International Geography 1972: Volumes 1 and 2 9781442615120

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International Geography 1972 La geographie internationale

Geomorphology Geomorphologie Climatology, Hydrology, G l a c i o ' 9 Y Climatologie, Hydrologie, G l a c i o ' ° 9 0

Biogeography and Pedology Biogeographie et Pedologie Regional Geography Geographie regionale Historical Geography Geographie historique Cultural Geography Geographie culturelle Political Geography Geographie politique Economic Geography Geographie economique Quality of the Environment Qualite du milieu

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International Geography

1972

La géographie internationale

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International Geography 1972 La géographie internationale

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Papers submitted to the 22nd International Geographical Congress, Canada Communications présentées au 22e Congrès international de géographie, Canada Edited by / sous la direction de W. Peter Adams and Frederick M. Helleiner Trent University

Published for the 22nd International Geographical Congress Publié à l'occasion du 22e Congrès international de géographie Montréal 1972 University of Toronto Press

University of Toronto Press 1972 Toronto and Buffalo Printed in Canada ISBN 0-8020-3298-2 (Cloth) Microfiche ISBN 0-8020-0271-4

Preface Préface This two-volume collection of papers contains most of the written submissions to the various events of the Centennial International Geographical Congress held in Canada in 1972. This Congress is the quadrennial conference of the International Geographical Union, the major world-wide association of geographers. The first Congress was held in Antwerp, Belgium in 1871, and the most recent ones in Stockholm (1960), London (1964), and New Delhi (1968). The volumes have been produced to provide a broad basis for the discussion-format proposed for this particular Congress. In order to strengthen this basis, authors were requested to submit 'short papers,' of no more than 1200 words including references, rather than the usual abstracts of their work. These papers provide delegates with a very substantial briefing on the principal topics of the meetings. In addition to having an important function in connection with the Congress, these volumes, unlike the usual collections of conference abstracts, provide a remarkable overview of work in progress in geography around the world. The work of geographers in over 60 countries is represented, and the topics encompass the entire range of the spectrum of modern geography. The articles published here include those submitted to the central programme of papers in Montreal and substantial groups of papers submitted to special Congress events, Commission meetings and Symposia, on specialized subjects such as 'Water Resources' and 'Geography in Education.' We feel that the inclusion of references with most of the papers greatly enhances the value of this collection as an 'overview' of geography. Preparing such volumes for use at a large international conference is an exceedingly complex operation. The editorial function must be performed within a wide range of constraints not normally encountered in collecting and publishing the works of others. Users of these volumes should therefore recognize the limitations and biases of the sample of geographical research represented by these papers. Even without the emphasis on discussion

Cet ouvrage en deux volumes renferme la plupart des communications écrites soumises dans le cadre des différentes activités du Congrès international de géographie qui se tient au Canada en 1972. Ce congrès est la conférence quadrienniale de l'Union géographique internationale, la principale association de géographes du monde. Le premier congrès s'est tenu à Anvers, Belgique, en 1871, et les plus récents congrès à Stockholm (1960), à Londres (1964) et à New Delhi (1968). On a préparé ces volumes dans le but de fournir un document de travail pour les réunions des diverses sections, commissions et symposiums du Congrès. Les auteurs avaient été invités à soumettre de cours textes, n'excédant pas 1200 mots, y compris les références, plutôt que les classiques résumés, afin que les délégués aient en main un important dossier sur les principaux sujets de discussion à leurs réunions. En plus d'être un outil de travail important lors du congrès, et contrairement aux compilations habituelles de résumés des communications, cet ouvrage présente une vue d'ensemble exceptionnelle de la recherche géographique actuelle à travers le monde. On trouvera dans ces volumes les textes soumis en vue des réunions des sections, ainsi que plusieurs textes qui seront présentés ou discutés lors des activités spéciales du congrès, de réunions des Commissions, et des symposiums; ces travaux traitent de sujets spécialisés tels que 'Les ressources en eau' et 'La géographie dans l'éducation.' Nous croyons que cet ouvrage offre une vue d'ensemble d'autant plus importante que, dans la plupart des cas, une liste de références est incluse à la suite du texte. Il est extrêmement compliqué de préparer un ouvrage en vue d'un congrès d'une telle envergure. On doit en effet tenir compte de multiples contraintes qui n'existeraient pas en temps normal dans la compilation et la publication de travaux collectifs. Les lecteurs de ces volumes doivent être conscients des limites de l'échantillon de la recherche géographique que représentent ces textes. Au cours d'un tel congrès, où l'accent sera mis sur la discussion, il est impossible que

vi / Preface which is a special feature of this Congress, it would be impossible to have all worthwhile submissions presented orally. Instead, organizers of the various Congress events were encouraged to use this volume as a means of distributing acceptable papers submitted to them while building their actual Congress programmes around a selection of those papers published. Thus many excellent papers published here will not be presented, in the formal sense, at the Congress. We hope that all authors who submitted papers to this Congress will be sympathetically aware of the difficulties faced by ourselves and by organizers of the programmes of papers involved. The event-organizers are dispersed widely over this rather large country and the sheer mechanics of correspondence between them and us, in addition to the complexities of communicating, in two languages, with over seven hundred persons scattered around the world, have inevitably resulted in mistakes and omissions. The absolute imperative of publication before the Congress has not simplified the task. Under other circumstances we would have consulted more closely with more authors before exercising our editorial prerogatives. Given the problems presented by the scale and complexity of the Congress, we could not attempt to produce here a collection of papers which is homogeneous in style and quality but, rather, we have assembled a readable collection which will provide a useful working basis for the Congress and, even more, one which will form a valuable reference work after the Congress. In preparing the volumes, our more important editorial objectives were : to publish 'short papers,' of about 1200 words including references (the majority of papers do not conform to this ideal but our editing has been undertaken with it clearly in mind) ; to publish in either French or English, as submitted; to include the name, initials, professional affiliation, and country of all contributors; to eliminate all except the most essential subheadings (this includes 'Acknowledgments' and 'References,' which appear, in this order, immediately following the text of each paper, as discussed below) ; to ensure that the text is 'readable' (no effort has been made to develop uniformity of style). All papers were read for content and tech-

les auteurs puissent tous présenter oralement leur travail, aussi important qu'il soit. On a donc encouragé les organisateurs des différentes activités du Congrès à utiliser cet ouvrage pour diffuser les communications qui leur ont été soumises et pour élaborer leur programme autour d'un choix de textes publiés. Ainsi, plusieurs textes d'excellente qualité qui font partie de cet ouvrage ne seront pas présentés de façon conventionnelle au congrès. Nous espérons que tous les auteurs qui ont soumis des communications écrites à ce congrès se rendront compte des difficultés auxquelles les organisateurs et les soussignés ont dû faire face. La complexité qu'impliqué la correspondance avec les organisateurs disséminés dans ce grand pays, ainsi que la nécessité d'entrer en communication dