmia, northeastern Poland. 1 - depressions of Neogene substrate, 2 - depressions of Quaternary substrate, 3 - area occupied by Holsteinian Sea, 4 - area occupied by Eemian Sea, 5 - main streams of Mazovian Interglacial, 6 main streams of Eemian Interglacial. Adapted from Marks (1988, fig. 19).
successive ice sheets. Deep-seated faults manifested themselves as flexures in glacial and glaciofluvial deposits. Through the whole Pleistocene, glaciostatic movements operated along meridional (N-S) and latitudinal (E-W) dislocation zones and faults. The most legible record of this effect is visible in the present-day relief of the Suwatki Lakeland, which was conditioned by ice streams, lobes and tongues that followed in turn patterns of the crystalline basement and its sedimentary cover, including Pleistocene deposits (Ber 1987, 2000; Ber and Ryka 1998). The consistency of the main orientations and present-day relief
137
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with the tectonic pattern of crystalline basement (fig. 9-14) demonstrates that pre-existing tectonic zones were rejuvenated in the Pleistocene due to differential crustal movements, which is also confirmed by photolineament analysis (Graniczny 1998). The Wi2ajny Elevation, located in the northern part of the Suwatki Lakeland, had a special influence on variability and movement directions of the Vistulian ice sheet, and also older ice sheets. The ridge developed on the pre-existing paleostructure composed of Cambrian and Valdai deposits. Owing to the tectonic situation through the whole Pleistocene history, it was an isolated basement high, similar to those observed in other areas of northeastern Poland (Szeskie Hills, Elbl~g Upland, Gdrowo Hills), in Belarus (Grodno Upland), and Lithuania (~emajtija Upland), as well as Latvia and Estonia. All of these are characterized by strong glaciotectonic deformations of surficial strata (fig. 9-17). The Vistulian ice sheet caused deformation of the basement, inducing static (vertical) stress and dynamic (horizontal) stress, as presented in a model by Rotnicki (1976). The stresses resulted in reactivation of tectonic fracture zones and mainly latitudinal faults controlling ice sheet extents and
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stagnation periods during recession (fig. 9-14). This gave rise to the formation of various ice-marginal landforms, of which festoon relief and valley-side glaciotectonics are distinctive examples (Ber 1987, 2000; fig. 9-18). Stressinduced deformation activity of the ice sheet, as well as the formation of festoons, occurred mainly in areas composed of till. It may be possible that development of glaciodepressions and glacioelevations, forming large festoons and lobes, was associated with mobility of individual tectonic blocks of the crystalline basement. Warmia and western Mazury Lakeland w The present relief of the Warmia and western Mazury Lakeland (northeastern Poland) was shaped by glaciotectonics and neotectonics connected with the structural plan of the old basement and activated along meridional (N-S) and especially by parallel (E-W) fractured zones. These zones were formed at the lithologic boundaries between Precambrian structural units of different density. Especially, the parallel geological boundaries of the crystalline basement and large faults connected with them had a basic effect on the parallel direction of many glacial forms of the recent landscape, such as marginal zones of the Vistulian Glaciation, hills of end
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and push moraines, and subglacial channels. Parallel boundaries likewise influenced zones of glaciotectonic deformations (Marks 1988; Ber 2000; Morawski 2004a). In the Warmia region, according to Marks (1988), the surface of the Quaternary substrate has a system of elongated elevations and depressions with general NNW-SSE direction, which are associated with features of the Neogene substrate and depressions and faults in the top surface of Cretaceous rocks (fig. 9-19). Similar direction (NNW-SSE) is marked in the system of river valleys of the Mazovian (Holsteinian) Interglacial (see fig. 9-22). However, the fault zones within the Mesozoic complex, known from vicinity of Lidzbark Welski (fig. 9-20) and ~uromin (fig. 9-21), showed at the same time a gradual vanishing of faults downwards and a change of throw or of its direction. This demonstrates that fault zones have been rejuvenated during the Pleistocene many times (Marks 1988). The isolated uplands in the present landscape in this area (Elbl~g Upland, G6rowo Hills and Dylewska Hill) contain strongly disturbed glaciotectonic internal structure and correspond with buried positive pre-Arenigian paleostructures composed of Cambrian and Valdai deposits,
named by Kotafiski (1977)" Elbl~g Elevation, G6rowo Itaweckie Hump and Dylewska Bank (fig. 9-22). In the present landscape of the southwestern Mazury, the Lubawa Upland with its highest peak of Dylewska Hill (312 m elevation) played a significant role. Dylewska Hill is composed of upthrusted Paleogene, Neogene and Quaternary sediments (Marks 1978; Gat~zka and Marks 2000) and corresponds with a pre-Arenigian high in tectonically predisposed deep basement. To the east of Dylewska Hill, the Quaternary substrate is disintegrated into some elevations and depressions, running NNW-SSE and highly concordant with the deeper, pre-Arenigian paleostructures. This coincidence connected with the activation of fault zones, according to Marks (1988), affected palaeogeography of the area during the Pleistocene, especially since the Mazovian (Holsteinian) Interglacial (see figs. 9-19 and 9-22). In the western Mazury Lakeland, deeply rooted glaciotectonic thrust structures, reaching a depth of about 300 m, have been found near Ortowo (Morawski 2004a). Lateral extent of these structures is up to 5 km and Oligocene through Pleistocene (Odranian Glaciation) deposits are involved. The maximum gradients of the thrust surfaces range from 15 ° to 20 ° and
139
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the thrust vergence is towards NE, opposite to the general direction of ice sheet movements in the Polish Lowlands. The thrusts, according to Morawski (2004b), seem to follow a structural slope, probably formed in the place of the buried faults occurring in the topmost Cretaceous through Miocene strata, and they developed probably as a result of vertical pressure exerted by the ice on the deeper basement. The main structural elements of the glaciotectonic deformations are glaciotectonic thrusts (slices) and diapiric upthrustings (see fig. 2-19). As Morawski (2004b) described it, slices thrust over each other, create a contractional duplex, with the vertical component of the stacking process being a derivative
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Figure 9-25. Geological section through Ostrzesz6w Hills. 1 - upper Triassic, 2 - middle Miocene, 3 - upper Miocene, 4 Pliocene, 5 - Middle Polish glacial strata, 6 - Quaternary formations on undisturbed substratum. Well-consolidated Triassic bedrock forms a horst beneath poorly or unconsolidated younger layers. Vertical scale in m above sea level. Adapted from Markiewicz and Winnicki (1997, fig. 2).
140
Aber and Ber
of pressure, which took place prior to or during Odranian (Drenthe) Glaciation.
NW ODOLAN6W BASIN --~
Ostrzesz6w Hills, Wielkopolska Lowland ~ The Ostrzesz6w Hills, situated at the southeastern end of the Wielkopolska Lowland, are the easternmost and highest part of the Silesian Rampart, which includes several other major glaciotectonic uplifts, such as Trzebnica Hills (fig. 9-23). Altitude of Ostrzesz6w Hills exceeds 280 m. Ostrzesz6w Hills extend 34 km N-S, are 4 to 10 km wide, and display an arcuate curvature, visible on Landsat images (Aber, Ruszczyfiska-Szenajch and Krzyszkowski 1995), concave toward the west (fig. 9-24). The area extending toward the west and northwest from Ostrzesz6w Hills is the Odolan6w basin with its bottom at the level ~120 m above sea level, which marks surficially the region of the buried depression within the Neogene substratum.
Rotnicki (1967) interpreted Ostrzesz6w Hills as a thrust end moraine, pushed from the west by the Wartanian (Saalian) Glaciation. Internal structure consists of steeply dipping imbricated blocks, pushed from the west, that include Pleistocene, Pliocene and Miocene sediments. West of Ostrzesz6w Hills the glaciodepression of Odolandw basin marks the source region for material thrust into the hills (fig. 9-24). Odolan6w basin and Ostrzesz6w Hills are localized within the dislocation zone of block structures above a horst within hard Triassic bedrock, beneath unconsolidated Neogene sediments (fig. 9-25). According to Markiewicz and Winnicki (1997) this situation is most important for the reconstruction of the factors stimulating the origin of deep glaciotectonic deformations occuring in this area, which were favored by diversified geological structure of the Neogene upland with positive (horst) structures of hard substratum (fig. 9-26). Glaciotectonic movements of the Triassic substratum probably were connected with processes resulting from irregular ice loads exerted repeatedly during multiple glaciations. Dislocations of block structures were influenced additionally by halotectonic movements of the underlying Zechstein salt beds, again enhanced by multiple glaciations.
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Figure 9-26. Hypothetical sequence of glaciotectonic and tectonic processes in the region of Ostrzesz6w Hills. Stages: a - ice sheet transgression into Odolan6w basin, b - advanced transgression of ice sheet, c - deglaciation. Pleistocene (Q): 1 - ice sheet, 2 - sediments of Nidanian Glaciation. Plioceneolder Pleistocene (Q+Pl): 3 - sand and gravel of Gozdnica Series. Upper Miocene (M3): 4 - clay and silt. Middle Miocene (M2): 5 - clay, brown coal and sand. Triassic and Jurassic: 6 - claystone, siltstone, secondary clay and silt. Adapted from Markiewicz and Winnicki (1997, fig. 8).
Chapter 10 Applied Glaciotectonics Introduction
Glaciotectonic deformation produces locally severe disruption of the normal geology and terrain, and therefore its recognition and structure are important for many human activities. The principal deleterious results of glaciotectonism are disruption and deformation of the stratigraphic sequence, reduction of sediment or rock strength, and major increase in substratum variability. Consideration of glacially thrust terrain is important for mine planning and operation, drift prospecting and mineral exploration, all manner of construction, soils mapping and utilization, and all aspects of ground-water development and protection (JCrgensen 2005). Glaciotectonic fracturing of bedrock presents special problems for mining and civil engineering compared to similar unfractured substratum. At the Heath Steele Mine in northern New Brunswick, for example, glaciotectonic fractures were implicated in failure of part of the mine wall. The mine was opened in a Pb-Zn ore body within porphyry and schist bedrock in the autumn of 1990. In the autumn of 1991, part of the 25-m-deep open-pit wall suffered a series of landslips and began to collapse, and the mine was closed (backfilled) the following year. The wall failure was attributed to a combination of circumstances (Park and Broster 1996). • Glaciotectonic fractures and reactivated joints with clay and sand fillings. • Presence of highly permeable and weathered joints that allowed water penetration. • Increased fluid pressure following several days of heavy rainfall. Glaciotectonic terrain may be expected to be texturally, lithologically, geochemically and geotechnically anomalous compared to the surrounding terrain. The terrain may provide data on the local subsurface stratigraphy, because thrust masses may include sediment and bedrock shoved up from the substratum. The depression upglacier from a thrust mass, if present, may serve as a window into the underlying stratigraphy. Ice-thrust terrain often has thinner drift cover than surrounding terrain, and in some places the bedrock itself may be exposed. Folding may produce inversion of stratigraphy and faulting may lead to repetition or omission of strata, that if unrecognized would cause problems during interpretation of drilling data.
Problems in soil mapping over glaciotectonic terrain originate from the high lateral variability of the deformed substratum. The parent material, which is of major importance for soil mapping, may change repeatedly from sandstone to shale, chalk, till, clay, or gravel over a few 10s of m. If bedrock or other subsurface sediments contain harmful substances, thrusting can bring these materials into the pedogenic zone. For example, bedrock thrust to the surface can increase soil salinity or acidity, because of its geochemistry compared to the normal till cover. One aspect of glaciotectonism often overlooked in applied studies is the depression left at the source of thrusting. This could be important during mineral-exploration and mineplanning phases of resource development. The false assumption that an ore body continues across a glacial depression would yield overly optimistic estimates of total reserves. Sudden discovery of the absence of ore during mining could cause disruption of the mining schedule with potentially serious economic implications. Use of a combination of test drilling and shallow subsurface geophysics between test holes is the most economical way to avoid a surprise discovery of this kind. In some coal mines of western Canada, glaciotectonism has locally removed the coal from areas up to 1 km 2. At one mine, ice thrusting extended deep enough to remove about 12 million m 3 of coal leaving a depression that is partially masked by the cover of later Quaternary sediment (fig. 101). With test holes drilled only at positions A, B, C and D, the interpretation would be that the coal subcrop is continuous. Additional drilling, such as holes E, F and G, would indicate that coal is missing. The transported coal remaining downglacier from the depression is of no use for mining, because the coal exists in small hills close to the surface. This position means the coal has been oxidized, so that its calorific value is reduced greatly and the moisture content increased. The following examples are situations where glaciotectonic disruption of normal bedrock stratigraphy and structure has created serious hazards or economic consequences for human activities. The examples are from large- and small-scale, open-pit mining operations, respectively in Alberta, Canada and Fur, Denmark, and from highway construction through ice-thrust ridges in Saskatchewan, Canada. They illustrate the importance of recognizing the special substratum conditions in glaciotectonic terrain.
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Figure 10-1. A - map showing subcrop of coal in relation to ice thrust depression and transported coal. B - schematic cross section showing the depression f o r m e d by glaciotectonism and downglacier hill including some of the coal transported from the depression. Taken from Aber, Croot and Fenton (1989, figs. 8-1 and 8-2).
Highwall failure, Highvale coal mine, Alberta Highwall failures in glacially thrust bedrock have occurred in several coal mines of western Canada, including the Highvale Mine west of Edmonton, Alberta (fig. 10-2). The mine was owned by TransAlta Utilities Corporation, and production exceeded 11 million tons per year. The adjacent Sundance Power Plant was the largest electrical generating station in western Canada (Tapics 1984). Open-pit mining here involves two basic operations" 1) removal of the overburden above the coal using a dragline
Figure 10-3. Schematic cross section of typical open pit mine showing overburden bench and highwall. Taken from Aber, Croot and Fenton (1989, fig. 8-4). (fig. 10-3), and 2) removal of the coal seam using power shovels (fig. 10-4). The dragline operates from a bench above the coal, removing the overburdern above and below the bench and casting it onto the spoil pile beyond the current mining pit. Structural integrity of the bench and highwall is a crucial factor for successful mining. Glacially thrust bedrock is weaker than undeformed bedrock, and as a result highwalls and benches cut into this material have a greater tendency to fail. The implications of highwall failure are serious; temporary benches excavated in the highwall serve as transportation corridors and as working surfaces for heavy equipment. In addition to the risks posed to men and machines, such failures also result in potentially sizable extra costs for rehandling overburden material, disruption of mining schedules, and outfight loss of minable coal.
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Figure 10-4. Power shovels removing the coal seam at the Highvale Coal Mine, Alberta, Canada. Photo by J.S. Aber (1984). The regional bedrock geology consists of flat-lying, upper Cretaceous and Paleocene, non-marine, coal-bearing strata. The coal belongs to the Ardley Coal Zone of the Scollard Member of the Paskapoo Formation (Carrigy 1970; Irish 1970; Holter, Yurko and Chu 1975). Six distinct and laterally continuous seams are mined. They are separated by shale and bentonite partings, with a cumulative coal thickness of about 10 m. Quaternary strata consist of a discontinuous cover of glaciofluvial and glaciolacustrine sediment over till (Andriashek et al. 1979). Glaciotectonism of the bedrock is extensive, although most disturbed sites are <1 km 2 in area. During the early 1980s, a series of highwall failures occurred in Pit 03 between ramps 3-1 and 3-2 (fig. 10-5). As a result, research was initiated to determine the cause and mechanism of the failures. Geologic data were collected through outcrop study, airphoto interpretation, surface geophysical methods, rotary and auger coring, and downhole geophysical logging.
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Figure 10-6. Geologic cross section A-A'(fig. 10-5) showing deformation. Adapted from Aber, Croot and Fenton (1989, fig. 8-6).
Ground-water data were obtained by setting and monitoring a number of piezometer nests. Much of the following information is summarized from Fenton et al. (1983, 1985, 1986) and Moell et al. (1985). The site stratigraphy is from the top down: till, sandstone, mudstone, interlaminated mudstone, coal and bentonite, and a six-seam coal unit. The till contains abundant clasts and lenses of sheared bedrock material. The underlying sandstone has been subjected to major folding and faulting, although the generally massive nature of this unit commonly prevents recognition of small deformation structures. The mudstone unit is massive and highly fractured with many of the fractures having polished or slickensided surfaces. This unit is also cut by numerous small shear zones. Where exposed in the highwall, each is about 1 mm thick, clay filled, concave upward, and is about 1 m 2 in areal extent. The contact with the overlying sandstone was in most places, when freshly excavated, the site of ground-water discharge. A mass of thrust bedrock overlies till in the western third of the area (fig. 10-5).
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Figure 10-5. Map of southeastern part of Pit 03 showing areas of glacially deformed sediment, sites of structural observations, and locations of geologic and hydrogeologic cross sections. Takenfrom Aber, Croot and Fenton (1989, fig. 8-5).
The glaciers advanced generally from the north, and the bedrock was folded, faulted, and crushed by compression and thrusting. The thrust and shear planes produced during deformation dip generally northward (fig. 10-6). Detailed data on the larger glaciotectonic structures were collected from the three locations where sections are transverse to the structural grain of the area. Averaging the data from each site indicates displacement directions toward 174 ° for pit A, 182 ° for ramp 03, and 160 ° for pit B (fig. 10-5). These directions agree with the average of 170 ° obtained from measurements of smaller shear planes exposed in the fractured mudstone unit. The base of disturbance is a shear zone that rises stratigraphically southward from a position that involves
Aber and Ber
144 deformation within seams i and 2 to a position above seam 1 near the haul road (fig. 10-6). The shear plane is believed to die out at some point south of the haul road. Where the shear plane immediately overlies seam 1, movement was along a 1-m-thick zone of interlaminated mudstone, bentonite and coal and is well illustrated by small folds in coal laminae. H y d r o g e o l o g i c data indicate the glacially deformed overburden is almost completely saturated with ground water and that ground-water flow is directed from south to north toward the highwall (fig. 10-7). However, stratigraphic relationships act to inhibit drainage of this saturated, poorly consolidated bedrock. Surface and subsurface data suggest that highwall failure was composite in nature with four types contributing to the overall result. These four are: 1) block sliding, 2) block rotation, 3) exfoliation or spalling, and 4) in-situ distintegration of large failed blocks (fig. 10-8). Block sliding resulted in a series of vertical fractures marked by cracks, up to 4 m deep. These fractures trended approximately parallel to the highwall face. They developed over the area between ramps 3-1 and 3-2 and extended into the land at least 50 m behind (south of) the upper edge of the highwall. The rotational slumps varied in size, but were generally 50 to 100 m long, roughly semicircular in plan, with direction of rotation approximately northward. One slump, which was observed within 12 hours of failure, appeared to have moved along a pre-existing surface, likely a glacial shear plane. The remaining two failure types were volumetrically less significant. Exfoliation consisted of falling sheets or slices of the sandstone about 1 m to 3 m thick. In-situ disintegration was the gradual disaggregation or reduction in size of sandstone blocks that had accumulated at the base of the slope as a result of the previous types of failure.
The failure cycle began as the dragline started a new cut at the eastern end of the pit and moved westward excavating the overburden bench as it moved. With time, the initial excavation site started to undergo the above types of failures. These progressed along the cut both in the direction of excavation and southward behind the cut, until within a few months the entire bench had been destroyed. At any particular site exfoliation occurred first followed by rotational slumping and block sliding. The rotational slumps, particularly the large ones, were likely the result of remobilization of pre-existing glacial shear planes (fig. 10-8). The block failures were likely caused by elevated pore-water pressures near the contact of the sandstone and fractured mudstone units. The situation was compounded by the saturated and weakened condition of the overlying sandstone. The dip of the sandstone/mudstone contact toward the pit undoubtedly also contributed to block sliding. Vertical fractures were a manifestion of the block sliding and may be the result of the sliding itself, but were more likely due to opening of pre-existing joints or were perhaps a combination of joints and shear planes. The base of failure, near the sandstone-mudstone contact, lies above the basal shear zone of the ice-thrust mass (fig. 10-6) demonstrating that failure took place within the thrust mass. In summary, highwall failure was the product of both the geologic and hydrogeologic regime at the site. Failure resulted from excavation into bedrock that had been crushed, folded, and faulted by glaciotectonism and that was situated in an area of ground-water saturation and elevated pore-water pressure. Orientation of the highwall also contributed to failure by exposing glacially induced shear planes that dipped
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Figure 10-8. Schematic drawing showing observational data and model for highwall failure. Takenfrom Aber, Croot and Fenton (1989, fig. 8-8).
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into the pit. Hydrologic conditions were caused by minimal hydraulic conductivity together with the northward hydraulic gradient into the mine. Foreknowledge about location, structure, dimensions, orientation, and hydrology of glaciotectonically disrupted terrain is important to both mine planning and management.
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Local geology consists of nearly flat-lying upper Cretaceous formations overlain by Quaternary strata (fig. 10-10). Upper Cretaceous bedrock is comprised of weakly consolidated sand,
5
MAYMONT
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Maymont is located in the North Saskatchewan River valley between the Thickwood Hills and the Eagle Hills uplands (fig. 10-9). The uplands exceed 600 m elevation, and the valley floor lies about 450 m elevation; total topographic relief is >200 m. During periods of maximum glaciation, ice movement over the region was generally from northeast to southwest. However during early and late phases of glaciation, when the ice was relatively thin, the Battleford ice lobe flowed toward the southeast within the valley.
4
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Highway construction, Maymont, Saskatchewan Recognition, structure, and composition of glaciotectonic terrain are important for geotechnical site investigations for highways, bridges, dams, canals, and related constructions. The following example is from the Maymont area of Saskatchewan (fig. 10-2). Construction began in 1973 on Highway 376 and a bridge to cross the North Saskatchewan River valley. When highway excavation was nearing completion, a massive failure occurred along one of the road embankments. Initial investigations indicated the failure was due to low shear strength of the glacially deformed bedrock (Sauer 1978; Krahn et al. 1979). A wide variety of geological and geophysical methods were employed for continuing, detailed study of surficial and subsurface conditions surrounding the site. On this basis, a thorough understanding of the glaciotectonic setting for the Maymont vicinity has emerged (Stauffer, Gendzwill and Sauer 1990).
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Figure 10-10. Section across the North Saskatchewan River valley showing main geologic units. Bearpaw, Judith River and Lea Park formations are all upper Cretaceous. A reconstruction of the Battleford ice lobe is depicted within the valley. Note large vertical exaggeration; adapted from Stauffer, Gendzwill and Sauer (1990, fig. 3).
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The upper gouge zone (465 m) is near the top of the Lea Park Formation. It forms the d6collement over which Judith River Formation was dislocated and thrust into ridges during a late phase readvance by the Battleford ice lobe. The lower two gouge zones mark a disturbed interval (438-433 m) that also was created by glacial shearing. Both upper and lower gouge zones were controlling surfaces for later landslide slippage. Stauffer, Gendzwill and Sauer (1990) identified two main landslide masses, each extending ca. 10 km along the river valley (fig. 10-13).
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• Upper landslide mass is about 1300 m wide, 30 m thick, and rests on the upper gouge zone (465 m). It probably was initiated first when the fiver eroded its valley below the gouge zone level or when the Battleford ice lobe retreated from the valley. It may still be active today, as suggested by open fissures on its slope.
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• Lower landslide mass is about 1000 m wide, 30 m thick, and rests on the lower gouge zone (438-433 m). It was activited second, about 11,000 radiocarbon years ago, but its movement ceased by 6800 years ago based on its relationship to dated alluvium in the fiver valley.
Figure 10-11. Log from test hole 81 (see next fig. for location) showing the interval in which three soft gouge zones are situated. W = natural water content, and W1 = liquid limit in water content as percentage of dry weight. Taken from Stauffer, Gendzwill and Sauer (1990, fig. 10).
Failure of the highway embankment happened near the northern margin of the ice-thrust ridge terrain (fig. 10-14) and was described in detail by Krahn et al. (1979). Site
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Figure 10-12. Cross section and well logs along Highway 376 approach to the bridge over the North Saskatchewan River at Maymont. Upper (465 m) and lower (438-433 m) gouge zones mark the basal surfaces for landslide slippage toward the river valley. Note near horizontal placement of the gouge zones and intertonguing of landslide debris with Holocene alluvium of the river valley (to right). Adapted from Stauffer, Gendzwill and Sauer (1990, fig. 13).
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geology consists of interbedded sandstone, siltstone and shale overlying a zone of highly brecciated and slickensided shale (fig. 10-15). The slickensided zone on the eastern side of the road cut dips west, toward the highway. Glaciotectonism deformed the bedrock to a depth of 45 m. Factors contributing to the failure were low shear strength of the disturbed bedrock and dip of the slickensided zone toward the excavation.
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Figure 10-13. Block model of structures associated with ice thrusting and later landslides at Maymont, Saskatchewan. JR = glaciotectonically disrupted material of Judith River Formation. LP = Lea Park Formation containing gouge zones of bentonitic clay beds weakened by glacial shearing. Upper landslide (circles) contains Judith River debris; lower landslide (triangles) includes Lea Park and Judith River debris. Adapted from Stauffer, Gendzwill and Sauer (1990, fig. 15).
The mo-clay strata are formally designated as the Fur Formation (Pedersen and Surlyk 1983) with age established as early Eocene. Diatomite of the Fur Formation is highly porous with a low density (0.8 g/cc) consisting of marine diatoms (opal), clay minerals (smectite) and volcanic dust (Pedersen et al. 2004). Dark-gray to black ash layers are for the most part basaltic tephras of glass and mineral particles. The ash layers are conformably interbedded with diatomite and contrast sharply with the light-colored diatomite layers. Individual ash layers are continuous and nearly constant in thickness throughout the Fur Formation, allowing for a detailed tephrachronology. The positive series of ash layers (+ 1 to + 140) is included in the Silstrup Member, and the : y !+~!~i!i~Ni!oi~i~i!~+i+~?i!~i~i~!i?!i~i~i!~!~!~+!~i!~i~i~i++~i!~:i~i~:+i negative series (-1 to -39) is within the Knudeklint Member (fig. 10-16). Gry (1965) subdivided the formation on the
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The classic study of glaciotectonic disturbance of mo-clay was published by Gry (1940, 1979), who applied methods of structural geology for analyzing the ice-push deformations. Major structures are large, rootless folds of mo-clay with or without deformed glacial strata. Disturbed mo-clay is present in composite ridges that make up the northern third of Fur (fig. 10-17). The ridges reach a maximum height of 75 m and define a gentle arc, concave northward. Close agreement exists between structural and topographic trends along the arc, which suggests the composite ridges were deformed by a northerly ice-lobe advance coming across the Limfjord basin.
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basis of gross lithology and industrial usage into four informal series. • Upper series - almost ash-free diatomite with thin ash layers (+ 119 to + 140), minor industrial use, about 8 m. • Ash series - thin diatomite layers interbedded with many, thick ash layers (+1 to +118), no industrial use, about 17 m.
The unusual properties of mo-clay have several industrial applications, so it has been quarried for many decades at various locations, principally on northern Mors and northern Fur. Primary use is for ceramic products. Structural disturbance of mo-clay is the main concern for planning and developing quarries. Quarries generally follow the strike of folded strata until those strata terminate against a fault or the overburdern becomes too thick.
Difficulty was encountered in excavation at the ManhCj quarry (fig. 10-18), during the early 1980s, as a result of irregular pockets of sand and gravel included within the mo-clay. This sand and gravel contaminates the mo-clay and diminishes its quality as a raw material, which has significance for ultimate extraction of the estimated mo-clay reserve and for day-to-day mining economics. The Geological Survey of Denmark was commissioned to investigate the problem and make recommendations to the quarry operators (Pedersen and Petersen 1985). The middle portion of the Fur Formation, totaling about 35 m in thickness from just above ash + 19 to the upper part of the claystone series, is exposed in the quarry. The portion between ash layers -13 to -19 is the quarry interval (see fig. 10-16). Normal thickness of the quarry interval is increased at this location as a result of reverse faulting. The overall structure consists of folds trending NW-SE with a culmination near the center of the quarry. Fold axes plunge about 20 ° SE in the eastern part of the quarry, whereas folds plunge about 15 ° NW in the northwestern part. This corresponds closely to the original land topography. Glaciofluvial sediment, consisting of sand and pebble- to cobble-sized gravel, fills channels cut into the folded moclay and in places contains reworked masses of brecciated mo-clay. The channels follow syncline troughs and are cut
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Figure 10-19. Portion of north-south, test-trench wall about 60 m west of ManhCj quarry. Note north-dipping sand wedges and gravel-filled channel. Modified from Pedersen (1986, fig. 3); copyright Geological Survey of Denmark and Greenland. down nearly to ash layer-11. Sand and gravel also forms steeply inclined wedges that extend to 10 m depth within the mo-clay (fig. 10-19). The wedges are up to 30 cm wide near their tops and become narrower downward to only a few cm width before pinching out into irregular fingers. Internal zonation of well-preserved wedges resembles that of ice wedges formed in connection with permafrost. The development of glaciofluvial channels and wedges was intimately related to folding of the mo-clay during glaciation (fig. 10-20). During the initial stage, anticlines and synclines were folded within mo-clay in front of advancing ice. Folding took place above a d6collement zone located at or near the base of the claystone series. Melt-water channels followed
syncline troughs in the glacial foreland, and fissures opened on anticline crests. Anticlines became overturned with continued deformation, and sediment filling of the open fissures with glaciofluvial sand and gravel took place. Thrusts developed along anticline axial planes and cut up into channel fills. The crests of
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Figure 10-22. Section of deformed mo-clay exposed in the quarry at Hesteggtrden, western Fur, Denmark. A presumed basal dgcollement is projected below the section, and discordant till caps the folded and thrust mo-clay (diatomite). Adapted from Pedersen (2000, fig. 3). anticlines collapsed along normal faults into the channels. In this manner, fissures were filled with sediment and were extended or reoriented by continued faulting and folding. Ultimately ice advance overran the area, truncated the folds, and deposited a till cover. Close relationship is evident between fold orientation and positions of wedge structures (fig. 10-21). The wedges fall in two dominant sets: 1) approximately 110-290 °, parallel to fold axes, and 2) roughly 60-240 °, oblique to the folds. Set 1 represents tension fractures formed on anticline crests, essentially normal to the direction of maximum compression (= local ice movement) from about 20 °. Pedersen and Petersen (1985) interpreted set 2 wedges as extension fractures; however, true extension fractures should be oriented parallel to maximum compression (Hobbs, Means and Williams 1976, fig. 7.31). Set 2 wedges are oriented 40 ° from maximum compression, an angle typical of shear fractures. A third, lesser wedge set is found at about 165345 ° . This is 35 ° from maximum compression and represents conjugate shear fractures in relation to set 2. Finally, there are some wedges oriented in various other directions, which Pedersen and Petersen (1985) thought were inherited from a permafrost polygonal pattern. Whether this permafrost was active or relict at the time of ice shoving is uncertain.
A newer quarry opened nearby at Hestegirden in 1996 displays intense imbricate thrusting and steeply dipping beds of mo-clay (fig. 10-22). In this case, the claystone series of the Knudeklint Member is exploited, mainly between ash beds-33 and-19. The southern portion of the section contains tight to near-isoclinal folds, and the northern part includes multiple thrust faults. Fold axial planes and fault planes dip steeply toward the north. These structures were created by the Norwegian ice advance coming directly from the north; the discordant till contains indicator erratics from the Oslo region. In addition, pre-existing normal faults of presumed neotectonic origin were reoriented during glaciotectonic disturbance of the mo-clay. As these examples from Fur demonstrate, intimate knowledge of glaciotectonic structures would be useful for planning future quarry operations. For example, development of sand wedges is largely a superficial phenomenon that extends no deeper than 10 m into the mo-clay. The distribution of wedges is closely controlled by fold structures, and so zones of sand and gravel contamination within the quarry interval may be predicted and avoided. Careful preliminary studies involving many test pits are necessary to properly evaluate quality and quantity of mo-clay reserves in ice-shoved terrain of the western Limfjord district.
This Page Intentionally Left Blank
Chapter 11 Distribution of Glaciotectonism Continental distribution of glaciotectonic phenomena Introduction m Glaciotectonic features were considered
unusual or rare glacial phenomena when first recognized more than a century ago. This point of view persisted for a long time, because the documented examples of glaciotectonic landforms and structures were few and far between. Recognition and description of glaciotectonic features became increasingly common beginning in the 1950s, leading Sauer (1978) to conclude that such features are probably widespread phenomena in the outer portions of glaciated regions. Our knowledge concerning the central regions of ice-sheet glaciations has developed more slowly. Fewer investigations of glacial geology were carried out partly for logistical reasons and because many geologists assumed that, aside from striations and scattered erratics, not much of interest could be found (Goldthwait 1971). This point of view must be modified substantially, based on systematic description and mapping of glaciotectonic phenomena during the past quarter century in Europe and North America. Vast areas hitherto poorly known, particularly Arctic regions of Canada, Alaska and Russia, have revealed surprising conditions of former glaciation. Recognition of glaciotectonic features has extended offshore onto the continental shelves as well. In the zone beneath the ice divide at the centers of former ice sheets, significant drift containing various typical glaciotectonic features is preserved. Thus, a general model for the continent-scale distribution of glaciotectonic phenomena includes three primary zones (Dyke and Prest 1987a; Aber and Lundqvist 1988; Aber, Croot and Fenton 1989). • Outer z o n e - all manner of large and small glaciotectonic phenomena in drift and soft sedimentary bedrock both onshore and offshore. • I n t e r m e d i a t e zone small, isolated glaciotectonic features found mainly in locally thick drift of the last glaciation. • Inner z o n e widespread, small- and moderate-sized glaciotectonic features developed in older drift. The zonal model is of course highly generalized; development of each zone may vary according to local geological circumstances. The boundaries between zones are sharp in some cases and transitional in other areas. Individual zones are not always fully developed or continuous around the whole area of glaciation. The three zones represent the cumulative results of multiple ice-sheet glaciations during the Pleistocene. The model is, thus, concerned with the general or overall
distribution of glaciotectonic phenomena, not their local presence or absence; factors related to local distribution will be addressed in the following sections. It should be noted also that basement faults and seismic zones are widely distributed throughout regions of former ice-sheet glaciation. They are more common in the intermediate and inner landscape zones, where the impact of lithospheric depression and rebound is greater than for the outer zone. The three-zone model can be demonstrated best for the last major glaciation: Laurentide Ice Sheet (Wisconsin) in North America (fig. 11-1) and F e n n o s c a n d i a n Ice Sheet (Weichselian) in northern Europe (fig. 11-2). The model applies to earlier glaciations as well, but glaciotectonic features related to earlier glaciations are preserved mainly in the outer zones, generally beyond the limits of Wisconsin/ Weichselian glaciation. Older glaciotectonic features may also be found in the inner zones of glaciation. The model applies to other ice sheets and mountain glacier complexes, such as the Cordilleran sector of North America and the Barents/Kara Sea portions of northern Eurasia. O u t e r zone - - The outer glaciotectonic zone in North America includes the Atlantic Coastal Plain and adjacent continental shelf of southern New England (Sirkin 1980), the Great Plains of the mid-continent beginning in the central United States and stretching northward across the Canadian Prairie (Moran et al. 1980), and the Arctic continental shelf of the Yukon (Mackay 1959). The outer zone reaches still farther west to the Bering Sea coastal zone of northwestern Alaska (Huston, Brigham-Grette and Hopkins 1990) and easternmost Siberia (Benson 1993).
The outer zone is underlain predominately by soft, poorly consolidated Mesozoic and Cenozoic sedimentary strata consisting mainly of Cretaceous or Paleogene bedrock of clastic composition. Thick and nearly continuous Quaternary strata mantle the bedrock. Large looped end moraines, drumlin fields, older drift, and multiple till sequences are common. All kinds of large and small glaciotectonic features are abundant, and many classic end moraines are now interpreted as ice-shoved features (Moran 1971; Oldale and O'Hara 1984). In Europe, the outer glaciotectonic zone begins with Ireland (Thomas and Summers 1984) and England on the west and extends eastward across the southern Baltic and Central European Plain. Various glaciotectonic phenomena are abundant in the Netherlands, Denmark, and northern Germany. The continental shelf is also included for the North Sea (Huuse and Lykke-Andersen (2000), Norwegian Sea
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(S~ettem 1990), and Barents Sea (Samem 1994). Across Poland, glaciotectonic features are likewise extremely common (Ruszczyfiska-Szenajch 1985; Ber and Krzyszkowski 2004), including many examples of buried forms (fig. 11-3). Glaciotectonic structures and landforms continue eastward in Lithuania, Latvia and Estonia as well as Belarus and the Ukraine (Matoshko 1995; Chugunny and Matoshko 1995). The outer zone continues farther to the northeast across Arctic Russia (Henriksen et al. 2001; Astakhov 2004), including the Pechora Sea continental shelf (Gataullin, Mangerud and Svendsen 2001). This outer zone is related to glaciation by the Barents Sea ice sheet (fig. 11-4). Eastward of the Urals, the outer zone stretches across the northern portions of both
West Siberia (Astakhov, Kaplyanskaya and Tarnogradsky 1996) and East Siberia (Grosswald et al. 1992). The outer glaciotectonic zone of Siberia is connected to the Kara Sea ice sheet. The outer glaciotectonic zone of Europe, as in North America, is underlain mainly by relatively soft Mesozoic and Cenozoic sedimentary strata. Cretaceous chalk and Paleogene and Neogene bedrock of clastic (clayey) composition are frequently deformed along with Quaternary clay-rich sediments (Jakobsen 2003). Thick drift, multiple till sequences, drumlins, and large end moraines are characteristic of this zone. Many of these end moraines have a glaciotectonic genesis (Ber 1987; van Gijssel 1987; Meyer 1987; van der Wateren 1987).
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The intermediate zone of glaciation presents a startling contrast to the outer zone. In Europe, the intermediate zone includes the Fennoscandian Shield of southern Sweden and Finland and the Caledonian Mountains of Norway and western Sweden (fig. 11-2). Drift is much thinner and discontinuous. Hard, mostly crystalline bedrock is exposed over large areas. Locally thick drift was left in scattered end moraines and eskers during the last deglaciation, but older drift is only rarely preserved in small, protected sites. Such strata do contain various small (<10 m Intermediate
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In both North America and Europe, the outer and intermediate zones are separated by transitional belts of varying width and lateral continuity. These transitional belts are underlain by consolidated Paleozoic sedimentary bedrock. In North America, this belt begins south of Lake Ontario and stretches westward across the Great Lakes into the Central Lowlands. The belt exhibits gradational features between the outer and intermediate zones. The inner portion shows evidence of strong glacial erosion with relatively thin drift and prominent drumlin fields, as in western New York (see fig. 9-10), but glaciotectonic features are rare (Andrews 1980). Toward the outer portion, drift becomes thicker, end moraines are common, and scattered glaciotectonic features are also more abundant (Barbour 1913; Howe 1968; Rieck et al. 1991).
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Figure 11-4. Generalized limits of Weichselian ice sheets in northern Russia and Scandinavia. Early~middle Weichselian glaciation was much more extensive to the east and south of Novaya Zemlya than was the late Weichselian ice sheet. Adapted from Gataullin, Mangerud and Svendsen (2001, fig. 1); reprinted with permission from Elsevier Hampshire (Stone and Koteff 1979). Small glaciotectonic features are present on Magdalen Islands in the Gulf of St. Lawrence and on mainland Nova Scotia, where slate, Permian
The northern limit of the outer glacial zone in Scandinavia is sharply marked in southernmost Sweden by a major structural lineament. Immediately to its north, a mosaic of crystalline and sedimentary bedrock is found in the Fennoscandian Border Zone, which together with Baltic islands of Bornholm, 01and and Gotland make up a transitional belt. Small glaciotectonic structures are present in drift on Oland (K6nigsson and Linde 1977) and in weakly consolidated Jurassic sandstone on Bornholm (fig. 11-5). Another well-marked transition zone is found in Estonia (EE, fig. 11-2), which is underlain by lower and middle
Distribution of glaciotectonism
157
Figure 11-5. Small thrust fault with drag fold in Jurassic sandstone (below) and bouldery glacial gravel at GallOkken, island of Bornholm, Denmark. Photo by J.S. Aber (1979). Paleozoic sedimentary strata. To the north and west, Finland and Sweden were subjected to strong glacial erosion that stripped away soft sedimentary strata to expose the Fennoscandian Shield. To the south and east, thick glacial sediments accumulated in Latvia, Lithuania, Russia, and Belarus. This transition is reflected in the glacial sediments of Estonia. The northern half of the country has generally less than 5 m of glacial cover, while glacial sediments more than 100 m thick are found in the southeast (Raukas and Kajak 1997). Various glaciotectonic structures, including bedrock rafts, ice-shoved ridges, and internally deformed drumlins, are common in Estonia (Rattas and Kalm 2004). Inner zone - - The inner glaciotectonic zone of Scandinavia extends from south-central Norway, across central and northern Sweden, into northern Finland, and continues eastward an uncertain distance onto the Russian Kola Peninsula (fig. 11-2). Drift is again moderately thick and continuous with drumlins and Rogen moraine common. Exposures of Precambrian and Paleozoic bedrock make up only a small part of the landscape. Interstadial and interglacial sediments are found in many places below till of the last glaciation, and multiple Weichselian till sequences are known (Lundqvist 1967; Ber and Kujansuu 1974; Hirvas, Korpela and Kujansuu 1981; Haldorsen and SCrensen 1987).
Glaciotectonic structures are common within such overridden deposits (Vorren 1979; Lundqvist 1985), and even small iceshoved hills are found in central Sweden (fig. 11-6). It is clear from the geological context that many of these glaciotectonic features were created in older Quaternary strata beneath the thick center of the Fennoscandian Ice Sheet (Aber and Lundqvist 1988). The deformation of overridden sediments occurred during all phases of the last glaciation: advancing, maximal, and recessional. Due to migration of the Weichselian ice divide, local directions of ice pushing
Figure 11-6. Glaciotectonic structures in pre-late Weichselian stratified drift near Ostersund, central Sweden. A - thrust fault with drag folds at Grytan; field of view ~10 m across. B - overturned and thrust core of a small cupola hill on AndersOn island; shovel for scale (lower right). Photos courtesy of J. Lundqvist. shifted greatly during the course of glaciation. It is also possible that some glaciotectonic features were formed in ice marginal positions during the final phase of deglaciation. Two inner zones are present in North America corresponding to the two main sectors of the Laurentide Ice Sheet (fig. 111). Drift cover is nearly continuous with till plains, Rogen moraine, and drumlin fields in many portions. Precambrian crystalline and sedimentary bedrock is rarely exposed at the surface. These two zones correspond to the major ice divides of the Laurentide Ice Sheet (Prest 1983). Unfortunately, little is known about the subsurface structure of the glacial sediments. Based on the central Swedish situation, glaciotectonic phenomena may be relatively common and probably will be discovered with more field work in the inner zones of northern Canada. A general impression exists that the inner zones in the regions of ice divides were stagnant areas in which the Fennoscandian and Laurentide ice sheets had little effect on the landscape,
Aber and Ber
158 aside from crustal depression and rebound (Sugden 1976). A few geologists conversely have suggested that deep crustal erosion of crystalline basement rocks took place beneath the ice-sheet centers (White 1972). Both points of view appear extreme. It is now apparent that moderate glacial erosion, deposition, and deformation all took place within the inner zones under dynamic conditions of ice movement.
Continental glacial landscape m o d e l - The distribution, types and sizes of glaciotectonic phenomena are related in a general way to availability of erodible or deformable strata, namely thick Quatemary deposits or soft sedimentary bedrock. Within each zone the local presence or absence of glaciotectonic features may reflect variations in glacier dynamics, permafrost, ground water, bedrock topography, lithology of the substratum, etc. The three zones are the results of multiple Pleistocene glaciations and represent long-term, cumulative modifications of continental substratum by ice sheets. Complex patterns of ice movement exist, particularly in the inner and intermediate zones of glaciation, due to changing ice centers and flow lines (Boulton and Clark 1990a). Thus, the landforms of older glacial movements may be overprinted, much modified, or completely obliterated by later ice flows. The results are cross-
SO0 .............
cutting geomorphic pattems over regions of former glaciation (fig. 11-7). The three glacial landscape zones may well be related to thermal regimes developed at the base of the ice sheets, but opposite opinions have been expressed on this issue. In North America, for example, Sugden (1978) envisioned the inner and outer zones as thawed-bed regimes, while the intermediate zone was cold-based ice during Laurentide glacial maxima (fig. 11-8). The pervasive erosion evident throughout the intermediate zone is hardly compatible with a frozen glacier sole, however. In contrast, Sollid and SCrbel (1988) interpreted the inner and outer zones as cold based, while the intermediate zone was at the pressure-melting point in Scandinavia during deglaciation (fig. 11-9). This viewpoint seems more compatible with the geomorphic evidence. The glacial landscape zones in North America and northern Europe are arranged in symmetrical patterns with respect to the Atlantic Ocean (fig. 11-10). From the continental shelves to interior mountain systems, a series of glacial geomorphic zones are found in the same general sequence over distances of several 1000 km. Glaciotectonic phenomena are found primarily in the outer (zone 1) regions toward the periphery
km -.= . . . . . . . . . .
Figure 11-7. Pattern of glacial lineations over central Canada as interpreted from Landsat imagery. Note complicated crosscutting relationships, especially in the Quebec and Keewatin sectors. Adapted from Clark (1993, fig. 14b). Copyright John Wiley and Sons Ltd. Reproduced with permission.
Distribution of glaciotectonism
159
~
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d97& fig. 9).
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t .,~~..~~
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---
Late
----
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Weichselian Oryas
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Figure 11-9. Possible thermal regimes of the last Fennoscandian ice sheet during its deglaciation. 1 = cold marginal zone, 2 = warm (pressure-melting) intermediate zone, 3 = cold interior zone. Adapted from Sollid and SOrbel (1988, fig. 5). Reproduced from Boreas by permssion of Taylor & Francis AS.
of this continental glacial landscape model. Toward the center, zone 3 represents a secondary landscape setting for glaciotectonic features. Finally zone 2 displays the least amount of glaciotectonism.
NORTH AMERICA {E ~w) ATLANTIC
APPALACHIAN
MTS,
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LANDSCAPE
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LAKE O R FJORD TROUGH
OEEP
Norizor~o i Scole Approximate Verfico~ Stole Adjus:~ed For Depressio~ Lo;rge Ve~ticQ! E~ggerotion
Figure 11-10. Model of glacial landscape symmetry for North America (east-west) and Europe (west-east) with schematic icesheet profile and substratum lithology. Pie diagrams represent local, relative contributions of glacial erosion, deformation, and deposition for modifying the three landscape zones; overall magnitude of glacial landscape modification increases toward the periphery. Modified from Aber (1992, fig. 14).
160
Aber and Ber
This continental distribution of glaciotectonism reflects the first-order occurrence of relatively weak substratum that could be deformed by glacier loading and movement. The general correspondence in North American and European glacial landscapes is most remarkable considering the geographic and climatic differences between the two continents. Repeated growth and decay of ice sheets has imposed a common geomorphic pattern, in spite of these differences, which represents a mature end-product of multiple Pleistocene glaciations.
Regional patterns of glaciotectonism The regional distribution of glaciotectonic features is considered primarily for the outer zone, in which such features are abundant and their pattern of distribution may be compared with other glacial phenomena. Various factors have been considered important for controlling the location and distribution of glaciotectonic features (Table 11-1). These factors vary greatly in time and space, which gives rise to a rich variety of glaciotectonic landforms and structures. Two basic regional distribution patterns are recognized for glaciotectonic features. • Sporatic distribution of megablocks, rafts, diapirs, and other features that have little or no morphologic expressions, along with small cupola hills and drumlins.
These features presumably were created in subglacial settings well behind ice margins. • Ice-marginal position of morphologically prominent hillhole pairs, composite ridges, and large cupola hills. These features were created at or near active ice margins, and their locations are closely related to development of ice lobes or tongues. The fundamental geometry of ice lobes was recognized in the late 19th century (fig. 11-11). Horberg and Anderson (1956) noted three main factors that controlled the form and extent of Pleistocene ice lobes in the north-central United States: 1) preglacial (bedrock) topography, 2) configuration of the ice sheet, and 3) deflections by adjacent ice lobes. Of these, the first was undoubtedly most important. Ice-shoved hills are arrayed in belts from 2-3 km to 3-5 km wide immediately at or inside ice margins. These belts define the positions of larger ice lobes. Individual hills within a belt often display their own looped shapes that reflect development of smaller ice tongues related to local valleys or aquifers. Behind these belts, the glaciotectonic hills become generally smaller and smoothed, giving way upglacier to streamlined terrain of low relief.
Table 11-1. Factors considered important for genesis of glaciotectonic phenomena. Compiled from many sources. Giaciotectonic factors
Subglacial
Proglacial
S
P
Lateral pressure gradient :::::::::::::::::::::::
:~=:
:, ,,,.,,,.,-,,--:
. . . . . . .
,.,..
:: .................
==:
......................
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: .................
................................
:::
......................
S
' .......................
tce advance over permafrost
,. . . . . . . . . . . . . . . . .
i""
..........
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,,,, ...................................
.....................................
,L,
S
P
S , . . . . . . . . . . . . . . . . .
P
S
P
Lithotogic boundaries in substratum
S
P
S
P
=:~:
......................
:: ..............................
=:.
Surging of ice [.obes I nterlobate position
:
Damming of prog!aciai lakes
,
S
S ~ a r i ~ fault btocks up into ice
S
Sq~e~
S
,, .........
,, .............
:into catty ....................
...........................
P
P
~ m p r e s s i ~ f!ow with basai drag
...........
, .....................
Subgtaciai meltwater erosion
0 0
3
l
P
T h r ~ t i ~ in front of ice
===
I
,,. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
ice adva.~e o ~ r buried ~rd-rock obstac!e
.................................................................................
52
I
:,,
. . . . . . . . . .
.................................................................................
Ice adva~e :against topographic obstacle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3
o
P
• '" . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2
...................
Excess: pore-water pressure (confi~d aquifer) ..............................
!
Km
Figure 11-11. Schematic map showing three phases of an ice lobe. 1 = outer, 2 = intermediate, 3 = inner ice-margin positions. Divergent flow lines shown for each phase. Based on Chamberlin (1886, fig. 25).
,
S
...........~ ........... ~=:..... 1 0 0 - 2 0 0
:.
................................
Distribution of glaciotectonism
161
,
Eskers are often associated with ice-shoved hills, for example at the small hill-hole pair near Anamoose, North Dakota (fig. 11-14). The 30-m-high hill is located immediately southeast of the source depression at Steele Lake. Steele Lake is situated on the margin of a partly buried melt-water trench, which contains a sand-and-gravel aquifer. A small esker begins at the edge of Steele Lake and extends along the northern flank of the ice-shoved hill. The esker supposedly was deposited by ground water flowing from the aquifer, when thrusting opened the downglacier edge of the aquifer. According to Bluemle and Clayton (1984, p. 285):
i
.
....~!~:,.
The formation of features like those found at Anamoose might be likened to popping the cork from a bottle of champagne; after initial release, the pressure is dissipated. The pressure in the bottle (aquifer) is released as the cork (hill) is removed from the bottle (depression).
BE N.SON WELLS
~ 0
Figure 11-12. Map of glacial features associated with the Martin ice margin in central North Dakota. Note icemarginal position of ice-thrust masses and subglacial placement of drumlins. Position of Anamoose indicated by asterisk. Adapted from Bluemle and Clayton (1984, fig. 21). Reproduced from Boreas by permssion of Taylor & Francis. Moran et al. (1980) and Bluemle and Clayton (1984) interpreted this basic pattern in North Dakota as simultaneous creation of streamlined and ice-thrust features beneath and behind a near-stationary ice margin (fig. 11-12). Thrusting presumably took place in a narrow frozen-bed zone, while streamlining occurred upglacier under thawed-bed conditions. Individual thrust blocks were "plucked" up by the ice. Permafrost was probably involved in many situations, but thrusting of unfrozen material also could occur, particularly above confined aquifers (fig. 11-13). In places where the ice lobe continued to advance, earlier ice-shoved hills were subjected to later streamlining.
,,~
.....................
i -
~,. ...........
5":
:.
COHESIVE "-ST R AT
. . . . . . . . . . . .
The distribution of ice-shoved hills and their relationship to other glacial features in North Dakota is seemingly consistent with genesis beneath thin, stationary margins of ice lobes.
ANAMOOSE
........4~
[ .......................................................................~ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
I
B
GL ACIER
"~",~
~ ; ' - " '':'~=~"
Figure 11-13. Schematic model for subglacial plucking of large cohesive block above distal end of a confined aquifer with excess ground-water pressure. Small arrows show direction of ground-water flow. Adapted from Moran (1971, fig. 5).
Figure 11-14. A - Sketch map of hill-hole pair and esker at Anamoose, North Dakota. For location, see Fig. 11-12; adapted from Bluemle and Clayton (1984, fig. 7). Reproduced from Boreas by permssionof Taylor & Francis. B - ground view over Steele Lake depression in the foreground with the ice-shoved hill on the horizon. Photo by J.S. Aber (1986).
Aber and Ber
162 This model is based primarily on the morphologic association of features, because little stratigraphic evidence is available for demonstrating the age or sequence of events. It is equally plausible that ice shoving took place during glacier advance and that the hills were subsequently overridden by ice (Carlson and Freers 1975). The fresh and uneroded appearance of many of the ice-shoved hills on the Great Plains indicates they were formed late in the deglaciation near the margin of active ice lobes and that subsequent ice movement was short or weak enough to prevent any remolding of the hills. The lobe-marginal distribution of ice-shoved hills in certain other situations has been interpreted as the result of proglacial thrusting during ice-lobe advance. This pattern of ice pushing is demonstrated best by those cases where ice did not override the hills subsequent to thrusting, for example the Dirt Hills, Saskatchewan and Utrecht Ridge, the Netherlands (chap. 5). In other cases, proglacially thrust hills were later modified by overriding ice, as at the Cactus Hills, Saskatchewan (chap. 5) and Aquinnah, Massachusetts (chap. 6). Generally the development of ice lobes, and thus locations of ice-shoved hills, was controlled by preglacial bedrock topography. Among the well-known examples is the western Baltic region (fig. 11-15). During late phases of each major glacial cycle, the Baltic depression guided ice streams that advanced across southern Denmark and northern Germany (Ehlers 1990). These ice streams originated in the northern Baltic region and flowed toward the south and southwest.
~:::~i~i',, v.,.~:~: ~;% ........................ ...~:~ ....................... -:i-..~,...
..............~:............................................ ..... ~:"............................ ~:.., ~:~,:~,, :~
Upon encountering the Pomeranian Upland of Poland, ice flow turned westward and passed through the opening between the islands of Bornholm and Rtigen (Stephan 2001 a). From there, ice flow fanned out toward the southwest, west, northwest and north. Permafrozen uplands to the south may have been another factor that diverted ice movement toward the west and northwest along lines of lesser basal resistance (Piotrowski 1993). Glaciotectonic disturbances associated with the Baltic ice lobe are particularly common in southern Denmark, northern Germany, and northwestern Poland. However, bedrock topography was not a factor in the Netherlands, where the Saalian ice sheet advanced over a relatively flat alluvial plain. Nonetheless, composite ridges display a well-developed lobate distribution around deep glacial basins that are now partly filled with younger sediment (fig. 11-16). The creation of these basins was the result of subglacial melt-water erosion combined with glaciotectonic thrusting (van den Berg and Beets 1987). Another important factor related to regional distribution of glaciotectonism merits discussion. Ice-shoved hills are frequently found where relatively soft surficial sediments thin or pinch out above more-resistant basement rock. This situation is particularly common in Poland (RuszczyfiskaSzenajch 1978, 1993). For example, Trzebnica Hills are located above a buffed horst in Triassic basement, over which Neogene sediments are much thinner than to the north (fig. 11-17). Ice advance thrust and stacked the softer strata material exactly at the northern edge of the horst. A similar situation is encountered in the easternmost Netherlands, where Neogene deposits come near the surface. Ice-pushed ridges are found at the position where Pleistocene sediments begin to pinch out (fig. 11-18).
Model for lobate pattern of glaciotectonism
. . . . . . . .
~
....... .,,,, ......... :
........
:::::: 2t~ .......................................
:,,,,:. ........ ..~.j
~:09tkm,
Figure 11-15. Flow of the Baltic ice lobe, late Weichselian glaciation in the western Baltic region. Two phases are shown: a) East Jylland-Sehberg-Mecklenburg advance, and b) Bcelthav-Warleberg readvance. Small arrows indicate local ice-push directions; short lines depict ice movement based on striations or till fabric. B = island of Bornholm, Denmark; R = island of Riigen, Germany. Adapted from Stephan (2001a, fig. 1).
A model for glaciotectonic deformation associated with ice lobes consists of two stages: 1) proglacial thrusting of iceshoved hills followed by 2) subglacial modification of overridden hills (fig. 11-19). Initial thrusting and shoving up of a composite ridge take place in front of the ice and above a detachment horizon or d6collement. The d6collement may be controlled by several features: lower boundary of permafrost, lithologic boundary, position of confined aquifer, etc. High ground-water pressure is presumably developed along the d6collement. Subglacial melt water may either erode tunnel valleys or deposit eskers, while proglacial melt water may erode spillways across the ice-pushed ridges and deposit outwash sediment on the distal side of the hill. Small and temporary lake basins may accumulate sediment between the ice and the hill or between individual ridges. Supraglacial debris moving down the ice front may be incorporated into the ice-
Distribution of glaciotectonism
163
Figure 11-16. Map of the Netherlands showing major glacial basins associated with ice-shoved ridge complexes. The western, southern, and eastern margins of these basis are marked by ice-pushed ridges composed largely of unconsolidated Pleistocene strata. Adpated from de Gans, de Groot and Zwaan (1987, fig. 1).
...... -,=
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: ...........
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.....
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Figure 11-17. Cross section of Trzebnica Hills, southwestern Poland. Note imbricate stacking of glacial strata and preglacial (Miocene) sediments above the edge of a buried horst (uplift) in hard Triassic bedrock. Based on unpublished information from Krzyszkowski (1993). Compare with Ostr~esz6w Hills (fig. 9-25).
& ! ::i!i:::i:: It~:::;iii::iii::ii i:: ~::i !::: ::i !::~ :::.i:: :::::::::::!:::::iii ::::::::i :::t::i:: ::::::~t~:: ::::i:: ::i: :i:, ::i::iiii ::i::i:: :::: ::::i:::.i:itl ::::::~:: :t::i:: :iti :i:i~ :i:: i:: ii:: itit~41i~iiti:: iii::~ ~:ii:#: ::::::::::iiii!t::i::ti :::::::::~:: ::::i :i::::! :iitii::i ~:i i:: ::i iiiiiiiiilli:: ::~ i:: iiitiiiiiiiT:i !i:: iii::ii iiiii iiii:: ii~: iilii::iii itii::i ):iii i:: iii:Jiitii i:: i:: ;iit:: i!i:: ::i ~i
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:
i ! ii i!!iili !i!i!iiiiii!iiiiiii~iiiiiiiii iiii i iiiii iiiiiiiliiiiiiiiiiiiiiiiiiiiiiiii!i ii ii!iiiiiiiiiiiiiiiiiiiiiiiiiiiiiiiii!iii!iiiiii~iiiiii iiiiii,-.-;. ~
~
shoved hill, and basal till may build up on the proximal side. Such features would ideally outline the frontal and lateral margins of the ice lobe at the time of thrusting and may be preserved in those cases where ice did not later overrun the hill. Continued glacier advance may eventually overrun the iceshoved hill, at which point a penetrative, highly sheared, metamorphic style of deformation may develop beneath the ice (Hart and Boulton 1991). Erosion of the hill provides
Q
2
4:
6
8
~iO
reworked sediment for a discordant till cover, and gradually a cupola-hill morphology develops. Further subglacial modification could produce streamlined, crag-and-tail, drumlin forms (Moran et al. 1980; van den Berg and Beets 1987). The ice-shoved hill ultimately may be destroyed completely in the subglacial environment. Under ideal circumstances, a series of ice-shoved hills may be created at different stages during ice-lobe advance and manage to survive throughout the glaciation. This results in
1
6
4
A
b
........
e
r
,.
and Ber
Figure 11-18. North-south cross section between Schoonebeek and Hengelo in the eastern Netherlands showing position of ice-shoved hill and basin at point where Tertiary strata rise toward the surface and Pleistocene sediments become thinner. Ice movement from north; note large vertical exaggeration. Adapted from van den Berg and Beets (1987, fig. 6).
Slcmbols: .....
'De¢o||emenl
~~; ~
High pore~wate:r
................~
-~""/
presSUre
~
Bas~l T|II
~
ice-con!L=¢t D~It
._.i/~
a glaciotectonic landscape (fig. 11-20), in which narrow belts or loops of ice-shoved hills alternate with wide, low basins. This situation is seen in the Netherlands, for example, where multiple stages of ice-shoved hills are found (see fig. 5-32). Glaciotectonic landscape is also well developed in northeastern Poland (fig. 11-21 and 22). The lobate pattern of glaciation and resulting glaciotectonic phenomena vary greatly in their types and geomorphic expressions. Nonetheless certain basic elements may be expected in subglacial, ice-marginal, and proglacial positions relative to the ice lobe. Melt-water drainage is a key factor for the reconstruction of paleo-ice lobes, as depicted by characteristic geomorphic forms: tunnel valleys, eskers, spillway channels, kames, outwash fans and deltas, etc. Such melt-water features are an integrated aspect of the lobate pattern of glaciation and are often closely connected with glaciotectonism. Glaciotectonic structures and landforms are, thus, part and parcel of the overall ice-lobe and melt-water impact on the landscape. Based on representative lobes of the southern Laurentide ice sheet, Clark (1992) recognized two principal types of ice lobes and their associated geomorphic features for the late Wisconsin glaciation. • Relatively thick lobes with steep margins, such as the Green Bay lobe (fig. 11-23). These lobes presumably advanced over permafrost by a combination of plastic flow and limited basal sliding. Such lobes were able to maintain continued movement for an extended period of time, which resulted in well-developed drumlin fields behind end moraines. Subglacial melt-water features, such as eskers and tunnel valleys, are generally lacking.
Figure 11-19. Schematic model for proglacial thrusting and subglacial modification of an ice-shoved hill during ice-lobe advance. A - initial proglacial thrusting, B - building icemarginal composite ridge, C - overriding and smoothing cupola hill. Note creation of ice-scooped basin (hole) behind the ice-shoved hill. Not to scale; subglacial melt-water features not shown. Modified from aber (1982, fig. 3).
• Relatively thin lobes with gentle margins, such as the James lobe (fig. 11-24). These lobes advanced rapidly over thawed ground by surging on water-lubricated beds or deforming beds. Repeated advances alternated with ice stagnation; hummocky moraine, eskers, tunnel valleys and other geomorphic features of stagnant ice are common.
Distribution of glaciotectonism
165
................
.~.~.~.~...,~. ...... ........
Figure 11-20. Long profile of idealized glaciotectonic landscape during ice-lobe advance and overriding. Note creation of ice-scooped basins between ice-shoved hills. Not to scale; adapted from Aber (1982, fig. 4).
° -
, %k,.._
. ~
f~
a
4X °.,;, • ~
~.~
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?:k~
? .........................................
In our experience, conspicuous constructional glaciotectonic landforms are mostly associated with ice-marginal positions for the latter type of ice lobe. Large hill-hole pairs, composite ridges, and cupola hills are characteristic of relatively thin
Figure 11-21. Sketch map showing festoon pattern of iceshoved ridges (a) and intervening basins (b) for the Suwatki Lakeland, Poland. Adapted from Ber (1987, fig. 6).
Figure 11-22. Festoon pattern in the landscape in vicinity Elk Upland, northeastern Poland. Shaded relief image derived from digital elevation model. Image courtesy of S. Ostaficzuk (2005). Scale in km.
22~55
22,6
22.65
22.7
22~75
22,8
22,85
22.9
1
6
6
A
i
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e
Figure 11-23. Major moraines and drumlins of the Green Bay ice lobe (A), reconstruction of ice-surface morphology (B), and profile of ice margin (C). Green Bay lobe is located in eastern Wisconsin, an area underlain by well-consolidated Paleozoic sedimentary bedrock. Adapted from Clark (1992, fig. 9).
88 °
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and Ber
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'11t 1: i t
~0
~\ \ :\: \ \\. \ "~,
25 50 (kmj
75
3
..................:................i................................. ,,,,,.,.~.............
Figure 11-24. Major moraines of the James ice lobe (A), reconstruction of ice-surface morphology (B), and profile of lobe axis (C). James lobe is located in eastern South Dakota, an area underlain mainly by weakly consolidated Cretaceous shale. Adapted from Clark (1992, fig. 7).
98 ° "ZEELAND
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44 °
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Distribution of glaciotectonism ice lobes that advanced rapidly or surged during late phases of glaciations. Surging is particularly implicated for many small Neoglacial push moraines (Boulton et al. 1999). This depends, of course, upon the presence of suitably weak substratum materials and appropriate geohydrologic conditions. On the other hand, concealed glaciotectonic structures, such as large megablocks, are more likely to have formed beneath thick ice lobes with frozen soles. As Patterson (1993) pointed out, however, some ice lobes were thick and slow during early phases of ice advance, and then changed to thin and fast lobes during later stages of glaciation. Considerable variability should be expected for the style of lobate glaciation and, thus, for the genesis of glaciotectonic features on a regional basis.
Glaciotectonic patterns in North America Glaciotectonic phenomena are widely found throughout the glaciated portions of Canada and the United States, as demonstrated with case studies in this book (chaps. 4-10). Distribution of glaciotectonic features over the continent is related primarily to the presence of thick deformable substrata. Such substrata generally include poorly consolidated Mesozoic and Cenozoic sedimentary rocks as well as unconsolidated Quaternary sediments. Basement faults and seismic zones are common in areas of hard bedrock, particularly the northern Appalachian region. The local presence or absence of glaciotectonic phenomena may reflect variations in glacier dynamics, land slope, aquifers, proglacial lakes or seas, permafrost, substratum lithology, buried hard-rock structures, or such other features that varied geographically or temporally during glaciation. Four regions are particularly noteworthy for their abundance of large glaciotectonic landforms and structures (Aber et al. 1995).
167 streamlined terrain. In front of the ice-shoved hills, spillway valleys, outwash fans, deltas, or stagnant-ice terrain are common. The Devils Lake area of eastern North Dakota demonstrates a representative association of this type in connection with multiple source basins developed over a major aquifer (see figs. 4-13 and 4-14). Proglacial lakes and seas were important elements in many situations. For example, Baldwin Peninsula is an exceptionally large push-moraine complex constructed entirely of middle Pleistocene marine, fluvial and glacial sediments. The peninsula is ~90 km long, located on Kotzebue Sound of the northwestern Alaskan coast (fig. 1125). The bulk of deformed strata exposed in cliffs consists of glaciomarine and glaciofluvial sediments laid down in front of three coalescent ice lobes that advanced southwestward from the Brooks Range. Baldwin Peninsula represents the terminal limit of the Anaktuvuk River glaciation (~500-600 ka BP), which had an areal extent nearly tens times greater than the late Wisconsin glaciation in the same region (Huston, Brigham-Grette and Hopkins 1990). Of these four regions, the northern Great Plains undoubtedly has the greatest geographic coverage and abundance of glaciotectonic phenomena (Aber et al. 1991; 1995). Nearly all of this region is underlain by weakly consolidated Cretaceous and Tertiary sedimentary strata. Three common associations are apparent; glaciotectonic features are frequently found in conjunction with 1) bedrock escarpments, 2) subsurface aquifers, and 3) identifiable ice-margin positions (fig. 11-26). Furthermore, ice margins are often located along bedrock escarpments, because of the
• Atlantic Coastal Plain including Long Island, New York and southern New England (see fig. 6-2; Oldale and O'Hara 1984). • Northern Great Plains of the United States and the Canadian Prairie region (Kupsch 1962; Moran et al. 1980). • Yukon Coastal Plain of the Northwest and Yukon Territories, northern Canada (see fig. 4-9; Mackay 1959). • Bering Sea coastal zone of western Alaska (Huston, Brigham-Grette and Hopkins 1990) and eastern Siberia (Benson 1993). Recurring geomorphic patterns are developed within these four regions in terms of the association of glaciotectonic landforms with other typical glacial features. Most large ice-shoved hills and ice-scooped basins were created at or just behind ice-margin positions. Various characteristic features are found in the subglacial zone behind the hills. These include tunnel valleys, eskers, drumlins, and
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Aber and Ber
168
Figure 11-26. Composite map of late Wisconsin ice-margin positions of the northern Great Plains region, showing prinicipal ice lobes and major escarpments (coteau). Adapted from Clayton, Moran and Bluemle (1980, fig. 33).
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topographic control of ice-lobe movement. Primary escarpments are the Prairie Coteau in eastern South Dakota and southwestern Minnesota, and the Missouri Coteau of North Dakota and southern Saskatchewan. The easternmost edge of the Great Plains glaciotectonic province is located in central Minnesota, where a thin, patchy veneer of Cretaceous shale is preserved in low positions on the Precambrian basement surface (fig. 11-27). The St. Croix, Stewart Lake, and Outing moraines contain numerous displaced masses of marine and coal-beating shales, saprolite derived from the basement, as well as deformed pre-late Wisconsin glacial materials (Mooers 1990; Knaeble 1998). Well-formed composite ridges and hill-hole pairs are situated within the three moraines (fig. 11-28). Ice-sheet advance upslope over poorly consolidated substrata containing confined aquifers were major factors for inducing widespread glaciotectonic deformation in the Great Plains. A narrow frozen bed at the ice-lobe margin is also considered to be an important factor in many cases (Mooers 1990). During late phases of the late Wisconsin glacial cycle, the southwestern margin of the Laurentide ice sheet probably experienced repeated surging of relatively thin ice lobes (Clayton, Teller and Attig 1985), for example, the James lobe (fig. 11-24). The likelihood of ice-lobe surging over deforming beds is corroborated by widespread clast pavements in the Great Plains region (fig. 11-29). These conditions, in combination, brought about numerous and diverse glaciotectonic structures and landforms throughout the northcentral United States and Canadian Prairie region.
Notable glaciotectonic differences are, nonetheless, apparent within the Great Plains (Aber et al. 1995). A great many large and small glaciotectonic features are present in North Dakota, for example. Southern Saskatchewan, in contrast, apparently contains fewer and larger glaciotectonic features. Part of this difference is real; the impressive ice-shoved hills and giant megablocks of Saskatchewan are among the largest in the world. Another part of the difference results from geologic mapping. Large-scale county mapping has been done for North Dakota, whereas southern Saskatchewan has received only reconnaissance mapping at 1:250,000 scale. Certain areas are conspicuous by their apparent lack of glaciotectonic features, such as eastern South Dakota and northeastern Montana. Casual observations suggest that iceshoved hills, source basins, and other glaciotectonic features are probably common in these areas (fig. 11-30), but they have never been mapped or described to our knowledge. We suspect many more glaciotectonic phenomena, especially concealed structures, are actually present within the northern Great Plains region.
Glaciotectonic patterns in central Europe Glaciotectonic patterns are described for the region stretching from the Netherlands on the west to the Ukraine on the southeast, and including portions of Germany, Denmark, Czech Republic, Russia (Kaliningrad District), Poland, Belarus, Lithuania, Latvia and Estonia. This region was glaciated several times during the Pleistocene by continental ice sheets, and therefore glaciotectonic features of different
Distribution of glaciotectonism
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Aber and Ber
170
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Figure 11-29. General distribution of clast pavements associated with the Laurentide ice sheet. Pavements (dots and lines) are found on sedimentary substratum (blank) surrounding the Canadian Shield (stippled). Note the particular abundance of clast pavements in the northern U.S. Great Plains and Canadian Prairie regions. Heavy black line depicts limit of late Wisconsin glaciation. Adapted from Clark (1991, fig. 1).
ages are found widely throughout the region both on land and on the North Sea and Baltic Sea floors. The recognition of glaciotectonic features in central Europe was possible through preparation of the Glaciotectonic Map of Central Europe (Ber and Aber 2003), based on compilation of glaciotectonic national maps elaborated for Germany (Stephan), Denmark (Jakobsen), Poland (Ber), Czech Republic (Nyvlt), Ukraine (Matoshko), Belarus (Karabanov), Lithuania (Bitinas and Aleksa), Russia- Kaliningrad District (Zagorodnikh), Latvia (Zelcs and Dzelzitis) and Estonia (Kalm and Rattas). The glaciotectonic features were classified in four main categories: • • • •
glaciotectonic landform expressed in modern relief. glaciotectonic buried or concealed landforms. glaciotectonically disturbed Quaternary sediment. glaciotectonically disturbed pre-Quaternary strata.
As in North America, distribution of glaciotectonic features over the territory of central Europe is limited within the outer zone of glaciation (Aber 1992, see fig. 11-2), and is related primarily to the extent of younger glaciations (i.e. Weichselian and Saalian) and thick deformable substrata consisting of unconsolidated Quaternary deposits and poorly consolidated Mesozoic and Cenozoic sedimentary rocks. The landforms and sediments were deformed mostly by active ice sheets through horizontal and vertical thrusting of the
Figure 11-30. Space-shuttle photograph of Prairie Coteau, Watertown vicinity, eastern South Dakota. The western margin of the Des Moines lobe is denoted by the Bemis moraine; a prominent lateral moraine with numerous large lakes marks the eastern edge of the James lobe. The interlobate zone remained ice free during late Wisconsin glaciation. Lakes have ice/snow cover in this late winter scene. Image adapted from NASA Johnson Space Center; STS36-152-110, 3/90, courtesy of K. Lulla. ice masses, as well as by vertical stress of the glacier loads. It is also probable that surficial morphology of the glaciated portion of central Europe reflects underlying influences of the configuration and geological structure of the crystalline basement, particularly along lithological boundaries, fractures, and fault zones (Ber and Ryka 1998; Ber 2000). These zones were activated by weight and movement of the advancing and departing ice sheets. Especially parallel geological boundaries of the older substratum and large faults connected with them had a basic effect on parallel direction of many glaciotectonic zones often marked in the present landscape as composite ridges and push moraines. The local presence or absence of glaciotectonic structures depends upon particular ice-sheet dynamics, ground-water and permafrost conditions, substratum lithology, and buried, consolidated, hard-rock structures. In the glacial landscape of central Europe there are two distinct zones both situated within the outer zone of glaciation (Aber 1992, see fig. 11-2). The first is fresh and expressive, created by the younger Weichselian glaciation. The second is older and changed distinctly by denudation and erosion; it is connected to Saalian and Elsterian glaciations. The extent
Distribution of glaciotectonism of the younger zone is marked by maximal limits of the Weichselian ice sheet; the second zone is located southward from maximal extent of the younger zone and is marked by maximal limits of the Saalian and Elsterian glaciations.
171 well logs. The island of Men displays impressive dislocated chalk masses in MCns Klint (see figs. 5-7 and 5-8), as well as in other cliff exposures and cupola hills (see fig. 6-12). Germany ~ The main features in the glaciotectonic patterns
Central Europe - - Weichselian Glaciation
The maximal limit of the Weichselian Glaciation is marked mainly by push moraines, composite ridges and cupola hills, and by associated glaciodepressions. In the westernmost part, the marginal line runs almost meridionally across the central part of Jutland Peninsula of western Denmark, then turns to the east across northeastern Germany, to northern Poland and Belarus, southern Lithuania, and lastly northern Russia (see fig. 11-2). Representative glaciotectonic phenomena are described from west to east.
of northern Germany (Schleswig-Holstein, MecklenburgVorpommern) are large and long end moraines systems, mainly push moraines of the Brandenburg, Frankfurt and Pomeranian Phases (fig. 11-33), connected to the distal margins of erosive glacial basins. In these Weichselian ridges and elevations, cores of older material are frequent and, according to Stephan (2001b), often partly deformed during the Saalian Glaciation. The remains of Saalian ice-pushed end moraines were overridden and reshaped by the ice sheet during the Weichselian Glaciation, for example Rauensche Berge, elevated grounds of Bad Freienwalde and Ftinfeichen.
Denmark ~ Several regions within Denmark are particularly
noteworthy for their abundance of glaciotectonic disturbances (see fig. 2-24). These glaciotectonic deformations are recorded in numerous well logs, geophysical imagery, landform morphology, and coastal cliff exposures. Deformed substrata consists of thick, unconsolidated Quaternary sediment, as well as poorly lithified Paleogene and Cretaceous bedrock. The glaciotectonic phenomena are related stratigraphically to multiple phases of Weichselian glacier advance and readvance from the north, northeast, east, southeast and south (see figs. 3-22 and 3-23). In the northernmost part of Jutland, situated within the limit of the Main Weichselian Glaciation are found related distinct composite ridges and large glaciotectonic structures in coastal cliff exposures, such as LCnstrup Klint (see figs. 1-4 and 137) and Rubjerg Knude (fig. 11-31). The imbricate complex at Rubjerg Knude was formed proglacially in front of an advancing glacier from a northerly direction. The ice sheet advance created diapiric intrusions of mud penetrating the upper strata, and then overrode the whole structure and deposited till unconformably upon the dislocated beds. Also in the western part of the Limfjord, glaciotectonic deformations are recorded in numerous well logs and in landform exposures (see figs. 2-23 and 5-21). The Young Baltic glaciation and the Ba~lthav readvance are connected with large glaciotectonic deformations situated along the east coast of Jutland (dislocated Paleogene and Eemian marine deposits with Quaternary drift), on the islands of Fyn, Langeland and ~ r ¢ , and in the northwestern part of Sja~lland Island (fig. 11-32). As described by Jakobsen (2003), in the central part of Sja~lland rafts of Cretaceous chalk and Paleocene Kerteminde Marl have been recorded. These rafts are seen in the well-log records, and they presumably were transported about 50 km from the southeast. In the southern part of Sja~lland, elongated composite ridges trend NW-SE, and the Copenhagen area shows glaciotectonism recorded in
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172
Aber and Ber
Figure 11-32. Space-shuttle photographs of southern and eastern Denmark, regions in which glaciotectonic structures and landforms are quite abundant. A - islands of Fyn, ./Ere and Langeland, southern Denmark (STS 45-87-25, March 1992). B island of Sjcelland and Store Bcelt, eastern Denmark (STS 45-152-153, March 1992). Compare with Fig. 2-24. Images obtained from NASA Johnson Space Center, Houston, courtesy of K. Lulla.
The Weichselian morainic landscape of Schleswig-Holstein represents the Eastern Hill Lands (Ostliches Htigelland). In this region push or thrust moraines are frequent, and especially most areas near the Baltic Sea show strong glaciotectonic deformations, as for example Heiligenhafen "Hohe Ufer" cliff section (Stephan 1997b; van der Wateren 1997) or cliff sites at the D~inisher Wohld peninsula between Kiel and Eckernf6rde (Piotrowski 1997; Stephan 1997a). The Heiligenhafen ("Hohe Ufer"), according to Stephan (1997a), is a hill exposed in a cliff, reaching 16 m above sea level, and covered by till of the Young Baltic Advance (upper till). The till covers Pleistocene (mainly Saalian) and uplifted Paleogene (mainly Eocene) deposits. The general structural pattern of the section is a chain of anticlines and synclines with axes striking N-S, and a large thrust slice reaches the surface of the hill. Structures were deformed at least two times: by the Weichselian Young Balic Advance and during the Saalian (Hennstedt) ice sheet. The exposures along the Baltic Sea cliff between Eckemf6rde and Kiel reveal glacial deposits and their glaciotectonic deformations over a distance of 30 km. The exposed sequence was subdivided by Piotrowski (1997) in two complexes (fig. 11-34). The lower, waterlaid complex, strongly glaciotectonically disturbed, corresponds to first (oldest) Weichselian advance. However, the upper complex derives from the youngest Weichselian ice advance. The formation
of the lower sedimentary complex can be explained as follows. In front of the advancing Weichselian ice sheet, a proglacial lake or a system of lakes existed, which was displaced farther to the south during progressive ice advance (fig. 11-35). Finally the glacier overrode the waterlaid sediment sequence, which was accompanied by minor glaciotectonic disturbances, truncation of the substratum, and intensive sheafing at the ice base. Subsequently the area again came into the icemarginal zone during glacier retreat, and superposed large glaciotectonic deformations occurred leading to the formation of high-amplitude folds (Piotrowski 1997). In the Mecklenburg-Vorpommern area, a prominent landscape feature is the main end moraine of the Pomeranian ice advance which stretches from northwest to southeast. In the Neukloster area, a ridge chain consists of push ridges, boulder agglomerations, end moraines, proximal sandurs, and the so-called "Hohe Burg" with strongly undulating elevations (Bremer and Mailer 1997). At Babst (also Neukloster area), numerous NW-SE oriented ridges of push moraines occur, which can be traced about 10 km northwest and 4 km to the southeast. According to Rtihberg (1997a), northwest of Grevesmtihlen a large push-moraine complex occurs and is characterized by numerous sub-parallel push ridges and abundant large boulders. Likewise along the coast of the Baltic Sea, huge basins shaped by glacial erosion contain common
Distribution of glaciotectonism
173
Figure 11-33. MODIS natural-color satellite image of northern Germany, southern Denmark and northwestern Poland. The Elbe River valley marks the approximate maximum limit of Weischelian glaciation. Image date 2/20/04, 250 m resolution. Obtained from MODIS Rapid Response System . Image downloaded Nov. 2005.
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Figure 11-36. Space-shuttle photograph of the island of Riigen, northeastern Germany (STS 45-87-27, March 1992). The chalk cliffs of Jasmund Peninsula are quite comparable to MCns Klint, Denmark. Image obtainedfrom NASA Johnson Space Center, Houston, courtesy of K. Lulla.
glaciotectonic structures, as seen in the cliff on the western part of the Isle of Poel (RUhberg 1997b) or the famous chalk cliffs on the island of Riigen (fig. 11-36). The latter bear a striking resemblance to MCns Klint, across the Baltic strait in Denmark. The Cretaceous chalk deformation of Rtigen, along the 18km coast between Jasmund and Glowe, displays overturned synclines with cores of Pleistocene deposits and overthrust upper limbs (Panzig 1997a & b; Panzig and Kanter 1997). According to Panzig (1997c), simultaneously with the ice advance the Cretaceous chalk broke into blocks in front of the glacier. Nextly these chalk blocks were cut bench-like by the advancing ice or were thrust over adjoining blocks, thus protecting the Pleistocene strata between the thrust blocks (fig. 11-37). Displaced Cretaceous chalk and Pleistocene deposits resulted in surficial landscapes on Jasmund Peninsula consisting of composite ridges, mostly with disturbed internal structures (folds and scales) running in two directions: nearly parallel (E-W) in the northern part of peninsula and from SSW to NNE in the southern part. These trends mark two separate glacier advances and stresses: one from north to south and second one with southeast to northwest direction.
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Figure 11-37. Schematic illustration for creation of main glaciotectonic structures (chalk thrust blocks) in the Jasmund Peninsula, island of Riigen, northeastern Germany. Based on Herrig (reproduced in Piotrowski 2002). Poland is representative of the southern Baltic region of central Europe, which has much in common with the glaciated Great Plains in North America. Several factors contributed to the wealth of glaciotectonic phenomena.
175
Distribution of glaciotectonism
• Poland is underlain by weakly consolidated or loose substrata of predominately Cretaceous and Cenozoic age, including upper Cretaceous chalk and Paleogene and Neogene clay-rich formations • Ice sheets advanced upslope, which blocked regional drainage, ponded proglacial lakes or seas, and confined aquifers. • Permafrost was present in front of advancing glaciers, particularly in upland regions south of the Baltic basin. • Repeated glaciations took place, each with multiple advances and readvances, often in the form of ice lobes or tongues following topographic lowlands. Late readvances were probably as surges (Aber and Ruszczyfiska-Szenajch
This combination of conditions facilitated widespread glaciotectonism, particularly in situations where buried hardrock structures, such as horsts, are present in the shallow subsurface. Glaciotectonic features are nearly ubiquitous across much of Poland. However, three regions are especially notable for Vistulian glaciation (fig. 11-38). • Wolin Island vicinity - displaced chalk and drift similar to nearby islands of Rtigen and Usedom, Germany and Men, Denmark. • Elbl~g Upland- huge glaciotectonic massif composed largely of unconsolidated Pleistocene strata displaced from the Gdansk embayment lowland (see chap. 6).
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Suwatki Upland- glaciotectonism was influenced strongly by basement fault zones, which were reactivated during successive glaciations (see chap. 9). 1880 ~ 65 >
Wolin Island is situated in the sublitoral zone of Pomorze Bay, northwestern Poland. A moraine ridge (115 m elev.) is connected with the latest phase of the last glaciation and is the main element of the relief. It is bordered by a high cliff that reveals numerous small and large glaciotectonic structures. A few huge blocks of Cretaceous strata (chalk and marl) occur within the moraine and outcrop in quarries and in the cliff. The blocks were transported as rafts by the last glaciation from the Baltic Sea floor. Much debate concerns the timing and directions of glaciotectonic deformation for this key geographic location (Persson and Lagerlund 1994; Lagerlund et al. 1995; RuszczyfiskaSzenajch 1999).
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In coastal cliffs on Wolin Island, glaciotectonic deformations are connected with the following sedimentary deposits: Cretaceous marls, forming scales and rafts; grey, Middle Polish (Grodno) till and brown till (Trz~sacz) of Late Vistulian age; and intertill sands and silts with gravels (fig. 11-39). In the bottom of tills, numerous glaciotectonic structures are present (Bor6wka, Goslar and Pazdur 1998). Wolin Island displays a distinct connection of surface landforms (morainic massifs) and their longer axes of elevations and glaciodepressions to azimuths of the main glaciotectonic structures. The surface of Wolin Island is diversified by numerous ridges of push moraines built with glaciotectonically disturbed Pleistocene deposits including Cretaceous rafts. According to Bor6wka, Goslar and Pazdur (1998), the geologic and geomorphologic data suggest that the glaciotectonic structures present in the cliff section are connected with the retreat (not advance) of the last (Vistulian) ice sheet. Vistulian glacier flow direction and deformation of Wolin Island deposits traditionally were interpreted as from NNENE directions. However, Lagerlund (1987, 1997; et al. 1995) and Bj6rklund, Lagerlund and Ing61fsson (1998) challenged this model. According to them, western and central portions of Wolin were deformed by ice movements from the west and northwest, which distinctly deviate from the predominating NNE-NE directions. Similar anomalous ice movement directions in Sweden (western Sk~ne) and northeastern Germany (Usedom, Riigen) are indicated (fig. 11-40). In eastern Wolin, in contrast, glaciotectonic structures suggest ice pushing from the northeast. Belarus m Only the northern part of the country is located within the limit of Poozerie (Weichselian) glaciation (fig. 11-41). The maximum limit of this glaciation is marked mainly by push moraines with glaciotectonically disturbed
Grodno till M i l l e Polish
Figure 11-39. Generalized stratigraphic sequence of Pleistocene deposits resting on dislocated Cretaceous strata from Wolin Island. Ages given in radiocarbon years before present. Adapted from Bor6wka, Goslar and Pazdur (1998, fig. 1).
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Figure 11-40. Anomalous directions of Weichselian ice movement in the vicinity of Riigen (R) and Wolin (W) islands. Arrows show groups of anomalous directions at different sites and different stratigraphic levels. Adapted from Lagerlund (1997, fig. 1).
177
Distribution of glaciotectonism
internal structures composed of dislocated tills and glaciofluvial deposits. According to Karabanov (2000), folded (Sopotskin) and injective ridge systems (Ushahi Lake group) are connected with faults in the crystalline and sedimentary basement of northern Belarus. Most common among glaciotectonic structures are folded dislocations, glaciodepressions, diapiric hills and ridges, and rafts composed of Cretaceous, Paleogene or Neogene strata. Glaciotectonic landforms often are associated with dislocations of Cretaceous rocks (e.g. Sopotskin, Grodno). The Grodno Highland area is a well-known example of a glacial inter!obate massif (cupola hill), similar to insular glaciostructural-accumulative highlands in Latvia and Estonia or the isolated Pleistocene elevations in Poland. It was influenced by the main ice streams during Middle and Late Pleistocene glaciations (Karabanov 1987). According to Pavlovskaya and Karabanov (2002), most of Grodno Highland is deformed in the shape of hill-hole pairs, stretching generally in W-E direction, with impressive dislocations of Cretaceous rocks at Grandichi, Pyshki, Melovaya Gora and Sopotskin. The chalk quarry at Pyshki, for example, reveals folded scales of upper Cretaceous chalk and chalky marl deposits, as well as upper Miocene and lower Oligocene (glauconite and quartz, sand and silt) strata, and injective forms and megablocks (Karabanov 2002a). These deposits are overlain by Pleistocene till and glaciofluvial sand (fig. 11-42). The Cretaceous deposits are normally located 80-100 m below water level in the Neman River, but where involved in glaciotectonic dislocations they are exposed at the surface.
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Injective structures in Belarus are represented by diapirs, diapiric hills and squeezed ridges, most often built of chalk, clay and clayey sands. The largest diapirs are composed of chalk and are up to 100-150 m high (Grodno). However, injective ridges are present in Chashniki Plain and in Ushachi Lakes area (Karabanov 2000). Baltic s t a t e s - Lithuania, Latvia, and Estonia occupy the eastern peribaltic sector of central Europe (fig. 11-43). In general, glacial erosion predominated toward the north of this region, whereas glacial deposition was more significant southward. The main glacial landforms consist of ridges and hills of accumulative and push moraines, widespread highlands (accumulative insular heights), interlobate complexes, and drumlins that mark ice-flow directions (fig. 11-44). Glaciotectonic phenomena of many types are widely distributed from the Baltic Klint of northern Estonia to southeastern Lithuania.
According to Aleksa and Bitinas (2000), Lithuania can be subdivided in two zones. The first zone of predominant erosion is located in northern and partly central Lithuania and is characterized by relatively thin Quaternary deposits less than 15 m thick. Drumlins and flutings dominate in the landscape. The southern part of Lithuania belongs to the second zone of glacial accumulation, mainly formed during the last glaciation (Nemunas, Weichselian), where glaciotectonized composite massifs and marginal ridges are conspicuous in the landscape. In this zone, a few hill-hole pairs are distinguished only according to their geomorphic expression. The ~emaitia Upland is noteworthy (fig. 11-44). It is a sizable glacial highland (~300 m elev.) that probably contains glaciotectonically disturbed internal structures (fig. 11-45). The ~emaitia Upland is similar to other accumulative insular heights in the eastern Baltic, such as Vidzeme in Latvia, and Haanja and Otep~i~i in Estonia. These features resemble isolated Pleistocene highlands in Poland, including Elbl~g Upland and Suwatki Lakeland (Mojski 1998).
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Figure 11-42. Anticlinal fold in Cretaceous chalk, Pyshki quarry, Grodno Highland, western Belarus. 1 - chalk, 2 quartz sand and silt, 3 - gravel, 4 - gravelly sand, 5 laminated sand, 6 - phosphorite concretions. Adapted from Pavlovskaya and Karabanov (2002, fig. 1).
178
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Figure 11-43. MODIS natural-color satellite image of the eastern peribaltic region, including Lithuania (LT), Latvia (LV) and Estonia (EE). This region was completely overrun by Weichselian and earlier glaciations. Image date 04/02/2004, 250 m resolution. Obtained from MODIS Rapid Response System . Image downloaded Nov. 2005.
179
Distribution of glaciotectonism
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In both zones, but particularly in the second one, preQuaternary rafts are present. Rafts of Cretaceous rocks are mostly present in southeastern Lithuania; however, older Jurassic, Triassic and Devonian rafts prevail in the western part of the country (Aleksa and Bitinas 2000). The distribution of rafts and their formation depended upon subQuaternary relief and physico-mechanical properties of the pre-Quaternary bedrock. In the territory of Latvia, according to Dreimanis and Zelcs (1995), Kartikapp (1997), and Zelcs and Markots (2004), glacial landforms occupy two general settings in the landscape--lowlands and uplands (fig. 11-44). Insular accumulative highlands, such as Latgale and Vidzeme, are most conspicuous and are considered to be glaciostructural
in origin. Within the lowlands, characteristic features are drumlins with glaciotectonically disturbed cores. The concealed structures of glaciotectonic plains in Latvia may be classified in three types (Zelcs 1998): • thin till and concealed folds, diapiric structures, and rafts of glacioaquatic sediment transported up to 6-8 km. • thin glacial deposits, including rafts and megablocks of upper Devonian dolomite transported short distances. • deformed Holsteinian and Eemian sediments (western Latvia and Gulf of Riga). The central part of the Baltic Sea coast is noteworthy in western Latvia (at Sensala, Ulmala and Strante), where 10to 15-m-high bluffs expose a glaciotectonically disturbed
Aber and Ber
180
Figure 11-45. Photographs of glaciotectonic structures. A recumbent folds in glaciofluvial gravel exposed in Pakalni~kiai quarry, northwestern Lithuania. B - till diapir in sand at the zvelsa Creek in the Klaipeda District, western Lithuania. Height of exposed section ~4 m. Photos courtesy of A. Bitinas.
THE GLACIOTECTONIC MAP OF ESTONIA: Lambert
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Distribution of glaciotectonism
181
Figure 11-47. Panoramic view of Haanja heights looking toward the northeast from the observation tower at Suur Munamiigi (317 m elevation). Photo by J.S. Aber (2000). sequence of Pleistocene deposits. The cliffs are built mainly of intertill sand and silt sediment, and clayey material is folded into diapirs (Dreimanis et al. 2004). Generally two types of glaciotectonic deformations are present in E s t o n i a m h a r d bedrock dislocations and soft-bed disturbances (Rattas and Kalm 2004). The distribution of both types is related to the thickness of the Pleistocene cover. The hard bedrock disturbances within carbonate rocks, for example, occur in the narrow zone next to the Baltic Klint (see fig. 7-15). Bedrock dislocations, fractures, folds and fissures, and large rafts of pre-Quaternary rocks are predominant in northern and central Estonia (fig. 11-46). To the south, however, glaciotectonically disturbed Quaternary deposits are the most common. Also the preQuaternary surface has been strongly glaciotectonized in the forms of folds, fractures, fissures, and overthrusts, as well as megablocks moved for short distances (2-3 km) from their sources. Especially the poorly consolidated Devonian sandstones and siltstones in southern Estonia have been folded and fractured by subglacial deformation. Several bedrock disturbances in northeastern Estonia also are known from boreholes (see fig. 7-16). Composite massifs and ridges built of glaciotectonically disturbed deposits, including push moraines and drumlins, are the most common glaciotectonic landforms (fig. 11-46). Most remarkable are the composite massifs such as Otep~i~i, Haanja, Karula and Sakala, which are composed of heaped and subsequently deformed drift masses. These landforms are characterized by great thickness of glacial deposits (i.e. tills up to 70 m thick) and complicated relief (fig. 11-47). Their cores often are built with glaciotectonically disturbed Late Ugandi (Saalian, Warthanian) tills (Rattas and Kalm 2004). The ice-marginal zones of the Late Weichselian deglaciation are marked mostly by ridges and hills of push moraines (fig. 11-46), which contain listric thrusts, faults and folds that were produced by proglacial compressive deformation and by shearing beneath moving ice (Rattas and Kalm 2004).
Central Europe ~ Saalian and Elsterian glaciations
The overall maximum limits of the Saalian and Elsterian ice sheets are nearly the same (fig. 11-2), although distinct differences emerge in southeastern Germany, Poland, and the Ukraine. Also distinct differences are visible in preservation of marginal landforms and sediments. The older Elsterian glaciation is marked poorly in present relief by strongly denuded and eroded glacial landforms, weathered tills, outwash plains or sandurs, as well as by erratic boulders. Evidence for the younger Saalian glaciation, in contrast, is morphologically distinct, and glacial deposits are well preserved along much of its limit. Glaciotectonic structures and landforms are reviewed beginning in the Netherlands on the west and proceeding eastward to the Ukraine. Netherlands m Reconstruction of the maximum extent of the Elsterian advance is difficult, because a morphologically distinct ice margin is absent. Elsterian glaciation probably reached the northern sector of the Netherlands and adjacent North Sea floor (Laban and van der Meer 2004). The modem landscape, including conspicuous composite ridges, was created during the Saalian glaciation, which reached the central portion of the country (see figs. 5-29 and 5-32).
Generally, according to Laban and van der Meer (2004), Saalian ice-pushed ridges are asymmetric with steep proximal slopes and consist mainly of Early and Middle Pleistocene deposits (central part of the country) and Neogene and Paleogene deposits (eastern part of the country). During the first (oldest) Saalian ice sheet advance, which reached the maximal extent of this glaciation in the Netherlands, Utrecht Ridge (fig. 5-29) and Veluwe Ridge (see fig. 12-9) were built. Also in this same region, ice-dammed lakes located within the glacial basins were developed. Germany ~ According to Ehlers et al. (2004), numerous distinct ice-shoved moraines (southern Dtibener Heide, Dahlener Heide areas) were formed during the Elsterian ice sheet advance, which also created the Muskau Arc. In the
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Figure 11-48. Main elements of the Muskau Arc (MFB) of glaciotectonic disturbances in eastern Germany and southwestern Poland (inset map). Gray shading with black lines indicates glaciotectonic deformations visible in surface morphology. Morphostructural arcs: KT- KOlzig-Tuplice, DT- Do'bern-Trzebiel, LM- L¢knica-Muskau. northern regions, Elsterian thrust moraines are mostly overlain by younger sediments, partly superimposed by younger deformations, or deeply buried (northwestern portion). Glaciotectonic deformations due to Saalian glacier advances typify a broad belt from south of Berlin to the northern Netherlands. The Muskau Arc (Muskauer Faltenbogen - MFB) is a c o m p o s i t e - r i d g e system extending as a large loop approximately 22 km wide and 20 km long, located near the town Bad Muskau (figs. 11-48 and 11-49; Kupetz 1997, 2000). The MFB is mainly in Germany and partly in Poland. A Miocene brown coal sequence up to 10 m thick is glaciotectonically deformed (folded and thrust) and locally uplifted to the ground surface. In the German part, such dislocated brown coal strata were extracted in 50 to 60 mines and open-cast pits from 1830 until 1970. The MFB consists of ~30 single linear structural elements, which are divided as follows (fig. 11-50).
• declined (vergent) structural elements, such as thrust sheets, vergent anticlines and synclines. • symmetrical (erect) structural elements, such as ejective folds and wide synclines. According to Kupetz (2000), thrust sheets with short transport distances are most frequent, and they increase in number from north to south. However, far-thrust, rootless rafts are unknown. The ejective fold genesis is interpreted as a result of tectonoclastic flow. The single elements of the MFB are locally cut by straight relaxation structures, marked mainly by clastic dikes consisting of sand and/or clay. Deformation of MFB took place in several marked phases during a single Elsterian glacial advance. The overall size and structural framework of MFB bear much in common with the Dirt Hills and Cactus Hills of Saskatchewan (see fig. 5-18).
Czech Republic - - Two or three major ice advances during the Elsterian and two during the Saalian have been
Distribution of glaciotectonism
183
Figure 11-49. Shaded-relief digital elevation model for the Muskau Arc (MFB) in eastern Germany and southwestern Poland. Compare with Fig. 11-48. Image courtesy of J. Ko~ma (2005).
distinguished in northern Bohemia. Traces of glaciotectonic structures from the older Elsterian (Dubnice) glaciation are known in the Fr~dlant vicinity (fig. 11-51). Gravel pits expose relics of ice-pushed and folded subglacial till together with overlying glaciofluvial gravel (younger Elsterian, Lvov~ glaciation), which were deformed by the advancing Saalian ice sheet. In northeastern Czech Republic, northern and southern push-moraine zones in the Opavska pahorkatina Upland are noteworthy (fig. 11-52). In the Opava Upland during early Elsterian glaciation, the ice sheet overrode the northern pushmoraine zone. Near Hlucfn, Ruzicka (2004) described large floes, 10s to 100s of meters in size composed of glaciotectonically deformed preglacial and Elsterian glacial sediments. According to Ruzicka (2004), the Saalian glaciation of the Czech Republic had less influence on relief than the Elsterian glaciation. The early Saalian (Jitrava) ice sheet entered from the Zittau basin into the Hradek and Nisou area and caused large glaciotectonic deformations of Elsterian sediments and
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Figure 11-50. Structural sections revealed in Miocene brown coal mines. A - imbricated thrusts with overturned folds, B - symmetrical folds. Adapted from Kupetz (2000, fig. 1).
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Figure 11-51. Maximal extent of continental glaciation in Bohemia, northern Czech Republic. Glaciotectonic structures are found in the Frajdlant vicinity. Adapted from Ruzicka (2004, fig. 1).
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of the underlying Miocene strata. In northern Moravia and Silesia (fig. 11-52), the Saalian ice sheet advanced twice into Ostrava basin. During the advance of the late Saalian ice sheet, the northern push-moraine zone was strongly deformed and morphologically enlarged. This marginal zone is the best preserved landscape depositional zone (i.e. Chuchelna push moraine) of the late Saalian glaciation. P o l a n d - The Silesian Rampart and Zielona G6ra Rampart in southwestern Poland are especially notable, as well as the Polish part of the Muskau Arc, within the limits of the older
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glaciations. The landforms in this region consist of a series of prominent composite ridges including Ostrzesz6w, Trzebnica, Wifisko, and Dalk6w (see chap. 9 and fig. 11-17). These ice-shoved hills were thrust above buried hard-rock horsts and include substantial volumes of dislocated preQuaternary strata (fig. 11-53). According to Badura and Przybylski (2002) the scale of disturbances decreases from north to south from >200 m thickness (Zielona G6ra Rampart) to 2-10 m depth in the intrasudetic valleys. Glaciotectonic deformations include mainly Quaternary and late Neogene
Distribution of glaciotectonism
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Figure 11-53. Cross section through the Wihsko Hills, Silesian Rampart, southwestern Poland. Deformation involves the Miocene Poznah clay and overlying Pleistocene strata. Vertical scale in m above seal level. Adapted from Krzyszkowski and Labno (2002, fig. 3).
deposits, but disturbances of Triassic, Oligocene and Miocene strata are noted as well. Research in the Mu~ak6w (Muskau) Arc (fig. 11-48) and in the Silesian and the Zielona G6ra Ramparts demonstrates the largest glaciotectonic disturbances of the South Polish (Elsterian) Glaciation. Deformed layers are discordantly covered by the Odranian (older Saalian) and the Wartanian (younger Saalian) deposits. The main outline of the Silesian Rampart was formed during recession of the second Elsterian ice sheet. The Mu~,ak6w Arc, which was formed during a rapid surge of an ice lobe, and the Zielona G6ra Rampart
were shaped during the next stages of the ice sheet recession. The origin of comparable arcs near Grgbocice and Syc6w in the Silesian Rampart (fig. 9-23) is probably similar (Badura and Przybylski 2002). This system of composite ridges of the Silesian Rampart bears a striking resemblence to the Missouri Coteau (see fig. 5-16). Belarus ~ The Dnieper glaciation completely overspread Belarus, and about two-thirds of the country was covered during the subsequent Sozh glaciation (see fig. 11-41). Glaciotectonic structures within the landforms or deposits
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are mainly created as large folded dislocations, glacial rafts, diapiric ridges, and glacial depressions (Karabanov 2000). The lower (older) buffed morainic complex of the Dnieper glaciation is strongly dislocated as result of pushing and squeezing during activity of subsequent ice sheets (Komarovsky and Vertinskaya 2000). The large folded dislocations of Belarus are generally arc-shaped and built with pre-Quaternary rocks. Their extents range from 2 to 30 km, and thickness of dislocated deposits varies from 80 to 250 m (Karabanov 2000).
and marl as well as Paleogene glauconite-quartz sand, Neogene clay and quartz sand, and Quaternary deposits (fig. 11-54). The dislocated deposits form at least 20 folds that are thrust over one another at 40-45 ° , and their thickness varies from 40 to 200 m. The Peski glaciodislocation is confined to an area where the crystalline basement, dissected by systems of faults, occurs at a depth of 1000 m. This suggests the fault system was activated during the Dnieper glaciation, when ice loading contributed to disruption of the Mesozoic and Cenozoic blocks (Karabanov and Levkov 1996).
Buried large folded glaciodislocations are present at many sites in western Belarus, as at Grodno, Volkovysk, and Novogrudok (fig. 11-41). In central and eastern regions, more typical are glaciotectonic internal structures within disturbed large uplands, for example at Minsk and Orga. Sizable glaciodiapirs, composed of Cretaceous chalk and up to 100-150 m high, are known in Grodno and Novogrudok districts.
Rafts, mainly composed of Ordovician, middle Devonian, upper Cretaceous, Paleogene and Neogene rocks, occur throughout Belarus (Karabanov 2000). In Gorna near Khotimsk, eastern Belarus (fig. 11-41), aggregations of large erratic blocks (6-12 m thick), are expressed clearly in the landscape, forming hills about 50-60 m long and 7-11 m high (fig. 11-55). These rafts are built mainly with middle Devonian limestone, dolomite and occasionally clay.
The most extensive folded dislocations, some 30 km long and 8-11 km wide, which are among the largest in Europe, are located in the Peski area (fig. 11-41). According to Karabanov (2002b), the dislocations at Peski (Krasnoselsky quarry) are arcuate, formed by thrust folds up to 15-20 m high, and mainly built with dislocated upper Cretaceous chalk
U k r a i n e - Only small areas in the northwestern Ukraine were affected during the Oka glaciation, and the Dnieper glaciation overspread larger areas (fig. 11-56). The main ice lobe advanced into the Dnieper drainage basin and is marked by ice-marginal landforms including push moraines, eskers
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Distribution of glaciotectonism
187 These diapirs coincide with a N-S striking regional fault. The diapirs, which differ in amplitude and reach maximum vertical dimensions of 140-160 m, consist of Jurassic clay that has intruded Paleogene arenaceous-argillaceous beds and lower/middle Pleistocene alluvial deposits (see fig. 8-4). According to Matoshko (2004) the diapirs are characterized by steep dips (40-60°), microfolding, microfaults, increased fissuring and slickensides.
and sandurs, spillway valleys, till and fluvioglacial sediment, as well as by local glaciotectonic dislocations (Matoshko 2004). These mostly consist of thrust systems cutting sections of different rock types, and may be separated into isolated diapirs and recumbent folds with average amplitudes of 515 m. Bedrock dislocations form knolls, mounds, and ridges, as well as areas of rugged topography. Southeast of Kiev, the Shevchenko valley is particularly noteworthy (fig. 11-56). This buried valley trends north to south and southeast, reaching a total length of 200 km and width of 10-12 km. The Shevchenko valley is thought to have originated by ice pressure in a tectonic (fault) zone with subsequent lateral extrusion of rocks and removal of large rafts, which were transported more than 40-50 km away (Matoshko 2004).
Glaciotectonic patterns in Arctic Russia During the past 15 years, substantial and numerous glaciotectonic phenomena have been described across the Arctic terrain of northwestern Russia and northern Siberia (fig. 11-4) in connection with investigations of Quaternary stratigraphy and chronology (Astakhov 2004). Glaciotectonic structures and landforms extend from the Pechora Sea continental shelf, to the Taymyr Peninsula (Hjort, M611er and Alexanderson 2004), and eastward. Large morainic arcs and
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Figure 11-57. Location maps for northwestern Russia and the Pechora Sea (a), lower Pechora River region (b), and HongureiMarkhida vicinity (c). Large glaciotectonic structures involving Quaternary strata are exposed in river bluff sections at Vastiansky Kon' and Hongurei. Adapted from Henriksen et al. (2001, fig. 1).
Aber and Ber
188
Pechora region in northwestern Russia (fig. 11-57). Large bluff sections along the Pechora River expose extensive glaciotectonic deformation of Quaternary strata (Tveranger et al. 1998). Upthrust and disrupted strata include ~80 m of marine clay, lacustrine silt, fluvial sand, peat, and till; the base of glaciotectonic thrusting appears to lie in Holsteinian marine clay. Following glaciotectonic thrusting, a discordant till was deposited by an ice lobe of the Kara Sea ice sheet, which advanced southwestward and reached the Markhida moraine approximately 60,000 years ago (Henriksen et al. 2001).
hill-hole pairs are comprised of deformed Cretaceous and Tertiary bedrock, as well as chaotic agglomerations of glacial sediments and massive ice. Depth of structural disturbance reaches 400 m in some cases. Considerable uncertainty and debate has revolved around issues of the extent, number, and timing of glaciations of this vast region by ice sheets derived from the Barents Sea and Kara Sea centers. The entire region was glaciated by pre-Weichselian ice sheets. The most surprising and significant discoveries are that maximum ice limits of the last glacial cycle were achieved during the early/middle Weichselian (fig. 11-4), and late Weichselian glaciation apparently did not reach the mainland of northern Russia east of the Kola Peninsula (Mangerud, Svendsen and Astakhov 1999).
The late Weichselian glaciation terminated on the Pechora Sea continental shelf between Novaya Zemlya and the mainland (fig. 11-4; Svensen et al. 2004), where Gataullin, Mangerud and Svendsen (2001) identified three distinctive glacial limits (fig. 11-58). Bedrock of the Pechora Sea shelf consists of Cretaceous clay, silt and sand >2000 m thick. Sea-floor drilling revealed the upper 20-25 m of bedrock is
Several representative situations serve to illustrate the nature and conditions of glaciotectonism, beginning with the
WATER i D E P T H ii 0-2Om i il
'i ??/? ................. ,. . . . .
%:,.+.:,~:~ ...............
/il:~:~:i~~:~
...................
Kolguev
'. . . . . . . ~'~
:
!~, !i
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..
.
Y 7 "
~ E~ty
~
owland
l~', ' "~'"r/
~
0
50
100
~~
150
200km
inferred ice sheet limits submerged shorelines
Figure 11-58. Bathymetry of the Pechora Sea floor showing three glacial limits during late Weichselian glaciation. The Kurentsovo Line contains large ice-thrust bedrock ridges formed by an ice sheet derived from the Barents Sea. Takenfrom Gataullin, Mangerud and Svendsen (2001, fig. 4); reprinted with permission from Elsevier.
Distribution of glaciotectonism
189 I
r
SE ......~ ..........~ /
~NW
ESE,.,.:--~.....N...N.W
i
SSE
~ .........w
ESE'~
E ~W
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E
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.
,
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: : 1 1 : !
:
i
i : 1 1 1
:
~.~
.
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1
:,
,,~
~_~
i
=
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'
Figure 11-59. Geologic section of strata and structures exposed in cliffs and recorded in boreholes along the Ob River near Atlym, western Siberia. Symbols: 1 - glacial strata, 2 - lower Miocene to upper Oligocene terrestrial silt, 3 - middle Oligocene terrestrial sand, 4 - lower Oligocene to Eocene marine clay, 5 - Eocene diatomite, 6 - Paleocene marine clay, 7 - faults observed (a) and inferred (b), 8- borehole. The profile is ~15 km long; adapted from Astakhov, Kaplyanskaya and Tarnogradsky (1996, fig. 16). Copyright John Wiley and Sons Ltd. Reproduced with permission.
t 2.~
~ 40 °
? .0~
. (~)
•
L~k~,2 Sevasti~n site
£ iii~~ii!~i!!iiiiiii¸~iiii,~i;i;~!i?Ji!i~¸~:i'!:iI:~:i~i~i:!i!5~i ~i~ii!:iii~!;:~i;!:~i~:;i¸ii~!i:i~!i!i
B~uga}akh S-..1.y S:~i~e
C~-) G~aciotectonic ~h~ust f~::~atures
-~.
Loc~ions o stri~,ed boulders
:~
D i s p I a c e d ( f o l d e d and faulted} beds o~ T,-~rtiary coal~beatin,~ deposits, S.ogo River basin Former meltwater detia, lower Khorogor Riv~.r va}le ~.,
I
~ a s t ~i ~ a ~
,,/TL/'~ t,.le41wa~er ower~low chariness t
Infe~~ed ..di~ectio.n of i::ce~!ow //
Strt~c~u~al ,,',eat,u~(-~s in P~sleozoic be(:!
0
5
10
1 5k.m
Figure 11-60. Glaciotectonic thrust systems at Lake Sevastian and Belugalakh Bay, Tiksi area, eastern Siberia. General location map (left). Box (T) indicates position of detailed map of Tiksi region on right. Based on Grosswald et al. (1992, fig.
1).
Aber and Ber
190 commonly folded, faulted, and sheared into mylonite by glacier action. The middle glacial limit, Kurentsovo Line, is a series of sharp-crested ridges, 20-30 km long, 2.5 to 10 km wide,, ld 20-25 m high. Seismic data lead to the conclusion that ridges are composed of upthrust Cretaceous bedrock. The Kurentsovo Line represents an ice-margin related to the Barents Sea ice sheet. Moving east of the Urals into West Siberia, glaciotectonic structures and landforms are ubiquitous over great expanses, including large thrust systems, megablocks, and ice-shoved ridges (Astakhov, Kaplyanskaya and Tamogradsky 1996). A particularly good example is demonstrated in fiver bluff exposures and borehole data along the Ob River in the Atlym vicinity (fig. 11-59). Glaciotectonic disturbances involve a thick sequence of poorly consolidated Tertiary strata, which are overlain unconformably by multiple tills and outwash. Thrust zones can be projected 300 m deep, ice-shoved hills stand up to 100 m high above the local datum, and the zone of deformation can be traced across a distance of 25 km. This site is situated beyond the limit of Weichselian glaciation and must be associated with older ice advances. In eastern Siberia, Grosswald et al. (1992) documented series of ice-shoved hills and associated source basins (hill-hole pairs) in the Tiksi vicinity (fig. 11-60). Two glaciotectonic landscapes are located around Lake Sevastian and adjacent to Belugalakh Bay. Ice-shoved ridges consist of giant slices of black Carboniferous shales displaced toward the southwest and thrust over Permian bedrock. Topographic basins mark discrete sources of shale as embayments on the shore of Lake Sevastian. This same area also contains displaced coalbeating Tertiary strata, which have been folded and faulted. It is clear that glaciotectonic deformations across Siberia took place in frozen substratum under cold-based glaciers, as permafrost is generally 100s of meters deep. According to Astakhov, Kaplyanskaya and Tarnogradsky (1996, p. 165):
Frozen sediments, if clayey and~or icy, can readily deform, thus translating basal glacial stress into sliding of the entire glacier/sediment complex along subglacial shear zones. Buried sediments with high ice content may deform plastically under shear stress, just as glaciers behave. Likewise, watersaturated clay may remain unfrozen and plastic at subzero temperatures. This is especially true where clay contains saline pore water, as in West Siberia. Thus clayey and icerich strata may serve as d6collements beneath a cold-based ice sheet, and the dynamic sole of the glacier may shift down to such d6collements with consequent deformation of the substratum (fig. 11-61). Astakhov, Kaplyanskaya and Tarnogradsky (1996) concluded that large compressional
Figure 11-61. Sequential model for ice-push deformation in a permafrozen glacier bed showing velocity profiles. During ice advance, the dynamic sole of the glacier shifts down into the clay strata. Adapted from Astakhov, Kaplyanskaya and Tarnogradsky (1996, fig. 19). Copyright John Wiley and Sons Ltd. Reproduced with permission. disturbances of this type are commonplace across the vast territory of West Siberia. At this point in time, it is premature to attempt comprehensive mapping of glaciotectonic phenomena across the huge expanses of Arctic Russia. Nonetheless, certain common elements emerge from current knowledge about the region. Major glaciotectonic features generally involve thick sequences of weak or unconsolidated Cretaceous, Tertiary or Quaternary strata, the Carboniferous shale of the Tiksi vicinity being a notable exception. Glaciotectonic features are all situated in the outer zone of glaciation connected with ice sheets derived from Barents Sea or Kara Sea centers. Where these ice sheets advanced into regions with soft substrata, they likely became relatively thin with low-slope surface gradients (Tveranger et al. 1999). Under such conditions, whether warm or cold based, the ice sheets were quite mobile and dynamic, which favored glaciotectonism.
Chapter 12 Dynamism of Glaciotectonic Deformation Fundamental cause of glaciotectonism It now is clear that glaciotectonic deformations may take place in a variety of settings--in front of the glacier, beneath the ice margin, or under the center of a thick ice sheet. Deformations may arise during advancing, maximal, or recessional phases of glaciation. All manner of material from crystalline bedrock to loose Quaternary sediment in both frozen and thawed conditions may be affected. Topographic settings vary from rugged mountains to continental shelves. Ice-shoved material was usually removed from a basin and piled into a hill of some kind. However, the opposite is also known, where an ice-scooped basin was partly filled with material dislocated from surrounding uplands ( R u s z c z y f i s k a - S z e n a j c h 1978). In short, glaciotectonic phenomena may be expected wherever susceptible crustal strata were overridden or pushed by glaciers or ice sheets. Many factors have been cited as important or necessary conditions for glaciotectonic deformation to take place (see Table 11-1). Of these factors, most are related to local topography, substratum material, ice dynamics, or water. These factors often vary considerably over short distances and times. Only the first--lateral pressure gradient-operates everywhere beneath a glacier, regardless of the local nature of topography, substratum material, or ice movement. The lateral pressure gradient is the fundamental cause of glaciotectonic deformation (Rotnicki 1976; van der Wateren 1985). Glaciotectonic deformation takes place when the stress (= pressure) transferred from the glacier exceeds the strength of the material subjected to the stress. A glacier imposes two kinds of stress on its bed: 1) vertical stress due to static weight of the ice column (= glaciostatic pressure), and 2) drag or shear stress due to movement of the ice over its bed (= glaciodynamic stress). The combination of these leads to what Jaroszewski (1991, 1994) called the static-kinematic conception for glaciotectonism. Hart (1995b) identified four zones beneath an ice sheet based on expected basal shear stress and the likelihood for subglacial deformation (fig. 12-1). Most glaciotectonic deformation takes place in the marginal and equilibrium zones, where shear stress is greatest; less deformation develops in the intermediate and divide zones. The manner by which basal shear stress is transferred from the glacier to the substratum
is conditioned by the ice/substratum interface. mechanisms are possible (fig. 12-2).
Three
• No basal slippage and no transfer of shear stress into the substratum. This typically occurs in frozen-bed conditions, and all glacier movement is accomplished by internal ice deformation. • Basal sliding, again with no transfer of shear stress into the substratum. Basal melt water allows the glacier to slip over a rigid bed. • Subglacial deformation, in which displacement takes place within a deformable bed. Thickness and nature of the deformable bed vary considerably; it may be thawed or frozen and may be underlain by a discrete d6collement. Glaciostatic pressure is given by ice thickness (H in m) times density (0.9 g/cc for ice) divided by 10 (see Table 12-1 for symbols and units)" ~zi =
0.9H/10 = 0.09H in kg/cm 2
(1)
This stress equals ~90 kg/cm 2 for each 1000 m of ice thickness. The shear stress created by ice movement is much less, for most situations only 1-2 kg/cm 2, with maximum values up to 10 kg/cm 2 where ice is frozen to its bed or flows around bedrock knobs (Weertman 1961; van der Wateren 1985). 2
0,2
i ~
a
0
I:o
=
.
.
.
.
.
.
.
deformation
Figure 12-1. Idealized profile across an ice sheet showing glacier thickness, basal shear stress, and base of deformation. Four zones: 1 - marginal (0-20 km), 2 - equilibrium (20-200 km), 3 - intermediate (200-800), and 4 - ice divide (8001000 km). Note large vertical exaggeration and different scales for ice thickness and depth of deformation. Adapted from Hart (1995b, fig. 1).
1
9
2
A
b
Table 12-1. Symbols and units used in analysis of glaciotectonic deformation. H
!ee thick~ss (~ight)in m
T
Thickness of substratum in m
p
Pressure or stress: P~ = iit~static pressure, P~= i~ergranular pressure:, P~,,= hydrostatic pressure
o:
kg/cm ~= - I bar or 1 atmosp~re pressure
~ r m a l stress or: pressur,e., Stress ~ i t used ~ r e is kg/cmL !
• i
-
:.-:
:.-:
::
;
;;
.
.
.
.
or to ice dMde; parallel to slope of ice sudace
o.
~ress compor~nt oriented perpendicular (:~rma!) to a plane
S ~ a r stress developed! par:al'!el to a surface of displacement; meas~ed in kglcm~', S ~ a r stress due to ice mo~,~ment over substratum (= glacioT~.,~ dynamic stress) . . . . . . . .
%
Co~sive stre~th of a material (= co~sion}
0
Ar~le of plane (fault) re!ati~ to: the maximum ~rmat stress
t~ .
The lateral pressure gradient can be calculated easily for situations where ice thickness is known (fig. 12-3). Part of the glaciostatic load on the substratum at any point is transferred to a horizontal stress. The relationship between vertical (Crzi)and horizontal (Crx)stresses is given by Poisson's
.
~ r m a l stress in ~rizontai direction; trans~rse to ice margin
o.
.
distance inside the margin. This inequality of ice loading generates a lateral pressure gradient regardless of the direction or rate of ice movement.
ratio (v):
G ia:ciotectonic stress operatic~ ~rizontaIly ,at ice/substratum o~ ,,co~act, Combination of ~ateral pressure gradient (ZAer,)+ :t~.,
.
and Ber
r
~ r m a l stress in ~rtica! direction; stress d ~ to ice badir@ (=: glaciostatic pressure, ~,~), or weight of o~rburden strata (cr~,)
. . . . . . . . .
.
e
A~te of inter~l friction; ~30 ~ for ~ s t rocks .
.
.
.
.
.
.
.
.
.
.
.
.
(~x = ~
((~zi)
For soft sediments and granular materials, Poisson's ratio is close to 0.2, so this value will be used for further calculations (van der Wateren 1985). Thus: (Yx= 0"25~zi = 0.0225H
AH~.~= H~ -H,~
v
Poisson's ratio; describes t ~ ratb of t~rizontal bu~i~ to vertical s~rtening for vertically stressed rock. E×perimentaf val~ ~O.2 for most sedi~,~ary rocks
The vertical load of ice creates a glaciostatic pressure on the substratum, which is irrespective of ice movement. Assuming constant ice thickness over a level surface, the glaciostatic pressure would be equal and uniform in all directions. However, ice thickness is not constant. Ice thickness varies, particularly near the margin, where thickness increases rapidly from zero at the margin to several 100 m a short
(3)
Because the vertical load varies with ice thickness, so the horizontal stress at the glacier sole varies from point to point. The lateral pressure gradient between two points results from the difference (A) in horizontal stresses: m(~xl/2 = (Yxl - (~x2 -"
C # e ~ e in vat~ between points of obserwation, for exampme:
A
(2)
1-v
0"0225(H1- H2)
(4)
For example, consider positions 7 and 8 on the diagram (fig. 12-3) to calculate the horizontal stress difference between two points (Table 12-2). A(~xT/8= (~x7 - Crx8= 13.5 - 10.1 = 3.4 kg/cm 2, or A(~xT/8= .0225(H 7 - Hs)= .0225(150)= 3.4 kg/cm 2
(5) (6)
The horizontal stress differences are cumulative over a distance. In other words, the stress difference over a given interval is passed on and added to the stress difference of the next interval, such that the total horizontal stress difference is:
Figure 12-2. Schematic illustrations for three mechanisms of basal movement of a glacier over the substratum. A -frozen bed with no basal sliding (U F), B - thawed bed with basal sliding (U s), C - subglacial deformation (U D ), which may be either thawed or frozen. Based on Boulton (1996, fig. 1).
FROZEN BED
UNFROZEN ROCK BED
ON ON
Dynamism of glaciotectonic deformation
ICE
o,~,~,.~
"
193
.
.
.
.
Figure 12-3. Cross section showing a hypothetical ice margin with surface gradient and subglacial pressure conditions for 1-km intervals. The summation of lateral pressure differences at the ice margin equals 22.5 kg/cm 2, ignoring the drag of ice movement. Adapted from Aber, Croot and Fenton (1989, fig. 10-1).
",,~+~ X
~*,,~
dX m
.~ kg/cm2
-~
lr
"t
I :1 i Y--,AO" x =
AGxl/2
........"r
t +
AO'x2/3
+
AGx;/4
+
...
. . . . . . . . . . . . . . . . . . . . . . .
TT (7)
In the example (fig. 12-3), the total lateral pressure at the ice margin is 22.5 kg/cm 2. This stress is independent of ice velocity and is controlled entirely by differential ice load. Lateral pressure ideally should be cumulative over long distance, from the center of an ice sheet to its margin. This cannot actually happen, though, because neither ice nor subglacial strata are ideal materials without cohesion or internal friction. Therefore, the distance over which lateral pressure may build is probably limited to a few 10s of km or less. The lateral pressure gradient is greatest near the ice margin, where surface gradient is steepest. This explains why most glaciotectonic features are created at or within a few km of ice margins. Maximum cumulative lateral pressure is probably on the order of 25 kg/cm 2, as it would be rare for ice thickness to change by much more than 1000 m over a distance of 10 kin. For ice sheet interiors, thickness changes over distance are naturally much smaller, but are still present transverse to ice divides and generate small lateral pressure gradients. A thickness change of only 100 m over a distance of 10 km could produce a cumulative lateral pressure of >2 kg/cm 2. Although smaller in magnitude, interior pressure gradients may also produce glaciotectonic disturbances in appropriate substratum materials.
r
"
Table 12-2. Stress factors for idealized glacierprofile, points 7 and 8 (see fig. 12-3). Point
H
7
600
54
:8
450
40.5
.........
....
13.5 10.1
substratum in many situations. To this lateral pressure, the glaciodynamic stress ('Cice)caused by ice movement may be added, assuming ice movement is usually in the same direction as the lateral pressure gradient. Thus, the total glaciotectonic pressure ((yg,) imposed horizontally on the substratum is given by: (~'gt : ZA(~x --I-1:ice
(8)
Given maximum lateral pressure and shear stress values of about 25 kg/cm 2 and 10 kg/cm 2 respectively, this means that maximum glaciotectonic stress ((~gt) is '~35 kg/cm 2. This maximum pressure would be realized in only a few restricted situations, for example where a steep ice front advanced upslope, perhaps over permafrost or during a surge. Many glaciotectonic features were not created under such unusual conditions; lower horizontal stress of <10 kg/cm 2 to only 1-2 kg/cm 2 was apparently sufficient to produce glaciotectonic deformations under many circumstances.
Initiation of thrust faulting The total lateral stress in the example (22.5 kg/cm 2) is more than ten times greater than normal subglacial drag and more than twice the maximum subglacial shear stress. It is clear that lateral pressure produced by unequal ice loading is the principal force for causing glaciotectonic deformation in the
Thrust faults are the primary structures of ice-shoved hills and megablocks. A subhorizontal thrust, or d6collement, separates displaced material above from undisturbed strata below megablocks. Deformed material comprising ice-shoved
Aber and Ber
194 hills is usually stacked along a series of inclined thrusts. Thus, any attempt to explain glaciotectonic deformation must deal with the initiation of thrusting. The h o r i z o n t a l pressure caused by glaciostatic plus glaciodynamic stress generates shear stress within the substratum. Shear stress is what initiates thrust faulting. Consider for example a plane of potential failure dipping at 30 ° and passing 50 m below the ice margin (fig. 12-4). The total lateral pressure gradient is 22.5 kg/cm 2. There is no vertical ice load directly below the margin, and we may ignore ice movement. The vertical stress is given by thickness (T in m) times density of the overburden strata (~2.0 g/cc for watersaturated, unconsolidated sand) divided by 10:
The horizontal and vertical pressures create two stresses on the plane" 1. (Yn - stress oriented normal to the plane. 2. x - shear stress acting parallel to the plane. The magnitudes of o and 1: can be calculated with the Mohr stress equations (see chap. 3). Applying the Mohr stress formulas to the example (fig. 12-4) gives: (Yn ~- 13.1 and x = 5.4 kg/cm 2. Maximum ~ would be achieved on a plane dipping at 0 = 45 °, so: °z)/2 = 6.25 kg/cm 2
(10)
This explains why most glaciotectonic disturbances occur close to the surface and near to the ice margin. With little
=:
The ideal orientation for thrust faults is at 0 = 45 °, where maximum shear stress occurs. However, most thrust faults actually develop at about 0 = 30 ° or less. The reason for this discrepancy is resistance to faulting within the rock mass due to internal friction and cohesion. These factors can be demonstrated by pushing a book over a horizontal table top (fig. 12-5).
(9)
O s = 2.0T/10 = 0.2(50) = 10 kg/cm 2
'~max -" ((Yx -
overburden or thin ice, o z may be small compared to Gx, and ~max is large. As substratum depth or ice thickness increase, however, so does o z, and 'l~ma x becomes smaller. When °z = °x, shear stress is zero, and thrusting cannot develop. Where °z > °x, the stress system is reoriented, such that maximum stress is vertical rather than horizontal, and low- to highangle normal faulting may occur (Croot 1987).
The normal stress acting on the book is simply gravity; shear stress is the horizontal pressure required to move the book. These two stress vectors define an angle (¢), called the angle of friction; its magnitude is proportional to frictional resistance. A simple relationship exists between 0 and ¢, such that: ¢ = 90 - 20, or 0 = (90 - ¢)/2
(11)
As many rocks possess an internal angle of friction of about 30 ° (Kulhawy 1975), the 0 angle of thrust faulting is also typically around 30 ° . Now suppose the book is glued to the table. An extra shear stress will be necessary to break the bond before the book can move. This additional stress is equal to the cohesive strength ('Co) of the b o o k / t a b l e bond. x ° values for
GRAVITY
i
,,C: 0 iiO 0 "~
E
@ ¢;
0
~0 ZO ...............~........................ ...
kg/cm2
30
////
~z cn
Figure 12-4. Enlarged view of ice margin in previous figure, showing stresses associated with a plane of potential failure at 0 = 30 °, 50 m below the ice margin. Adapted from Aber, Croot and Fenton (1989, fig. 10-2).
Figure 12-5. Diagram illustrating angle of friction (0) and stresses for pushing a book across a table. Adapted from Aber, Croot and Fenton (1989, fig. 10-3).
Dynamism of glaciotectonic deformation
195
unconsolidated sediments are essentially zero, whereas most well-consolidated sedimentary rocks have values in the range 250 kg/cm 2 to >300 kg/cm 2 (Kulhawy 1975). Deformation of better-consolidated rocks actually takes place along pre-existing weak discontinuities, such as bedding planes, shale seams, claystone or lignite interbeds, and so on. Glaciotectonic pressure is not capable of deforming "solid" sedimentary rocks. Rather it is unconsolidated sediment or low-cohesion discontinuities within lithified strata that may fail. According to the Coulomb principle, thrust faulting will occur when the shear stress along a plane of potential failure equals the shearing resistance of the material (Hubbert and Rubey 1959):
Lithostatic pressure is normally supported largely by intergranular pressure. However, in some situations hydrostatic pressure may approach or equal lithostatic pressure, in which case the overburden material essentially floats on a high-pressure fluid cushion that has no cohesive strength ('c° = 0). As hydrostatic pressure reaches lithostatic pressure, the shear stress required for thrust faulting approaches zero. Suppose that a confined aquifer exists below the ice margin in the example (fig. 12-4), and this aquifer is fed from an upglacier water table within the ice 100 m above the substratum. This extra water head is equal to height of the water above the substratum (in m) times the density of water divided by 10: Ph = 100( 1)/10 = 10 kg/cm 2
I: = I: + (y tan ¢ o
n
Applying this formula to the example (fig. 12-2): ¢ = 30 °, "c = zero for unconsolidated sand, and (Yn= 13.1. I: required for thrusting is 7.6 kg/cm 2, which is more than the actual shear stress acting on the plane (5.4 kg/cm2). So, thrust faulting could not take place under ordinary conditions, yet such thrusts are commonly observed under seemingly similar conditions. o
The mechanism for thrusting large rock masses over long distances was a puzzle to geologists for many years. The necessary physical conditions were explained elegantly by Hubbert and Rubey (1959), and Mathews and Mackay (1960) quickly applied this explanation to ice-shoved features. The total lithostatic pressure (Pt) exerted at any level within the substratum below a glacier results from weight of the overburden strata plus the ice cover ((Yzs+ °z~)" This lithostatic pressure is comprised of two components-intergranular pressure (Pi) and hydrostatic pressure (ph). I n t e r g r a n u l a r pressure is a mechanical stress transmitted from solid grain to grain within the rock mass. Hydrostatic pressure is a fluid stress transmitted through the water column within connected pores and fractures in the rock mass. Below a glacier, the substratum is assumed to be completely saturated with ground water, so substratum pressure can be represented as: Pt = Pi + Ph
(13)
Hubbert and Rubey (1959) demonstrated that the critical shear stress for movement along a thrust fault depends only on intergranular pressure, such that" = Pi tan ¢~, or I: = (P- Ph) tan 0
(15)
(12)
(14)
For a given plane, Pt is equal to (y,, and we will assume "co is zero at the time of thrusting. Thus" I: - (o n - Ph) tan ¢
(16)
For the example (fig. 12-4), this gives: "c = (13.1 - 10) tan 30 = 1.8 kg/cm 2
(17)
Thrust faulting would take place readily under these conditions of shear stress, hydrostatic pressure, and material cohesion. No ice movement is required; the lateral pressure gradient combined with elevated ground-water pressure is sufficient to initiate thrust faulting. Any ice flow toward the margin simply augments the available shear stress. In a glacial context, elevated hydrostatic pressure can develop in two manners: 1) c o m p a c t i o n of i n c o m p e t e n t and impermeable strata, such as claystone or lignite, and 2) transmission of water into a confined aquifer under a pressure head. The former occurs wherever glacier advance loads a sequence of interlayered sediments, and the incompetent beds are compacted faster than water can escape. The latter is often observed around modem glaciers, where fountains, jets, and sudden floods of melt water emerge from below the ice margin (Gustavson and Boothroyd 1987). M a n y authors have e m p h a s i z e d the i m p o r t a n c e of overpressurized aquifers beneath the glacier margin as a key ingredient for glaciotectonism (e.g. Bluemle and Clayton 1984). Often this is combined with a narrow frozen marginal zone or permafrost to trap water in confined aquifers (Mooers 1990; Piotrowski 1993; Boulton and Caban 1995). In other cases, such as ice-contact glaciomarine and glaciolacustrine fans or deltas, glaciotectonic deformation takes place in completely thawed settings (fig. 12-6; Hart 1996).
Aber and Ber
196
Figure 12-6. Model for glaciotectonic deformation in an icecontact submarine fan or apron, based on Weichselian and Holocene fjord basins in Norway and Svalbard. Deformation at the ice margin (4) takes place under completely thawed conditions. Adapted from LOnne (1995, fig. 1); reprinted with permission from Elsevier.
I
FORESETBEDS BUILT OF RESEOtMENTE!D S U ~ C ~ A L T~LL ~ D OUTWASH MATER~AL{2} @ LOC-~.L.~ZEDSUSPENSIONFALLOUT FROM BUO¥~T MEL~ATER PLUME (:3)
@ G'LAC!OTECTON!C DE,FORMAT:~ON !N THE
I
HEAD ZONE .(4) COMMON~CE.,,~FTED DEBRIS (5)
~,~ ....~ ,,d
~..,, .....
!"
k LODGEMENT T~LL
B~ROCK T~RESHOLD
Thrust faults ideally develop at about 0 = 30 ° within homogeneous material. However, the position of thrusting is often controlled by a pre-existing weakness within the substratum: bedding plane, incompetent bed, lithologic boundary, unconformity, permafrost boundary, etc. Such discontinuities are approximately horizontal in many sedimentary sequences, so subhorizontal thrust faults are quite common (Jaroszewski 1991).
Continuation of thrust faulting The continued movement along a thrust fault requires maintenance of the critical shear stress and may involve overcoming increasing resistance to fault movement. Consider a flat megablock moving horizontally over level
terrain beneath a glacier (fig. 12-7). As long as the pressure conditions that caused initial thrusting continue to exist, the megablock may be displaced horizontally. Long-distance transportation over glaciated plains is possible (chap. 7). Many megablocks in western Canada were moved over and deposited on till of the same glaciation. Movement will cease when either hydrostatic pressure or glaciotectonic stress decreases, for whatever reason, or a topographic obstacle is encountered. Most megablocks were probably moved as individuals; that is, displacement of one megablock was independent of other megablocks in time and space. Continued movement on dipping thrust faults poses a different problem, for here the fault block is pushed up an inclined
Figure 12-7. Schematic diagram illustrating continued displacement of thrust blocks in two settings: horizontal subglacial megablock and inclined ice-marginal thrusts. Not to scale; adapted from Aber, Croot and Fenton (1989, fig. 10-4).
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Dynamism of glaciotectonic deformation ramp (fig. 12-7). Part of the thrust block--its toemmust be lifted above the original surface; this is what creates an iceshoved hill. The weight of the thrust toe represents increased resistance to fault movement, and the toe weight grows as thrusting continues. Thus, shear stress must be large enough to overcome initial shear resistance on the fault as well as to lift the thrust toe (van der Wateren 1985). The effect of toe weight explains several common observations about ice-shoved hills. First, the amount of vertical uplift is limited, for most large hills to between 100 and 150 m, with maximum vertical uplift ~200 m (Kupsch 1962). Given a thrust-plane dip of 30 °, only around 200 to 350 m of horizontal displacement is necessary along the ramp itself. This may at first seem like too small a figure for horizontal movement, but the ramp itself also may be displaced as part of a series of imbricated thrust blocks. The second effect of toe weight is to increase the load on the substratum below the front end of the thrust block. This will initiate a second, more distal thrust fault (fig. 12-7). When thrusting of the first block ceases, it will be carried forward on the back of the second thrust block, and so on, in an imbricated manner. In this way, composite ridges are constructed, in which proximal thrust blocks may be moved some km, whereas the most distal will move only a short distance.
197 During thrusting of composite ridges, individual thrust blocks undergo various internal deformations and may be pushed into steeper positions as they pile up in front of an advancing ice margin. Eventually the combined toe weights and internal resistance to thrusting of the composite ridge will exceed the glaciotectonic stress and thrusting will cease, Continued glacier advance may override the stabilized thrust zone, and the whole thrusting process could be repeated farther downglacier (fig. 12-8). The final effect of toe weight is to place severe limitation on subglacial ramp thrusting. In addition to lifting the toe weight, glaciotectonic stress must also displace the column of ice above the toe. It is simple enough to say that subglacial thrust blocks may be "incorporated" into the glacier (Moran 1971). However, it is mechanically quite difficult to develop sufficient stress to lift large thrust blocks beneath a glacier. Glaciostatic pressure, which is by far the greatest stress acting at the glacier sole, prevents uplift of more than a few 10s of m. Subhorizontal thrusting or development of low-angle normal faults is to be expected, rather than ramp thrusting, in subglacial settings (Croot 1987). So, whereas megablocks and small ice-shoved hills may be created subglacially, large ice-pushed hills can originate only at ice margins (van der Wateren 1992). According to LCnne (1995, p. 31-33), glaciotectonic
Another effect of toe weight is to cause collapse of the uplifted part of the thrust block. Once raised above the original surface, lateral support for the toe is lost, and the maximum stress (gravity) is vertical. Normal faults, slumps, fissures, and related structures may form, as a result of relaxation of glaciotectonic stress. Such secondary collapse structures are most common in the upper portions of ice-shoved hills and may be the only structures visible in shallow exposures. Some care for interpretation is advisable, as these structures may not represent the primary genesis of an ice-shoved hill.
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deformation is an inherent and important attribute of glacier termini ... major compressional deformation is probably the best diagnostic criterion for the recognition of an advancing glacier terminus. As an example, Bakker and van der Meer (2003) illuminated the role of thrusting for genesis of the Veluwe Ridge in the eastern Netherlands, based on surface exposures and shallow geophysical methods. Structural disturbances reach >200 m in vertical extent and stretch over a distance at least 12 km transverse to ridge trend. They identified three structural styles across the ridge complex (fig. 12-9).
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Figure 12-8. Multiple phases (I and II) of thrusting associated with continued glacier advance. DD - zone of diapirism. Compare with Fig. 11-22. Takenfrom Jaroszewski (1991, fig. 19).
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• Style I ~ proximal, ice-contact zone consisting of imbricated thrust faults dipping toward the presumed source basin at relatively high angle (20-25°). • Style II ~ inner proglacial zone containing both folds and low-angle thrust faults. Folds have wavelengths of ~250 m and amplitudes of about 40 m. • Style III ~ outer proglacial zone with long-wavelength (-400 m), low-amplitude (~30 m) folds, but lacking thrust faults. Toward the distal end, folds tend to flatten out and disappear. This large, ice-shoved ridge is composed of pre-existing Pleistocene fluvial strata, which make up the largest portion of the ridge, as well as penecontemporaneous glaciofluvial deposits. The structures were created in an ice-marginal and proglacial environment, in which deformation reached >5 km beyond the maximum limit of glacier advance.
Scale models of glaciotectonism Large ice-shoved hills are not known to be forming actively anywhere today, and so we have no modern analog for comparison with Pleistocene features. This is part of the reason for the variety of opinions on genesis of large Pleistocene ice-shoved hills. Fortunately, some scale models, both natural and artificial, may be relevant to ice-shoved hills in general. Gry (1940) made an early attempt to model icepush deformation by applying horizontal pressure on snow layers. And K6ster (1957, 1958, 1959) conducted numerous scale-model experiments related to glaciotectonic deformation. A major difficulty for such artificial models is correctly scaling down all physical factors, from sediment characteristics to rate of deformation, in order to produce a realistic simulation. Mulugeta and Koyi (1987) convincingly modeled the
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199
Dynamism of glaciotectonic deformation
"piggyback" style of thrusting. They subjected thin (0.1 to 0.2 mm) layers of fine (0.08 to 0.18 mm) sand to 40% bulk shortening in a conventional squeeze box. The deformed sand layers were then exposed by using a small vacuum nozzle to "erode" into the thrust mass. Three domains of deformation result under these conditions (fig. 12-10). The distal domain (a) consists of low-angle thrusts, between which thrust blocks develop overturned drag folds, extension fractures, and slumps. Thrust blocks are rotated into steeper positions in the intermediate domain (b), and back kinking develops in the upper folded portions of thrust blocks. The most severe deformation occurs in the proximal domain (c), where thrust blocks are rotated into near-vertical positions and are laterally compacted. Back thrusts or kink zones, that dip outward, result in underthrusting and further thickening in the proximal domain. This model was developed to analyze and demonstrate the kinematics of piggyback thrusting, not as an analog for a specific type of thrust terrain. Nonetheless, it represents a good approximation of glaciotectonic structures. Most iceshoved hills exhibit domain a and b features. Domain c features imply much greater uplift and lateral compression than is usually developed, although such features are present in the proximal portions of some composite ridges.
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The model deals strictly with deformation of pre-existing strata. In glaciotectonic settings, the ice and melt water also deposit new sediments, which may become intermingled with older material as deformation proceeds. Furthermore, continued glacier advance may overrun the ice-pushed hill, at which point a new stress regime combined with glacial erosion or deposition would be imposed. A life-sized artificial model of glaciotectonism was created as a byproduct of mining activity near Mach6w in southern Poland (Jaroszewski 1991). In this case, the waste dump from a sulfur mine loaded underlying Tertiary clay such that clay was thrust horizontally more than 500 m away from the dump front. Comparison of thrusting at the mine dump with ice-marginal glaciotectonic thrusting demonstrates remarkably similar structures formed by static loading of susceptible strata (fig. 12-11). The most realistic scale models of large ice-shoved hills are smaller natural ones that we can observe in connection with modern glaciers. For example, Eybergen (1987) investigated a push-moraine complex created in front of Turtmannglacier in Wallis, southern Switzerland. Turtmannglacier is a typical alpine glacier, which readvanced about 100 m during the 1971-86 period. The push moraine is comprised of small composite ridges, 5-11 m high, developed in lobate plan moreor-less parallel to the glacier snout. It is 100 m long east-
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Figure 12-11. Models for thrusting associated with static loading of poorly consolidated strata. A - mine dump (left) and thrusting of strata to right. 1 - Tertiary clay, 2 - Quaternary sand, 3 - mine waste pile. B - idealized glacier margin. Note the resemblance of structural styles. Sections not to scale; adapted from Jaroszewski (1991, figs. 12 and 14).
200 west and nearly 40 m wide. The push moraine is composed of till plus coarse, sandy gravel and rests on a drumlin of older lodgment till. The eastern portion of the push moraine contains mainly till derived from deforming older lodgment till and squeezing material from beneath the ice margin. Gravel dominates in the western portion; the gravel is derived from thrusting of at least partly frozen older outwash deposits and melting out of englacial and subglacial debris. Volume of the push moraine is estimated at 12,000 m 3, of which <1000 m 3 is material supplied from the glacier during the period of readvance (Eybergen 1987). Hence, the moraine is primarily a glaciotectonic feature consisting largely of deformed, older glacial sediments. During readvance of the Turtmannglacier, proglacial thrusting took place in both till and outwash gravel (fig. 1212). Thrusting was facilitated by development of increased pore-water pressure within the overridden sediment and in confined beds of the sediment wedge in front of the glacier.
Aber and Ber Piggyback-style thrusting migrated outward, with youngest deformation in the most distal portion. The structures correspond to domain a and b deformation. When the Turtmannglacier receded during the 1990s, the moraine was protected from melt-water erosion because of its position on a drumlin (fig. 12-13). Humlum (1983) and Krtiger (1994) described small festooned push-moraines that formed during a readvance in the 1980s in front of K6tluj6kull, an outlet glacier of M3~rdalsj6kull, Iceland (fig. 12-14). The moraine is composed partly of unfrozen floes thrust up from proglacial outwash and partly of supraglacial debris and mud-flow deposits derived from the glacier front. Following the advance of ~100 m distance, marginal retreat of the glacier in the 1990s left ice-cored moraines marked by steep-sided mounds and ridges (figs. 12-15 and 16). Initial thrusting took place in front of the advancing ice margin. The average angle of thrusting (0) is 26 °, giving an angle of friction (¢) of 38 ° for the sand and gravel sediment. This larger-than-usual angle of friction is because of the
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Figure 12-13. Photographs of the snout of Turtmannglacier. A - 1989, when the glacier margin was next to the push moraine. B - 1998, after the glacier had retreated to the left. Photos courtesy of J.J.M. van der Meer.
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Figure 12-16. Sketch map of the general pattern of moraines and related features at the margin of KO'tlujOkull. 1 - debriscovered ice, 2 - steep ice walls, 3 - groove, 4 - channel with supraglacial stream, 5 -frontal ice-cored ridges, 6 - push moraine, 7- ice-contact fan. Taken from Kriiger (1994, fig.
185). demonstrates proglacial thrusting of unfrozen and unconsolidated sediment with a rather high angle of friction.
Figure 12-15. Small push moraines in front of KO'tlujOkull date from minor ice advances of the 1950s (A - moss covered) and 1980s (B - bare). KO'tlujOkull has since retreated to right. Myrdalssandur outwash plain and the Atlantic Ocean are visible in the background. Photo by J.S. Aber (1994).
volcanic origin of the sediment; the grains are angular in shape and thus more abrasive than most other types of sedimentary grains. Once formed, individual floes were pushed up the ice front, as new floes appeared beneath, in a sequential manner (fig. 12-17). Meanwhile, supraglacial debris moved down the ice front, which resulted in composite interlayering of material in the push moraine. At its maximum advance, ice overthrust part of the moraine, before retreating. This situation
During subsequent melting of dead ice beneath the moraine in the 1990s, the terrain lowered some 25-30 m, and the once conspicuous push moraine became a single-crested ridge only 4-7 m high (Kriiger, Kjaer and van der Meer 2002). Evidence for its thrust origin was destroyed during melting and burial by supraglacial sediment, so that now it resembles a dump moraine left during a pause in glacier recession. In fact, it was formed by pushing and thrusting, and it marks the maximum extent of K6tluj6kull's advance during the 1980s. These various natural and artificial scale models demonstrate the kinds of structures and morphology that may be expected in larger Pleistocene ice-shoved hills, where both drift and bedrock were deformed. Most ice-shoved hills apparently were created close to or in front of ice margins, and therefore other ice-marginal processes must be taken into account when interpreting the origin of such hills. Megablocks and some small ice-shoved hills were formed subglacially. In many cases, features that originated in proglacial settings were later modified by overriding ice. The conditions of subglacial pressure, particularly hydrostatic pressure and the lateral pressure gradient, were of paramount importance for genesis of glaciotectonic phenomena.
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Chapter 13 Glaciotectonic Analogs Introduction Many structures within the Earth's crust are created in a manner analogous to glaciotectonic deformation. Wherever an advancing mass imposes an increasing load on weak substratum material, the conditions for deformation may develop both in front of the mass as well as beneath it. Glacier ice is only one kind of mass whose distribution and resulting crustal load may spread through time. Local increase of crustal loading may come about in many other non-glacial situations--growth of an alluvial fan or a delta, buildup of volcanic deposits, landslide or mass movement of material, and convergence of lithospheric plates. The scale of these analogous structures ranges from superficial cm-size deformations in soft sediments to major mountain ranges. In a general way, the size of structures and depth of disturbances are related to the size of the advancing load and the time interval during which that load was effective in reducing shear strength of the deformed material. Displacement above a d6collement with elevated hydrostatic pressure is an essential ingredient in all situations from ice pushing of small composite ridges to thrusting of major mountain ranges. Glaciotectonic deformation is not unique from this point of view; it is similar to many other crustal disturbances, the only significant differences being temporal and spatial scales.
mudlump uplift has migrated outward >2 km during the past century at South Pass distributary. This corresponds to southward extension of the distributary-mouth bar. Morgan, Coleman and Gagliano (1968) concluded that exposed mudlumps are spines on a series of linear diapiric folds that develop peripherally to the deforming load of the sand bar. Mudlumps range in size from pinnacles to small islands with maximum areas of about 8 hectares. They are mostly oval in form, with length usually 3-4 times the width. The surficial portion of a typical mudlump is formed by an anticline trending parallel to the mudlump's long axis (fig. 13-3). A narrow, shallow graben marked by many small normal faults runs along the crest of the anticline. Strata within the graben are highly irregular and confused. Active mudlumps rise at rates >60 cm per month (Morgan 1961). Extrusion of mud volcanoes and venting of methane-rich gas during uplift indicate that excess hydrostatic pressure must be developed within the clay cores of mudlumps. Mudlumps form only at the mouths of major distributaries; they do not develop in connection with lesser or shallow distributaries. Mudlumps undergo three developmental stages: 1) initial uplift as a submarine shoal, 2) growth into an island, and 3) erosion and truncation by waves. Many mudlumps experience episodic uplift, which invariably
Discussion of analogous non-glacial deformation provides information and ideas relevant to further understanding the genesis of ice-pushed structures. This is a two-way exchange, for knowledge of glaciotectonic features could likewise inspire the interpretation of non-glacial structural deformation (Pedersen 2005). Three glaciotectonic analogs of increasing size are presented in this chapter. First are mudlumps of the Mississippi Delta, which are similar in size to composite ridges. Glaciotectonic structures are next compared with thinskinned thrusting of mountains. Finally similarities between convergent plate-boundary and ice-push deformation are discussed.
Mississippi Delta mudlumps, Louisiana The Mississippi Delta is a classic, large, bird-foot style delta (fig. 13-1). Among its lesser known features are mudlumps-small islands or shoals formed over uplifted clay structures near bar-finger sands at the mouths of major distributary channels (fig. 13-2; Coleman 1988). Mudlumps are active features with life spans lasting a few decades before they become buried by prograding delta sand. The zone of
Figure 13-1. Landsat image of the Mississippi River and Delta vicinity, southern Louisiana. Dispersal of muddy water shown by intermediate gray tones within the river, passes, marshes, and coastal zones of the delta. Landsat MSS band 3 (near-infrared); date April 1983. Image acquired from EROS Data Center, U.S. Geological Survey.
Aber and Ber
204
Test drilling near South Pass distributary revealed the subsurface stratigraphy and structure of several mudlumps (fig. 13-4). Beneath bar-finger sand, which may reach up to 120 m in thickness, a sequence of prodelta clay units rests on an algal reef zone ~180 m deep. The algal reef zone is dated at 26,500 years ago, and a shell layer between clay units I and II is dated at 15,500 years ago. These dates indicate a late Wisconsin age for the prodelta sediments that were deposited during the late- and post-glacial rise in sea level. The clay units have been deformed into a series of asymmetrically thrust anticlines or diapirs with the algal reef zone acting as a structural basement. The zone of failure is situated in clay units above the algal bed. Clay strata exposed in mudlump islands contain foraminifera derived from at least 120 m or more in depth (Andersen 1961). The mudlumps of South Pass distributary have developed during the past 100 years, and thick sand has built up within the subsiding synclines between mudlumps during the same time. Assuming distributary sand accumulates near sea level, then synclines have subsided at average rates > 1 m/year (Morgan, Coleman and Gagliano 1968). This subsidence compensates for uplift in mudlumps. The asymmetry of folds, faults, and diapirs reflects differential loading by the accumulating sand mass.
i Figure 13-2. Sketch map of Mississippi Delta showing larger distributaries and zones of mudlump development. Based on Morgan, Coleman and Gagliano (1968, fig. 1); adapted from Aber (1988d, fig. 3). coincides with river flooding, when rapid sedimentation occurs at the distributary mouth. Some mudlumps have risen 3-5 m during a single flood cycle (Coleman 1988). Mudlump development, thus, is connected intimately with the sediment load created by seaward growth of distributaries (Morgan, Coleman and Gagliano 1968).
Figure 13-3. Schematic, transverse cross section showing surficial features of a typical mudlump. H.W.M. = high water mark; L. W.M. = low water mark. Adapted from Morgan (1961, fig. 6).
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Mudlump uplift proceeds sequentially in front of advancing distributary-bar sand in a manner akin to piggyback thrusting (fig. 13-5). Loading, compaction, and over-pressuring of the prodelta clay strata take place as thick sand accumulates at the delta front. Clay begins to flow outward and upward toward the zone of lower pressure. Asymmetric diapirs and thrusts develop in the clay, and uplifted anticlines emerge at the surface. Flowage of clay into the diapir core leaves a subsiding basin behind the mudlump that is filled by delta sand. Mudlump uplift continues until the original source of clay is depleted or until the mudlump is buffed by sand. The
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distribution of active mudlumps follows lobate patterns around the ends of prograding distributaries. Mudlumps are created by growth of a static sand load; no horizontal pushing or drag is exerted by the sand on underlying clay strata. The driving force for deformation is the lateral pressure gradient produced by unequal loading of relatively heavy sand over relatively light and plastic clay. The pressure conditions at the base of sand are easily calculated, according to the method of Bouwer (1978), assuming maximum sand thickness of 120 m, grain density of 2.6 g/cm 3, sediment porosity of 50 percent, and no overpressuring within the sand. Total pressure equals 21.6 kg/cm2; hydrostatic pressure is 12.0 kg/cm 2, and intergranular pressure equals 9.6 kg/cm 2. The latter represents the increased load imposed by the sand, which resulted in collapse and deformation of the subjacent clay. The size, style, and pattern of Mississippi mudlump deformation is quite comparable to composite ridges generated at the margin of an advancing ice lobe. Prograding distributary-bar sand serves the same role as advancing ice. In both cases, growth of static loads generates lateral pressure gradients that are responsible for deformation around the
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frontal and lateral margins of the ice lobe or distributary bar. The advancing load may eventually overspread the earlierdeformed structures. Mississippi mudlumps demonstrate that structural features closely similar to composite ridges may be created entirely by differential loading in a totally thawed environment (Aber 1988d).
Thin-skinned thrusting There are many examples of tectonic deformation involving displacement of surficial sedimentary strata over basement rocks by orogenic processes, so-called thin-skinned thrusting. The most intensively studied mountains include the Canadian Rockies (Bally et al. 1966; Dahlstrom 1970, 1977; Perry, Roeder and Lageson 1984), Appalachians (Rich 1934; Harris and Milici 1977; Perry, Roeder and Lageson 1984; Thomas 1990), and Alps (Graham 1978; Lemoine 1973; Siddans 1979, 1984). At a descriptive level, many close similarities exist between glaciotectonic landforms, particularly composite ridges, and orogenic thrust masses. These similarities extend beyond the morphological comparisons made by Gripp (1929), to similarities of internal structures,
Table 13-1. Structural styles, dominant stresses, and transport modes for thin-skinned deformation in glacial settings. Based on Harding and Lowell (1979) and Lowell (1985); adapted from Aber, Croot and Fenton (1989, Table
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Figure 13-5. Schematic model for structural development of mudlumps during advance of delta sand from left to right. Short arrows show direction of clay flowage; long arrows indicate differential loading by sand. Symbols same as fig. 13-4; based on Morgan, Coleman and Gagliano (1968, fig. 23F); adapted from Aber (1988d, fig. 5).
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Figure 13-7. Thrust system of LOnstrup Klint, northwestern Denmark. Listric, imbricated faults increase in dip from the distal (right) end toward the proximal (left) end. Ice pushing from north to south. Adapted from Pedersen (1987, fig. 10). Compare with Fig. 1-4. geometry of the systems, and models proposed to explain their genesis. As in glaciotectonic settings, it is common to find orogenic thrusting and folding associated together in the same system, with complex spatial and temporal relationships between each style of deformation (Dahlstrom 1977; Lowell 1985). The overlap between styles and geometries of movement enable us to look beyond the level of descriptive comparison of individual elements to whole thrust systems and to a variety of models for their development (Table 13-1). Many composite-ridge systems are composed of imbricated sheets of sediment displaced by glacial activity from flatlying strata at or just beyond the ice margin. Such imbrication is also the hallmark structural framework of many montane foreland zones (figs. 13-6 and 7), in which thrust faults are the dominant structural features (fig. 13-8). This gives rise to a starting point for the thin-skinned thrust analogy. Other common elements include the following features (fig. 13-9).
Figure 13-8. Tilted and deformed sedimentary strata of the Madera Formation (middle Pennsylvanian) are capped by crystalline rocks (early Proterozoic), Sange de Cristo Mountains, southern Colorado. Rocks of the Madera Formation are brecciated and contorted in the fault zone which marks the Culebra Thrust Fault (Wallace and Lindsey 1996). Forested range in background is part of a huge crystalline thrust sheet. Photo by J.S. Aber (2003).
• A basal d4collement, usually a weak zone or horizon, from which all thrusts propagate forward and upward (Boyer and Elliott 1982). In many cases, multiple d6collement levels may exist for incompetent layers within a thick sedimentary sequence. • Major thrust surfaces are bed-parallel in less competent strata forming flats, and cut upsection across competent
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As illustration of typical orogenic thrust systems, the southern Appalachian Mountains serve well (Thomas 1990). The thrust belt of northeastern Alabama is an allochthonous mass of Paleozoic strata displaced over Proterozoic crystalline basement (fig. 13-10). The regional d6collement is near the base of the Paleozoic sequence in Lower and Middle Cambrian fine-grained clastic rocks; massive Upper Cambrian and Lower Ordovician carbonates form a stiff unit that dominates thrust sheets. Structural features trend/strike
• Individual thrust sheets are laterally discontinuous along strike (limited axial length), and form ridges that rise and die out again. Lateral ramps or tear faults bound the margins of thrust sheets. Such discontinuties are transverse zones (Thomas 1990).
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frontal and lateral ramps, and associated folds, which are integral components of the overall kinematic system (Thomas 1990).
• Domain I - multiple-level thrust sheets stacked in imbricated and duplex structures.
The role of pre-existing basement faults is recognized here and elsewhere for controlling geometry of orogenic thrust belts. Basement faults of the southeastern United States originated during the late Proterozoic as normal faults and rift zones that were reactivated during later tectonic events (Marshak and Paulsen 1996). In this case, well-documented northeast-striking, down-to-southeast basement faults underlie major frontal ramps in the thrust belt (fig. 13-10). This situation is analogous to glaciotectonic thrusting of soft strata above buried hard obstacles, such as demonstrated by Trzebnica Hills, Poland (see fig. 11-17).
• Domain II - folds associated with frontal thrust ramps; structural relief >4000 m. • Domain III - broad, flat-bottomed synclines and narrow, asymmetric anticlines; structural relief <3000 m. These three zones bear a striking resemblance to the three glaciotectonic styles of composite ridges (see fig. 12-9). Parallel to strike of the thrust belt, faults and folds are offset within the Anniston transverse zone, which extends across the entire thrust belt in a band ~30 km wide. Thrust sheets are truncated by lateral ramps, and folds are deflected across the transverse zone (fig. 13-12). The thrust belt, thus, comprises a three-dimensional network of d6collement fiats,
In order to explain the genesis of the Appalachians, Alps, and other thrust mountains, geologists employ three main mechanistic models: 1) gravity sliding, 2) gravity spreading, and 3) pushing from the rear (fig. 13-13). The applicability of such models has been debated vigorously by Elliott (1976),
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Chapple (1978), and Siddans (1984) among others, and Pedersen (1987) provided a useful introduction to the debate from another point of view. G r a v i t y sliding or gliding is a mechanism in which deformation is caused by movement of the sediment package downslope under its own weight (Pedersen 1987). In a glaciotectonic setting, such a model implies that deformation of sedimentary strata in front of an advancing glacier occurred because the proximal zone was uplifted, for example as a forebulge, providing a surface gradient away from the glacier. As far as we are aware, few structures of this origin are observed in glacial settings, and this mechanism will not be considered further.
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G r a v i t y s p r e a d i n g is also induced by a change in gravitational forces acting on a once stable sediment package. In this case, a normal (static) load applied by an increased weight (ice, sediment or rock) is translated into a lateral pressure gradient, which results in squeezing the adjacent sediment away from the load and upward (Elliott 1976; Pedersen 1987, 2005). Such increases in weight may be caused by glacier loading, sediment accumulation, volcanic buildup, orogenic loading, or even by dumping man-made waste (see fig. 12-11; Ruszczyfiska-Szenajch 1986). The style of deformation in sediments adjacent to the load point is cylindrical or listric (curved, concave-upward) thrusting that dies out with distance from the load (Jaroszewski 1991). In a glaciotectonic setting the implication is that the major stress is due to differential ice loading (normal stress), not drag created by glacier advance (Rotnicki 1976). The final model, pushing from the rear, was proposed by Chapple (1978) as an explanatory model for orogenic structures identical in style to those that Elliott (1976) explained by gravity spreading. Pushing-from-the-rear models, which at first glance appear to be most logical, have met with considerable criticism from structural geologists. Many have argued consistently that thin, broad sheets of sediment cannot transmit sufficient stress to remain undeformed while undergoing thrusting. Chapple's mechanistic treatment appears to contradict this, and he argues that compressive stress (by a pushing movement) is a fundamental feature of thrust belts (1978, p. 1196). His model includes a weak basal layer and a wedgeshaped sedimentary package. The implication for glaciotectonic deformation is that the forward motion of the glacier is critical, not the static ice load. In other words, the glacier acts like a bulldozer, pushing and piling up material at the ice margin during its advance (van der Wateren 1992). Siddans (1984) considered each of these models in different configurations of surface slope and sediment package in an alpine orogenic setting, and he argued that each model is mechanically possible and could produce tectonic deformation, providing the underlying assumptions are
210
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plates, whose sizes, shapes, positions, and boundaries are constantly changing. The lithosphere is composed of the crust plus the uppermost mantle with an average thickness of 100 km, thicker under continents and thinner under oceans. weak/incompetent layer
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Plates interact along their mutual boundaries, which may be either divergent, convergent or transform. Mountain systems are created where convergent boundaries develop within or next to continental lithosphere. Convergent boundaries involving continents take two general forms (Molnar 1986).
layer
Figure 13-14. Ice-marginal thrusting in relation to depth of a potential ddcollement along a weak layer. Thrusting is more likely where the weak layer is relatively close to the surface compared to thickness of the advancing glacier margin. Taken from Aber, Croot and Fenton (1989, fig. 1112).
fulfilled. In a glacial setting, several critical factors may influence the mechanism of deformation. • Slope of the proglacial area, toward or away from the ice margin. • Depth to a suitable weak layer, a potential drcollement (fig. 13-14). • Relationship of glacier snout to proglacial sediment. • Thickening of the proximal end of the sediment package. Croot (1987) argued that pushing from the rear explains the piggyback development of composite ridges in Iceland; whereas, van der Wateren (1985) and Pedersen (1987, 2005) considered the gravity-spreading model most appropriate for the same style of glaciotectonic deformation. Van der Wateren (1992) emphasized that the glacier margin must be substantially thicker and steeper than the ice-shoved hill for the bulldozer effect to function. In all likelihood, pushing from the rear and gravity spreading operate together, to a greater or lesser degree, as both derive from the gravitational instability of the glacier. So, many composite ridges may be considered as results of both mechanisms acting in concert, which fits well with Jaroszewski's (1991, 1994) statickinematic concept for glaciotectonism.
• Subduction z o n e - where thin oceanic lithosphere descends below a continental plate and a volcanic mountain chain develops on the overriding continent. Such mountains are supported by deep, but relatively warm and weak crustal roots. The Andes Mountains of western South America represent a prime example. • Collision zone - where thicker oceanic or continental lithosphere moves into a subduction zone and thrusts under the overriding continental plate. Such mountains are held up by thinner, but colder and stronger crust. Typical examples include the Alps and Carpathians of western and central Europe and the Appalachians of eastern North America. Both types of convergent boundaries are comparable in a general way to ice-pushing of composite ridges. Ignoring the great differences in time span of development, lithology of affected rocks and size of structures, an overriding continent behaves like an advancing ice sheet--where an obstacle is encountered, thrusting of that obstacle may occur in front of and below the leading edge of the continent or ice sheet. Two primary arcs of deformation chararcterize fully developed subduction zones (fig. 13-15). • Hydraulic arc - low temperature/high pressure, blueschist and greenschist metamorphism. Small sedimentary volcanoes are fed by cool diapirs of mrlange, salt, shale, and fluid derived from the accretionary wedge. Mud seamounts, composed of serpentine and fluids derived from the subducted lithosphere, also may be created by diapiric extrusion or fault uplift.
Convergent plate boundary
• Magmatic arc - high temperature/pressure, amphibolite and migmatite metamorphism. Large volcanoes are fed by intrusions of intermediate (andesitic) silicate magma derived from partial melting within the subducting and overriding plates.
The Earth's largest topographic features--mid-ocean ridges, deep-sea trenches, and mountains--are the results of plate tectonics. The rigid outer shell of the Earth, called the lithosphere, moves slowly as several large and many small
In a glacial context, the accretionary wedge is most relevant (van der Wateren 1992). The advancing glacier serves the role of the continental plate, which scrapes material off the oceanic plate and builds an accretionary wedge on its leading
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edge. This wedge is a m61ange of rock and sediment stripped off the subducting plate. High-pressure fluids expelled from the subducting plate give rise to various diapirs, mud volcanoes and springs (fig. 13-16). The accretionary wedge gains mass both at its front and on its underside; the latter process is called underplating. Similar mechanisms operate in ice-shoved hills, where excess ground-water pressure facilitates thrusting, feeds diapirs and intrusions, and may erupt as hydrodynamic blowouts (Bluemle 1993). As the source basin is deepened and lower d6collements are activated, the ice-shoved mass may gain material in a manner analogous to underplating (fig. 13-17). Both types of convergent plate boundaries as well as the glaciotectonic analogy are demonstrated by development of the Himalayan Mountains and Tibet Plateau. At the beginning of the Cenozoic (fig. 13-18A), India was moving Distance J:rom,T}enc!s A:×is (km i} 50
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rapidly northward approaching Asia. The two continents were separated by an ocean basin--the Tethys Sea, which was closing due to subduction beneath Asia. Volcanic mountains grew above the subduction zone and the crust was greatly thickened by intrusions to form most of what is now the Tibet Plateau. Along the southern edge of Tibet, landderived sediments along with oceanic material, scraped off the subducting plate, built up an accretionary-wedge complex. Meanwhile, continental-shelf sediments continued to accumulate along the northern edge of India. Approximately 50 million years ago (fig. 13-18B), India began to move into the subduction zone and underthrust Tibet. Owing to its low density, continental crust cannot be subducted deeply, so subduction and accompanying volcanism soon ended. However, India has continued moving northward relative to Asia. This continued movement is facilitated in part by southward thrusting of the Himalayas. Major uplift of the Himalayas was underway by 35 million years ago (fig. 13-18C), at which time the Main Central Thrust was active. This thrust transported rocks of the Tibetan accretionary wedge and Indian continental shelf as well as Indian basement rocks.
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~ ................................ .................................................................... ..... i! ~,~,, o,, .......o,,,, ~oo,,..... Figure 13-17. Schematic model for underplating during development of an ice-shoved hill, as the source basin is deepened and a lower ddcollement becomes active. Adapted from van der Wateren (1992, fig. 5.7).
212
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to ice-pushing. Tibet fills the role of an ice sheet, whose forward movement imposed an increasing load on the substratum. The accretionary-wedge complex is analogous to ice-contact and proglacial sediment deposited during glacier advance. India represents pre-existing strata that were thrust in front of the advancing ice, and lithospheric depression occurred in response to both tectonic and glacial loading.
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F i g u r e 13-18. S c h e m a t i c cross s e c t i o n s s h o w i n g developmental stages of the Himalayas and Tibet Plateau: A - 60 million years ago, B - 50 million years ago, C - 35 million years ago, D - 20 to 10 million years ago. The Himalayas have been thrust southward as India underthrust Tibet toward the north. Based from Molnar (1986); not to scale, vertical dimension approximately 75 km.
Sometime between 20 and 10 million years ago, this thrust stopped moving and a new, deeper fault, the Main Boundary Thrust, became active (fig. 13-18D). This fault movement uplifted rocks of the first thrust sheet still more and also overthrust mountain-derived sediment on the Ganges Plain. Indian lithosphere was depressed by the added load of these thrust sheets. Major uplift and considerable erosion of the Himalayas took place in the Miocene and again in the Pleistocene (Amano and Taira 1992; Clark et al. 2005). As a result of thrusting, many Himalayan peaks are capped with sedimentary rock that originated in the Tethys Sea. This indicates minimum vertical uplift, not counting eroded material, of at least 8 km. Himalayan underthrusting was produced by India's northward movement against the Tibet Plateau, or conversely the advancing load of thick Tibetan crust caused overthrusting of the Himalayas. The latter point of view shows the similarity
India's northward movement is also accommodated in another, quite different way by displacement along major transform faults to the east, north, and west (Tapponnier et al. 1982). North-south transform faults bound the eastern (Burma-Bay of Bengal) and western (Pakistan-Arabian Sea) sides of the Indian plate. These boundary faults are analogous to the lateral margins of certain ice-pushed hills, where tear faults, low ridges or elongated drumlins are developed parallel to ice movement and transverse to composite ridges (chap. 4). However, the mosaic of faults and rift zones that disrupts a vast region extending north of India, across China, as far as Lake Baikal in eastern Russia does not have any glaciotectonic comparison. In a manner similar to the Himalayas, the southern Appalachians and Ouachita Mountains of eastern North America were built during the Paleozoic by continental collisions involving ancestral North America and Gondwana. During the collisions, thick thrust sheets were pushed 100s of km onto the American continent. The thrust sheets include igneous and metamorphic rocks along with a large volume of sedimentary rock that originated on the former continental margin. Prior to thrusting, the sedimentary strata were porous, full of fluid, and contained abundant hydrated minerals. Oliver (1986) speculated on the fate of fluids involved in Appalachian and Ouachita thrusting. Some of the fluid migrated up into the metamorphic interior of the mountains; some escaped in springs along the mountain front; some remained in d6collements to facilitate thrusting; some may have descended into the basement; and some fluid was expelled into permeable sedimentary strata of the foreland beyond the mountains. These latter fluids carried heat, dissolved minerals and hydrocarbons into the mid-continent region with widespread geologic consequences. The zonation of coal metamorphism, hydrocarbons, and secondary mineralization shows definite relationships to the mountain belts (fig. 13-19). Oliver (1986) compared the action of the thrust sheet to a giant squeegee or roller that drives fluids ahead as it advances. This situation is analogous to the outer zone of glaciation, where ice sheets thrust poorly consolidated and saturated sedimentary strata. Ground water is forced to flow from under the glacier toward the ice margin. This fluid is driven out of
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From a mechanical point of view, creation of major thrust mountains along convergent plate boundaries is really no different from thrusting of ice-shoved hills. As these examples demonstrate, young and old mountain systems around the world exhibit structural styles and basement controls that mirror composite ridges created by glaciation, which amply confirms the notion that ice-shoved hills are natural scale models of orogenic thrust belts (Hopkins 1923; Berthelsen 1979).
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References Aarseth, I. and Mangerud, J. 1974. Younger Dryas end moraines between Hardangerfjorden and Sognefjorden, western Norway. Boreas 3, p. 3-22. Aber, J.S. 1979. Kineto-stratigraphy at Hvideklint, MOn, Denmark and its regional significance. Geological Society Denmark, Bulletin 28, p. 81-93. Aber, J.S. 1982. Model for glaciotectonism. Geological Society Denmark, Bulletin 30, p. 79-90. Aber, J.S. 1985. The character of glaciotectonism. Geologie en Mijnbouw 64, p. 389-395. Aber, J.S. 1988a. Structural geology exercises with glaciotectonic examples. Hunter Textbooks, WinstonSalem, North Carolina, 140 p. Aber, J.S. 1988b. Spectrum of constructional glaciotectonic landforms. In Goldthwait, R.R and Matsch, C.L. (eds.), Genetic classification of glacigenic deposits, p. 281-292. A.A. Balkema, Rotterdam. Aber, J.S. 1988c. Bibliography of glaciotectonic references. In Croot, D.G. (ed.), Genesis of glaciotectonic phenomena, p. 195-210. A.A. Balkema, Rotterdam. Aber, J.S. 1988d. Ice-shoved hills of Saskatchewan compared with Mississippi delta mudlumpsmImplications for glaciotectonic models. In Croot, D.G. (ed.), Genesis of glaciotectonic phenomena, p. 1-9. A.A. Balkema, Rotterdam. Aber, J.S. 1991. The glaciation of northeastern Kansas. Boreas 20, p. 297-314. Aber, J.S. 1992. Glaciotectonic structures and landforms. Encyclopedia of Earth System Science, vol. 2, p. 361-378. Academic Press, San Diego. Aber, J.S. (ed.) 1993a. Glaciotectonics and mapping glacial deposits. Canadian Plains Research Center, Canadian Plains Proceedings 25, vol. 1, 310 p. Regina, Canada. Aber, J.S. 1993b. Expanded bibliography of glaciotectonic references. In Aber, J.S. (ed.), Glaciotectonics and mapping glacial deposits. Canadian Plains Research Center, Canadian Plains Proceedings 25, vol. 1, p. 99137. Regina, Canada. Aber, J.S. 1993c. Geomorphic and structural genesis of the Dirt Hills and Cactus Hills, southern Saskatchewan. In Aber, J.S. (ed.), Glaciotectonics and mapping glacial deposits. Canadian Plains Research Center, Canadian Plains Proceedings 25, vol. 1, p. 9-35. Regina, Canada. Aber, J.S. 1999. Pre-Illinoian glacial geomorphology and dynamics in the central United States, west of the Mississippi. In Mickelson, D.M. and Attig, J.W. (eds.), Glacial processes past and present. Geological Society America, Special Paper 337, p. 113-119. Aber, J.S. 2003-04. Applications of small-format aerial photography in North Dakota. North Dakota Geological Survey, Newsletter 30/2, p. 116-19. Aber, J.S. 2006. Ice-thrust terrain of the Missouri Coteau in
southern Saskatchewan. In Schlichtmann and Sauchyn (eds.), Centennial geography of Saskatchewan. Canadian Plains Research Center, Canadian Plains Proceedings [in press]. Aber, J.S. and Aarseth, I. 1988. Glaciotectonic structure and genesis of the Herdla Moraines, western Norway. Norsk Geologisk Tidsskrift 68, p. 99-106. Aber, J.S., Abdelsaheb, I., Nutter, B., Denne, J.E. and MacDonald, W.D. 1988. Composition, paleomagnetism, and age of the Kansas Drift. Kansas Academy Science, Abstracts 7, p. 1. Aber, J.S., Bluemle, J.R, Brigham-Grette, J., Dredge, L.A., Sauchyn, D.J. and Ackerman, D.L. 1993. Glaciotectonic data base and mapping of North America. In Aber, J.S. (ed.), Glaciotectonics and mapping glacial deposits. Canadian Plains Research Center, Canadian Plains Proceedings 25, vol. 1, p. 177-200. Regina, Canada. Aber, J.S., Bluemle, J.R, Brigham-Grette, J., Dredge, L.A., Sauchyn, D.J. and Ackerman, D.L. 1995. Glaciotectonic map of North America, 1:6,500,000. Geological Society of America, Maps and Charts Series, MCH079. Aber, J.S., Bluemle, J.R, Sauchyn, D.J. and Ackerman, D.L. 1991. Great Plains glaciotectonics. North Dakota Geological Survey, Miscellaneous Map 31, scale 1:2,500,000. Aber, J.S., Croot, D.G. and Fenton, M.M. 1989. Glaciotectonic landforms and structures. Glaciology and Quaternary Geology Series, Kluwer Academic Publishers, Dordrecht, Netherlands, 200 p. Aber, J.S. and Gatgzka, D. 2000. Potential of kite aerial photography for Quaternary research in Poland. Geological Quarterly 44, p. 33-38. Aber, J.S. and Kalm, V. 2001. Remote sensing of eskers from Vormsi and V~iinameri vicinity, northwestern Estonia. Geological Quarterly 45, p. 365-372. Aber, J.S. and Lundqvist, J. 1988. Glaciotectonic structures in central Sweden and their significance for glacial theory. G6ographique Physique et Quaternaire 42, p. 315-323. Aber, J.S. and Ruszczyfiska-Szenajch, H. 1997. Glaciotectonic origin of Elbl~g Upland, northern Poland, and glacial dynamics in the southern Baltic region. Sedimentary Geology 111, p. 119-134. Aber, J.S., Ruszczyfiska-Szenajch, H. and Krzyszkowski, D. 1995. Landsat interpretation of glaciotectonic terrain and regional lineaments in Poland. Quaestiones Geographicae, Special Issue 4, p. 1-11. Aber, J.S., Sobieski, R., Distler, D.A. and Nowak, M.C. 1999. Kite aerial photography for environmental site investigations in Kansas. Kansas Academy Science, Transactions 102, p. 57-67. Aber, J.S., Spellman, E.E. and Webster, M.R 1993. Landsat remote sensing of glacial terrain. In Aber, J.S. (ed.), Glaciotectonics and mapping of glacial deposits. Canadian
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Subject and geographic index alluvial delta/fan 203-205 Anaktuvuk River glaciation 167 Antarctica 1, 130
Antiquity of Man l, 2 Archean 11 Asia Himalaya Mountains 211-212 Kara Sea/shelf 153 Tibet Plateau 211-212
Laundry Hill 105-106 Oldman River 105-106 Monitor 69 Neutral Hills 69 southern Alberta megablocks 104-107 Sundance Power Plant 142 Thickwood Hills 101 Wolf Island 105-106 Wolf Lake/Hill 48, 49 Canada, British Columbia 42
balanced cross section 35-37, 40, 41 pin point 35, 41
Canada, general 4, 15, 125, 153, 158, 167-168, 196 Alberta Plain 67-68, 107 Hudson Bay 125 Manitoba Plain 79-80 Prairie region 153, 170 Rocky Mountains 63, 205-206 Saskatchewan Plain 67-68, 79-80, 103
Barents Sea ice sheet 154, 188, 190
Canada, Magdalen Islands 156
Belarus 45, 136, 154, 157, 168, 170-171, 176-177, 185186 Chashniki Plain 177 Gorna 177, 185-186 Grandichi 177 Grodno Upland 137, 177, 186 Melovara Gora 177 Minsk 186 Neman River 177 Novogrudok 177, 186 Orga 177, 186 Pyshki 177 Sopotskin 177 Ushahi Lakes 177 Volkovysk (Peski) 177, 185-186
Canada, Manitoba Brandon Hills 63, 79-82 Duck Mountain 80 Little Souris River 80 Riding Mountain 80 Tiger Hills upland 79-80, 82
Atlantic Ocean 125, 130, 158-159, 200
borehole (see well) Burma 212 Cambrian 109-110, 135, 137-138, 207 Canada, Alberta l, 4, 45, 69, 102, 104-106, 141-143 Catchem 105-106 Cooking Lake 101 Driftwood Bend 105-106 Highvale Coal Mine 142-143 Killarney Lake 69 Kipp 105-106
Canada, New Brunswick 125-128 Heath Steele Mine 141 Miramichi earthquake 128 Saint-Jacques 126 Saint John 126-127 Canada, Nova Scotia 126-127, 156 Cape Breton Island 126 Canada, Northwest Territories Bell and Boothia arches 125 Bylot Island 60, 62 Foxe-Baffin crustal block 125 Foxe Channel 125 Canada, Ontario 42, 125 Canada, Quebec 125-128, 158 Charlevoix seismic zone 128 Chicoutimi earthquake 128 Montreal 11 T6miscouata-Madawaska Valley 126-128
236 Canada, Saskatchewan 4-5, 69, 141-142, 168 Chaplin Lake 69 The Coteau 69 Crestwynd 28 Dirt Hills & Cactus Hills 29, 63, 67-72, 162, 182 Eagle Hills 145 Esterhazy 102-105 Lancer 69, 71 Maymount 145-147 North Saskatchewan River 145-147 Qu'Appelle Valley 103-104 Radville 69 Regina Lake Plain 67 Rocanville Valley 103-104 Steen 24 Thickwood Hills 145 Thunder Hill 13 Canada, Yukon Coastal Plain 4, 50 continental shelf 153 Herschel Basin 50 Herschel Island 14, 48, 50-52 Kay/King Point 50 Canadian Centre for Remote Sensing 31 Carboniferous 190, 207 Cenozoic 10, 101, 113, 137, 153-154, 167, 170, 175, 186, 211 ceramics 148 champagne 161 China 212 continental shelf 26 coteau, general 15, 168 Cretaceous, general 13, 28, 45, 53, 63, 70, 83-85, 87-89, 96, 102-103, 106-107, 135, 137-139, 153-154, 166-168, 171, 175-177, 179, 185-186, 188, 190 Ardley Coal Zone 143 Bearpaw Formation 69, 72, 106, 145 Campanian 89-90 Eastend Formation 69, 72 Grand Rapids Formation 101 Hell Creek Formation 59, 61 Judith River Formation 145-147 Lea Park Formation 145-147 Maastrichtian 64-65, 90, 107 Odanah Member 103, 105 Scollard Member, Paskapoo Formation 143
Aber and Ber
Ravenscrag Formation 69, 72 Riding Mountain Formation 13, 80, 103, 105 Senonian 112 Whitemud Formation 69, 72 crust (see lithosphere) Czech Republic 6, 168, 170, 182-184 Bohemia 183-184 Chuchelna 184 Fr3~dlant 183-184 Hlucfn 183 Hradek 183 Moravia 184 Nisou 183 Opava Upland 183 Ostrava basin 184 Zittau basin 183 Danske Atlas 3
Denmark 2, 6, 30, 32, 43, 44, 107, 153, 162, 168, 170, 171, 173 Als 32, A3r¢ 24, 32, 89, 171-172 Bornholm 156-157, 162 Copenhagen 171 Fakse Bugt/Bay 67, 93 Fan¢ Bugt/Bay 26, 27 Feggeklit, Mors 39-41 Flade Klit, Mors 63, 72-75 Fur 141, 147-151 Fyn 32, 89, 171-172 Great Belt/Storeba~lt 32, 67, 172 Hanklit, Mors 12, 72-77 Hesteg~rden quarry 151 Hjelm Bugt/Bay 67 Hundborg 45, 46 Hvideklint, Men vi, 83, 89-93 Hcje Men 63-67 Jylland (Jutland) 32, 43, 45, 171 Langeland 171-172 Limfjord 2, 24, 72-75, 77, 147-148, 151,171 LCnstrup Klint 2, 4, 31, 111, 171,206 ManhCj quarry, Fur 148-150 Men 32, 63-68, 107, 111,171,175 Men block 66, 68 MCns Klint 1, 3, 31, 63-68, 83, 171,174 Mors 148 Oresund 67 Ringkcbing-Fyn-Falster high 66 Ristinge Klint, Langeland 83, 88-89 RCgle Klint, Fyn 18-20 Rubjerg Knude 171 SalgerhCj 72-75
237
Index
Sams¢ 43 Sj~elland 32, 43, 171-172 SjCrring SO 45-46 Stevns 107 Systofte, Falster 114, 120-121 Viuf valley, Jylland 27 Devonian 97, 100, 135, 179, 181, 185-186 Onondaga Limestone 133 dragline 142, 144 Eemian 89-90, 93, 107, 171, 179 Elsterian glaciation 19, 75, 163, 170-171, 175, 181-186 Dubnice glaciation 183 Lvovfi glaciation 183 Oka glaciation 186 South Polish glaciation 185 Eocene 19, 28, 45, 63, 72, 172, 189 Fur Formation 12, 39, 40, 73-75, 147-149 Knudeklint Member 147-148, 151 Silstrup Member 147-148 Estonia 6, 102, 136-137, 1154, 156-157, 168, 170, 177-178, 180-181 Baltic Klint 109-110, 177, 181 Haanja Upland 177, 181 Karula 181 Otep~i~i Upland 177, 181 Pandivere Upland 99 Peipsi Lake/basin 99 Saadj~irve 83, 97-100 Sakala 181 Sinim~ied 109-110 Suur Munam~igi 181 Europe, general 6, 153-156, 158-160, 168-188 Alps 1,205, 208, 210 Baltic Sea/region v-vi, 6, 11, 30, 65-66, 68, 93, 96-98, 108, 153, 162, 170, 172-173, 175-179 Barents Sea/shelf 57, 107, 153-154 Caledonian Mountains 155 Carpathian Mountains 210 central Europe 6, 32, 168-187 Central European Plain 153 earthquakes 129 Fennoscandian Shield 109, 125, 127, 155, 157 Fennoscandian Border Zone 107, 156 Gulf of Finland 110 Jura Mountains 1 North Sea 26, 30, 153, 170, 181 Norwegian Sea/shelf 48, 55-57, 153 Scandinavia 1, 31, 127, 129, 156-158
excavator/excavation 17-18 Fennoscandian Ice Sheet 4, 130, 153, 157, 159 Finland 31, 157 fossils foraminifera 108, 204 microfossils 107, 109 France Mont Blanc 1 Alps 1 geographic information systems (GIS) 5, 6, 32 digital elevation model (DEM) 32, 47 Geological Survey of Denmark 148 geometric analysis 17-32 strike and dip, definition 19 structural analysis 12 trend and plunge, definition 19 geophysical methods 22-28, 133, 141., 1.43, 145, 1.71, 197198 acoustic survey 126 electrical resistance 23, 26 geo-electrical survey 73, 76 gravity survey 26, 134 ground penetrating radar (GPR) 24, 25 transient electromagnetic (TEM) 25-27, 73, 77 seismic survey/data 26, 27, 55, 106, 112, 118, 134, 190, 208 spontaneous potential (SP) 23 Germany 30, 44, 49, 98, 107, 112, 153, 162, 168, 170-176, 181-183 Bad Freienwalde 101-102, 171 Bad Muskau 182 Berlin 182 D~inisher Wohld 172 Dahlener Heide 181 Dtibener Heide 181 Eckemf6rde 172 Elbe River 173 Ftinfeichen 171 Grevesmtihlen 172 Heiligenhafen 172 Kiel 172-174 Mecklenburg-Vorpommern 171-172 Muskau Arc 181-185 Neukloster 172 Poel 174 Rauensche Berge 171
238 Rtigen 1, 107, 162, 174-176 Schleswig-Holstein 171-172 Usedom 175-176
Aber and Ber
glaciocataclasite 11 glaciodynamic 7-8, 44, 192-194 sequence/event, definition 44
glacial tectonics, first use 1 glacial theory 1, 5 glaciation annual moraine 63 buried valley 75, 77 clast pavement 168, 170 deposition 1, 5, 7, 44, 83, 158-159, 199 drumlin 15, 18, 19, 20, 31, 48, 49, 83, 89, 95, 97-100, 130-132, 153-154, 156-157, 160-161, 163-164, 166-167, 177, 179-181,200, 212 dump moraine 201 end moraine 13, 19, 61, 68, 109, 140, 153-156, 164, 171172 erosion 1, 5, 7, 44, 83, 132-133, 137, 156-159, 163, 177, 199 erratic 1, 41, 51, 126, 151, 153, 186 fjord 1, 42, 131-132, 196 flute/fluting 48, 177 forebulge 209 kame and kettle 41, 45, 51, 80, 82, 120, 164 lineation 158 moraine 1, 51, 72, 107, 164, 166 nunatak 72, 88 roches moutonn6es 130 Rogen moraine 157 s6rac 1 striation/groove/abrasion 1, 18, 20, 41, 101, 108, 115, 118-120, 126, 153, 162 glacier ice 6, 7, 130, 191-193, 195-196, 199, 203 block movement 49 calving 120 crevasse 121-124, 136 dome 44, 120 dynamic sole 190 frozen/cold bed 49, 115, 158-159, 161, 167-168, 190192, 195 ice divide/shed 153, 157, 179, 191-193 ice fan 117 ice lobe/tongue 30, 42, 43, 54, 56, 66, 79, 82, 117, 160168, 175, 179, 205 ice stream 43, 100, 162 melt water 72, 79, 120, 123, 130, 132, 150, 160, 162164, 191,195, 199 surge/surging 24, 35, 44, 56, 79, 98, 117, 120, 160, 164, 167-168, 175, 193 thawed/warm bed 49, 98, 115, 121, 158-159, 161, 164, 190-192, 195-196 velocity profile 190
glaciofluvial 77, 109, 121-122, 124, 143, 148, 150, 1.67, 177, 185, 187, 198, 200-201 esker 19, 20, 31, 45, 59-60, 80-82, 121-124, 130, 136, 155, 161-162, 164, 167, 186 sandur/outwash 19, 35, 41, 77-78, 84, 130, 163-164, 167, 187, 190, 200 spillway 19, 45, 53-55, 72, 78, 103, 162, 164, 167, 187 tunnel valley 19, 20, 53-55, 78-79, 122, 134, 136, 162, 164, 167 glaciokarst 7 glaciolacustrine 114-117, 126, 128, 130, 143, 173, 179 fan/delta 51, 164, 167, 195 Lake Madawaska 126 Lake Merrimack 155 proglacial lake 72, 98, 117, 160, 167, 172, 174-175, 181 varve 40, 106, 185 glaciomarine 117-118, 120, 167, 195-196 marine transgression 120, 204 glaciotectonic analog 203-213 comparison to mountains 12, 17, 203, 205-206, 212-213 convergent plate boundary 203, 210-213 delta mudlump 203-205 fluid migration 210-213 gravity sliding/gliding 208-209 gravity spreading 208-210 pushing from the rear 208-210 thin-skinned thrusting 203, 205-210 glaciotectonics, applied 141-151 coal mine/mining 141-144, 182-183 drift prospecting 141 highwall failure 142-144 highway construction 141,145-147 mine planning/operation 141, 145 mineral exploration 141 soil mapping 141 soil salinity/acidity 141 glaciotectonic, bibliography 5 glaciotectonic, case-history approach 15 glaciotectonic, definition 6-10 glaciotectonic deformation 1, 5, 7, 20, 44, 191-201,203 brittle 10-12, 37
Index
Coulomb principle 195 domainal 8, 41, 44, ductile/plastic 1, lO-12, 34, 37, 130, 164, 190 extra-domainal 41, 44, fundamental cause 191-193 geometric analysis 17-32 glacier bed 9 hydrofracturing 111 internal friction 192-194 liquefaction 111 penetrative 8, 9, 41, 44 piggyback thrusting 199-200, 204, 210 rotation 37, 38 scale model 198-201, 213 shear/shearing 146-147, 163, 172, 181, 190-197 static-kinematic conception 191, 210 superimposed 38-40, 182 thrusting 54, 57, 146-147, 163-164, 170, 184, 188, 193201,209-211 underplating 211 glaciotectonics, distribution 153-190 continent-scale 153-160 geomorphic setting 15 inner zone 153-154, 157-158 intermediate zone 153-156, 158 landscape symmetry 158-159 lobate model 162-167 outer zone 153-156, 158, 160, 170, 190, 212 regional 160-162 transition belt 156-157 zonal model 153, 158-159 glaciotectonic landforms 13-15 cave 11 composite ridge 14-15, 28, 35, 36, 59-83, 88, 138, 145148, 160, 162, 164-165, 168, 171, 174, 181-182, 184185, 197, 199, 203, 205-206, 208, 21 O, 212-213 crates de chevauchement 13 cupola hill (kuppelbakke) 14, 15, 83-98, 120, 138, 147, 157, 160, 163-165, 171, 177 depression/basin 35, 45-57, 83, 141-142, 164-165, 167169, 171-172, 176-177, 181, 186, 190 festoon pattern 136-137, 165, 200 glaciotectonic landscape 164-165 hill-hole pair 14, 15, 45-57, 147, 160-161,165, 168, 177, 188, 190 ice-shoved hill/ridge 12, 13, 56, 134, 156-157, 161-165, 167-168, 190, 193, 197-199, 201,210-213 irregular hills 14, 83 materials 14-15 megablock plain 105 moraine de chevauchement 13 moraine de pouss6e 13 pseudo-moraine 13
239 push-moraine 2, 13, 24, 25, 60-61, 63, 138, 167, 171172, 176, 181, 183-184, 186, 199-202 roches moutonn6e 1 Stauchrticken 13 Stauchendmor~inen 13 Stauchmor~inen 13 stuwmorenen 13 stuwwallen 13 transverse ridges 59 trellis drainage 51 glaciotectonic structures 10-11 allochthonous 7, 59, 207 anticline/syncline 40, 52, 65, 101, 115, 118-120, 133, 150-151, 172, 174, 177, 182, 203-204, 207-208 antithetic fault 35, 41 apophysis 82, 113 arrowhead fold 38, 39 augen 18, 34 autochthonous 7 avalanche/debris/mass flow 126, 128 basement fault 125-140, 153, 167, 170, 176-177, 186, 208 boudin/boudinage 8, 10, 18 box fold 40 breccia/brecciation 7, 10, 39, 44, 53, 78, 88, 90-91, 95, 101, 103, 107, 109-110, 145, 147-148 cataclasis 44 chalk-till m61ange vi, 11 cleavage 7, 8, 10, 18 competence/incompetence 69, 72 concealed 32 conjugate faults 34, 40, 92, 151 d6collement 7, 8, 12, 27, 36, 40, 41, 57, 72-73, 78-79, 89, 102-103, 146, 150-151, 162, 190, 193, 203, 205-213 diapir 10, 11,109-110, 111-113, 115-120, 123, 130, 139, 160, 171,177, 179-181,186-187, 197, 203-205, 210-211 dike/vein 18, 112-113, 182 drag fold 89, 120-121, 157, 199 duplex 139-140, 208 endiamict 8-10 exfoliation 144 exodiamict 8-10 extension fracture 33, 151 fault/faulting 7, 10-12, 18, 19, 39, 40, 49, 73, 81-82, 87, 90-91, 95, 100-101,119, 122-124, 134-138, 141,143-145, 147-148, 151, 181, 187, 189-190, 204, 208, 210 fault-line scarp 127 fissure/fissuring 11, 18, 112, 146, 150-151, 181, 187, 197 floe/raft 12-13, 83, 88-91, 93, 96, 101-109, 111, 157, 160, 171, 176, 177, 179, 181-183, 186-187, 200-201 fold/folding 7, 10, 12, 19, 28, 39, 40, 42-44, 52, 69, 73, 78, 85, 92, 96, 100-101,106-107, 109, 113, 115-116, 118125, 133, 141,143-145, 147-148, 150-151,172, 174, 177, 179-183, 186-187, 190, 198, 203-206, 208
240 fold axis 18, 21, 35, 101, 116, 118, 120-122, 125, 148, 150, 172 foliation 8, 10, 18, 89, 96, 113, 118 fracture/fracturing 11, 39, 95, 141, 143-144, 150-151, 181, 198-199 gouge 145-147 graben 147 groove 113 horst 129, 139-140, 162-163, 175, 184 identification of 7 intrusion 73, 111, 113, 116, 118, 120-121, 123, 211 joint 11, 40, 96, 141,144 kink/kinking 82, 198-199 line/linear 18, 21 lineation/lineament 8, 10, 45, 48 megablock 13-14, 93, 101-109, 160, 167, 177, 179, 181, 190, 193, 196-197, 201 m61ange 8, 10, 11, 90, 92, 106, 210-211 mylonite 7, 8, 10-11,103, 190 normal fault 35, 36, 43, 78, 82, 93, 151, 194, 197, 205, 208 omission of strata 23, 141 plane/planar 18, 21 plug 116 primary 7, 18, 41 ramp 197 refolded fold vi, 38, 39 release fracture 33 repetition of strata 23, 52, 141 reverse fault/faulting 127, 148 scale 13, 85, 87-89, 96, 174, 177 secondary 8, 18 seismic zone 125-127, 130, 153, 167 shear plane/zone/band 11, 12, 33, 52, 73, 78, 92-93, 96, 99-101, 106-107, 126, 143-145, 150-151, 160, 190 sill 111,116 slickensides 7, 10, 52, 73, 78, 95-96, 103, 106, 143, 145, 147, 187 slide/slump 144, 146-147, 197-199 stock/plug 111 tear fault 45, 48, 49, 207, 212 thrust block/mass/sheet 12, 35, 49, 65, 69, 73, 75-76, 78, 107, 109, 112, 141, 143-145, 161, 172, 190, 197, 199, 209, 212 thrust/thrusting 10, 24, 28, 35, 37, 40-44, 52, 69, 71-72, 75, 80, 87-89, 93, 100-101,111, 113, 115, 117, 119-120, 122-123, 133-134, 138-140, 143, 150-151,157, 160-162, 174, 181-183, 185-188, 190, 193-201,204-208, 211-213 unconformity 18, 44 underthrusting 199, 211-212 vergence 35, 123, 208 wedge 111-113, 120-121,150-151 xenolith 113 glaciotectonite 8, 10, 39
Aber and Ber
Gondwana 212 Greenland 1,130 ground water 158, 170, 195, 212 pressure 11, 54, 88-89, 100, 111, 116, 141, 144, 162, 195, 200, 211,213 aquifer 48, 60, 88, 98, 160-162, 167-168, 175, 195, 213 hydraulic conductivity 98, 145 hydrodynamic blowout 55, 111,211 piezometer 143-144 protection 141 saline water 190 Spiritwood aquifer 54, 55 spring/fountain 129, 195, 212-213 halokinesis/halotectonic 130, 133, 140 Holocene 126, 146, 185, 196 Little Ice Age 35, 60, 62 Neoglaciation 2, 167 Holsteinian/Mazovian interglaciation 19, 137-138, 179, 188 iceberg 1, 6 Iceland, general 210 Eyjabakkaj6kull 35, 36 K6tluj6kull 200-202 M3~rdalsj6kull 61-62, 200 M3~rdalssandur 200 S16ttj6kull 61-63 Vatnaj6kull 35 Illinoian glaciation 125 Montauk Till 84-85, 87 Independence glaciation 114-117, 125 Dakota lobe 117 Independence Formation 113-115, 117 Minnesota lobe 117 India 211-212 Ganges Plain 212 Main Boundary Thrust 212 Main Central Thrust 211 Indian Ocean Bay of Bengal 212 Arabian Sea 212 INQUA Commission on glaciation 5 Congress in Canada 5 Congress in Germany 6
241
Index
Congress in United States 6 Work group on GAGE 6 Work group on glacial tectonics 5 Ireland 153
pressure shadow 37 sigmoidal grain 37, 38 Miocene 26, 28, 85, 87, 112, 139-140, 163, 177, 182-185, 189, 212 Poznafi Clay 185
Italy, Alps 1 Mississippian 207 Jurassic 101, 107, 112, 135, 137, 140, 156-157, 179, 187 Kara Sea ice sheet 154, 188, 190 karst 109
Mohr stress equation/circle 33, 34, 40, 194 failure envelope 34, 35 Neogene 28, 109, 112, 135, 137-138, 140, 154, 156, 162, 175, 177, 1.81, 185-186
key section 18, 31, 42 neotectonic 38, 151 kinematic analysis 33-44, 199 kineto-stratigraphy 40-44 principle of 2, 41 landslide 1, 145-147, 203 Latvia 137, 154, 157, 168, 170, 177-179 Gulf of Riga 179 Latgale 179 Sensala 179 Strante 179 Ulmala 179 Vidzeme Upland 177, 179 Laurentide Ice Sheet 4, 82, 88, 153, 157-159, 164, 170 Foxe sector 125 Keewatin sector 159 Quebec sector 159 southern margin 32, 125, 168 lithosphere (crust) 210-212 depression/rebound 7, 52, 120, 125, 127, 130, 135, 137, 153, 158, 212 Lithuania 101, 133-136, 154, 157, 168, 170, 177-180 Pakalnigkiai 180 Klaipeda 180 ~emajtija Upland 137, 177 megageomorphology 31 Mesozoic 10, 101,113, 125, 137-138, 153-154, 167, 170, 186 micromorphology 36-38 galaxy structure 37, 38 microfabric 36 microstructure 36 plasma 36, 37
Netherlands 2, 153, 162-164, 168, 181-182 Gelderse Vallei 75, 77-78 Kwintelooijen 75, 77-78 Rhine/Meuse River 77, 79 Utrecht Ridge 63, 75, 77-79, 162, 181 Veluwe push moraine 24, 181, 197-198 Nidanian glaciation 140 North America, general 4-6, 31, 153-160, 167-168, 170 Appalachian Mountains 125, 133, 155, 167, 205, 207210, 212-213 Appalachian Plateau 131-133 Atlantic Coastal Plain 83-84, 88, 153, 167 Atlantic coastal zone 125 Atlantic continental shelf 84, 153 Beaufort Sea 50 Canadian Shield 125, 155, 170 Central Lowland 156 Great Lakes 156 Great Plains 4, 6, 15, 42, 49, 114, 153, 162, 167-168, 170 Missouri Coteau 29, 59, 67-69, 168, 185 Ontario Lowland/basin 132-133 Ouachita Mountains 212-213 Prairie Coteau 117, 168, 170 Saint Lawrence zone/lowland 125-126 Susquehanna basin 132 Norway 6, 127, 129-130, 155, 157, 196 fjords 1, 117, 120 FuglCybanken 55-57 glaciation 1 Herdla Moraines 114, 117-118 Oslo 151 Tra~nabanken 55, 56 Voss 112-113
242
Aber and Ber
Oligocene 19, 26, 28, 102, 138, 177, 185, 189
Pliocene 87, 139-140, 185
Ordovician 11, 97, 99-100, 109-110, 135, 138, 186, 207
Poland 1, 2, 44, 45, 125, 133-140, 154, 1.62, 164, 168, 170171, 173-177, 181-185 Dalk6w 184 Dylewski Hill 138 Elbhtg Upland 18, 83, 93-98, 135, 137-138, 175, 177 Elk Upland 165 Gdansk Bay/embayment 93, 97-98, 175 G6rowo Hills 135, 137-138 Gr~bocice 185 Itawecki Hump 138 Kronowo esker 114, 121-124 Lidzbark, Warmia 137-138 Lubawa Elevation 138 Lukow 101 Mach6w 199 Mazury Lakeland 137-138 Mu~ak6w Arc 185 Odolan6w basin 139-140 Ortowo, Warmia 26, 28, 112, 138 Ostrzesz6w Hills 139-140, 163, 184 Pomeranian Upland 162 Pomorze Bay 176 Silesian Rampart 140, 175, 184-185 Suwatki Upland/Lakeland 133-138, 165, 175-177 Syc6w 185 Szeskie Hills 135, 137 Szczytno-Okr~tgte Lake 135-136 Trzebnica Hills 139-140, 162-163, 184, 208 Warmia 137-138 Wielkopolsa Lowland 140 Wifisko 184-185 Wista (Vistula) delta/river/valley 94, 96, 135, 156 Wi~ajny elevation 134-135, 137 Wolin Island 175-177 Zielona G6ra Rampart 175, 184-185 ~uromin 137-138
Pakistan 212 Paleocene 28, 45, 143, 189 Danian 107-108 Kerteminde Marl 171 Paleogene, general 23-24, 27, 28, 39, 112, 135, 137-138, 140, 147, 153-154, 156, 171-172, 175, 177, 181,185-187 Holmehus Formation 39, 40, 75 Olst Formation 39 Paleolithic artifacts 77 Paleozoic 11, 90, 102, 107, 109, 114, 125-126, 132, 135, 156-157, 166, 207, 212 Caledonian 135 Pennsylvanian, general 12, 101,207 Bethany Falls Limestone 101 Dennis Limestone 101-102 Elmont Limestone Member 101 Galesburg Shale 101-102 Stark Shale Member 102 Tarkio Limestone Member 101 Willard Shale 101 Winterset Limestone Member 102 Zeandale Formation 101 permafrost 4, 11, 15, 24, 48, 52, 53, 72, 79, 89, 96, 104, 113, 158, 160-162, 164, 167, 1.70, 175, 190, 193, 195-196, 21.3 continuous 50 discontinuous 107 ground/massive ice 188, 190 ice wedge 113, 150 patterned ground 61, 113, 151
Poozerie glaciation 176 Permian 135, 156, 190 Zechstein salt beds 140
power shovel 142-143
petrographic microscope 37
pre-Illinoian glaciation 32, 125
Pleistocene 42, 50, 63, 75, 83-85, 87, 94, 102, 112, 114, 135-138, 140, 153, 158, 160, 162, 164, 167-168, 172, 175177, 179, 181, 185, 187, 197-198, 201,212 Aquinnah conglomerate 85 Bow Valley gravel 105-106 Cromer Tills 111-112 Floral Formation 145 Gozdnica Series 140 Kedichem Formation 77-79 Urk Formation 77-79
pre-Quaternary strata 14, 23, 32, 44, 50, 59, 170, 179, 18 l, 184, 186 Proterozoic (precambrian) 11,125, 127, 134-135, 137, 157, 168, 207-209 Sioux Quartzite 117 Quaternary 2, 5, 14, 23, 28, 40, 44, 48, 56, 57, 59, 61, 75, 89, 96, 101,109, 111, 134-135, 137-139, 141, 143, 153-154, 156-158, 167, 170-171,177, 181,184-188, 190-191,199
243
Index
radiocarbon (C14) date 61, 71, 82, 85, 97, 108-109, 120, 146, 176, 204 remote sensing 26-31 aerial photography 27-28, 31, 48, 49, 51, 52, 54, 66, 7071, 74, 80-81, 91, 143 blimp aerial photography 31, 47 digital orthophotograph 30, 47 Ikonos 31 kite aerial photography 31, 75, 88, 99 Landsat imagery 29, 30, 53, 54, 98, 140, 158, 203 manned-space photography 30, 31, 170, 172, 174 MODIS imagery 173, 178 multi-concept 32 photogrammetry 31 photolineament analysis 137 Radarsat 31 Skylab missions 30 small-format aerial photography 31 rock/sediment types algal reef 204-205 amphibolite 210 andesite 210 anorthosite 134-135 argillite 113, 125, 187 ash 147-148, 151 bentonite 40, 69, 106, 143-147 blueschist 210 chalk 10, 11, 63-66, 68, 73, 89-93, 107-109, 111-112, 141, 154, 171, 175-177, 185-186 chert/flint 11, 64, 69, 73, 107-108, 116 clay/silt 10, 11, 19, 23, 24, 35, 50, 52, 53, 56, 65, 73, 75, 77, 85, 88, 93, 96, 98, 100-102, 104-109, 111-119, 121, 140-141,145-147, 154, 173, 175-177, 181-182, 185-190, 199, 203-205 claystone 70, 96, 103, 140, 148, 150-151, 195, coal 10, 12, 106, 140-144, 168, 182-183, 190, 212 concretion 102, 148, 177 conglomerate 12 crystalline 5, 7, 11, 69, 84, 90, 107, 134-137, 155-158, 170, 191 diatomite 12, 39, 40, 73, 147-148, 189 dolostone/dolomite 11, 69, 97, 99, 109, 179, 185-186 glauconite/greensand 85, 177, 186 glimmersand 102 gneiss 11 granite 1, 11 greenschist 210 hydrocarbon 212 igneous/metamorphic 12, 210-212 kaolinite 69, 85, 88 lignite 10, 69-70, 85, 88, 195, limestone 11, 34, 97, 99, 101-102, 107, 109-110, 114, 125, 186
loess 113, 163 magma/lava 111 marble 35 marl 96, 100, 176-177, 186 migmatite 210 moler (mo-clay) 73, 147-151 montmorillonite 73 mudstone 10, 68, 72, 143-144 oil shale 110 opal 147 peat/gyttja 22, 23, 50, 71, 77, 188 porphyry 141 quartzite 11, 69 quick-clay 111, salt 111, 125-126, 130, 132-133, 210 sand/gravel 10, 11, 19, 25, 35, 50, 61, 65, 73, 75, 77, 82, 85, 92, 96, 98, 100, 102, 104-108, 111-123, 140-141,145146, 148, 150-151,157, 161,173, 176-177, 180-182, 185186, 188-189, 198-200 sandstone 10, 11, 68, 70, 72, 100, 106, 109, 113, 125, 141, 143-144, 147, 156-157, 181 saprolite 168 schist 11,141 serpentine 210 shale 11, 12, 24, 53, 68, 80, 101,103, 106, 114, 141,143, 147, 166, 168, 190, 195, 210 siltstone 96, 100, 103, 140, 147, 181 slate 125-127, 156 smectite 147 soil/turf/loam 71, 77 sulfur 199 tephra 35, 39, 40, 73, 147 till (see below) volcanic 200, 203 Russia 6, 31, 153, 157, 171 Arctic region 154, 156, 187-190 Atlym 189-190 Belugalakh Bay 189-190 Hongurei-Markhida 187-188 Karelia 127, 129-130, Kaliningrad 133-134, 168, 170 Kola Peninsula 157, 188 Lake Baikal 212 Lake Sevastian 189-190 Novaya Zemlya 156, 188 Ob River 189-190 Pechora River 187-188 Pechora Sea continental shelf 154, 187-188 Siberia 153-154, 167, 187, 189-190 Taymyr Peninsula 187 Tiksi 189-190 Ural Mountains 154, 190 Vastiansky Kon' 187
244 Saalian glaciation/till, general 19, 39, 75, 78-79, 89, 93, 162-163, 170-173, 175, 181-186 Dnieper glaciation 113, 185-186 Drente Formation 77-78 Gelderse Vallei lobe 78 Grodno till 176 Hennstedt glaciation 172 Jitrava glaciation 183 Odranian glaciation 138, 140, 185 Sozh glaciation 185 Ugandi 181 Wartanian glaciation 140, 181, 185 Sangamon 126 Pelukian 51 Silurian 97, 99-100, 135 Salina Group 132-133 Slovakia, Tatra Mountains 6 soft-sediment deformation 111, 117-118, 203 South America, Andes Mountains 210 Soviet Union 4, 6
Aber and Ber
normal stress 33, 192, 194 Poisson's ratio 192 shear stress/strength 33, 34, 37, 56, 57, 78, 116, 145, 147, 190-197, 203 tension/extension 33, 34, 150-151,205 survey methods 17, 18 altimeter 17 compass 17, 18, global positioning system (GPS) 17 Svalbard 2, 57, 196 Scott Turnerbreen 24-25 Spitsbergen 2, 42, 120 Sweden, general 107, 109, 127, 129-130, 155-157 Alnarp Graben 107-108 Anders6n 157 Gotland 156 Kvarnby/,~ngdala 102, 107-109 Lake Kamasjaure 129 P~irve fault 127, 129-130 Skhne 1, 31, 107-108, 176 Tornquist Line Oland 156 Oresund 108 Ostersund 157
squeeze box 198-199 stereographic projection 17, 20-22, 93, 96, 115, 119, 121, 150 Pi diagram 22 Schmidt stereonet 20, 21 Wulff stereonet 20 strain 9, 33, 35, 37 dilation 33, distortion 33 stratotype 114-115, 117 stress/pressure 33, 34, 117, 122, 137, 187, 191-197, 205, 209 angle of failure 33, 34, 40 angle of friction 33-35, 194, 200-201 cohesion/cohesive strength 33-35, 192-195 compression 34, 160, 181,190, 197, 205, 209 differential stress/loading 33, 78, 125-126, 130, 133, 136, 140, 160, 186, 192-193, 204-205, 209 ellipsoid 34 glaciostatic 7, 170, 191-192, 194, 197 hydrostatic 53, 192, 195-196, 201,203, 205 intergranular 192, 195, 205, lateral 78-79, 123, 160, 191-193, 195, 201,205, 209 lithostatic/confining 33, 34, 192, 195
Switzerland 113 Alps 1, 63, Turtmannglacier 199-200 Unteraar Glacier 1 tectomict 9 tectonics accretionary wedge 210-212 collision zone 210 hydraulic arc 210-211 mud volcano 210-211 magmatic arc 210-211 plate tectonics 63, 130, 203, 210-213 ramp/flat 206-208 regional tectonics 125 rift zone 208, 212 subduction zone 210-211 transform fault 212 transverse zone 207-208 Tertiary 26, 83-85, 96, 107, 164, 167, 188, 190, 199 Bow Valley gravel 105-106 Cannonball Formation 59, 61, 70 Rupelian 101-102 Tethys Sea 211-212
245
Index
thin section 37 till, general. 12, 19, 23, 24, 40, 41, 44, 51, 59, 61, 63, 65, 73, 75, 80, 82-84, 90-96, 99-109, 112, 114-118, 120-121,129, 137, 141, 143, 145-146, 151, 153-154, 157, 163, 171173, 176-177, 179-181, 183, 187-188, 190, 1.96, 200 fabric 7, 18, 20, 37, 41, 42, 107-109, 115, 117, 121,162 flow till 121-124 deformation 8-10, 42, 164, 168 diamicton 37, 97, 99, 113, 121, 123-124, 126, 185 comminution 8 lodgement 10, 63 Triassic 135, 137, 139-140, 162-163, 179, 185 Ukraine 6, 112-113, 154, 168, 170, 181, 186-187 Kiev 187 Shevchenko valley 186-187 United Kingdom England 1, 11, 153 Isle of Man l, 3 Norfolk 1, 4, 111 West Runton 11 Weybourne 11
United States, Massachusetts 88 Aquinnah, Martha's Vineyard 16, 83-88, 162 Cape Cod 84 Gay Head, Martha's Vineyard 83 Menemsha Bight 88 Nantucket Island 84 Vineyard Sound 87 United States, Michigan 22-23 United States, Minnesota l, 168-169 Big Rice Lake 169 United States, Missouri 117 New Madrid zone 125 United States, Montana 168 United States, Nebraska 117 Nemaha zone 125 United States, New England 1, 32, 153, 167 United States, New Hampshire 156 United States, New Jersey 125
United States, Alabama 207 Anniston 207-209 United States, Alaska 153, 167 Baldwin Peninsula 167 Bering Sea coastal zone 153, 167 Brooks Range 167 Kotzebue Sound 167 Malaspina Glacier 88 United States, general 32, 125, 160, 167-168, 208 United States Geological Survey, EROS Data Center 29, 30 United States, Illinois 12 New Madrid zone 125 United States, Iowa 10 l- 102, 117 United States, Kansas 6 Atchison 114-117 Nemaha zone 125 Topeka 101 Wathena 116-117 United States, Louisiana Mississippi Delta 203-205 South Pass distributary 203-205
United States, New York 88, 125, 133, 156 Lake Ontario 1, 131-132, 156 Finger Lakes 125, 130-133 Keuka Lake 131-132 Long Island 84, 167 Niagara escarpment 131-133 Seneca Lake 131, 133 United States, North Dakota 4, 15, 32, 45, 161, 168 Anamoose 161 Antelope Hills 45, 46 Crow Hill 53, 54 Devils Heart Butte 53, 54 Devils Lake 48, 53-55, 167 Devils Lake Mountain 45, 47, 54 East Devils Lake 53 Geologic map of 4, 19 Knife River 60 Pelican Lake 60 Prophets Mountains 59-61 Sullys Hill 53-55 United States, Ohio Anna zone 125 Cleveland 125 Lake Erie 125-126 United States, Pennsylvania 133
246 United States, Rhode Island 88 Block Island 84 United States, South Dakota 117, 166, 168 Watertown 170 United States, Wisconsin 4, 166 Vineyard orogeny 85 volcanism 1,209-211 Weichselian/Vistulian glaciation 13, 19, 20, 30, 42-44, 63, 67, 75, 98, 122, 134-138, 153, 156, 170-181, 188, 190, 196 AllerCd 120 Ba~lthav/Storeb~elt advance/lobe 43, 67-68, 89, 93, 121, 162, 171 Baltic ice lobe 162 Brandenburg phase 171 East-Jylland phase 89, 121, 162 Frankfurt phase 171 ice-dome model 44 ice-lobe model 43-44, 162 Jydelejet phase 67 Kurentsovo line 188, 190 Leszno phase 135 Main Weichselian advance 43, 67, 75, 93, 112, 124, 171 Nemunas 177 Norwegian advance 43, 75, 151 North Polish, Baltic phase 94 North Polish, Pomeranian phase 97 Old Baltic advance 43, 67 Oresund advance/lobe 67-68 Otep~i~i phase 99 Pandivere phase 99, 109 Pomeranian phase 135, 171-172, 175 Young (Low) Baltic advance/lobe 43, 67, 89, 93, 108, 121, 171-172 Younger Dryas glaciation 117-118, 120 Wigry phase 135 Weichselian/Vistulian stratigraphy/till 2, 39, 90, 108, 157 B~elthav Till 67 East Jylland Till 67 Elbl~tg Clay 95-96 Hanklit glaciogenetic Group 73, 75
Aber and Ber
Krastudy interstade 95 Kvarnby Till 107, 109 Mid Danish Till 67, 93 North Sja~lland Till 67, 93 Ristinge Klint Till 67, 93 S. Sallerup Till 107, 109 Sunnan~ Till 108-109 Tromscflaket diamicton 57 Trz~sacz till 176 well drilling/logging 13, 17, 22-24, 26, 27, 32, 95, 103, 105-109, 113, 122, 133, 137, 139, 141,143, 145-146, 156, 171, 181, 188-190, 204, 208 Wisconsin glaciation 13, 57, 84-85, 104, 117, 125-126, 142, 153, 167-168, 170, 204 Ardill end moraine 72 Assiniboine sublobe 82 Avonlea tongue 70-72 Battleford ice lobe 145-146 Battleford Till 145 Bemis moraine 170 Buckland glaciation 50-52 Cape Cod Bay lobe 84 Cold Lake glaciation 49 Cooperstown margin 54 Des Moines lobe 170 Galilee tongue 70-72 Green Bay lobe 164, 166 Heimdal margin 54 James lobe 164, 166, 168, 170 Lostwood glaciation 71, 82 Marchand phase 82 Martin ice margin 161 Menemsha tongue 88 Narragansett-Buzzards Bay lobe 84, 88 •North Viking margin 54 Outing moraine 168-169 Pekin margin 54 Primrose lobe 49 Spring Valley tongue 70-72 St. Croix moraine 168 Stewart Lake moraine 168 Valley Heads Moraine 132-133 Vineyard moraine 84, 88 Weyburn lobe 71
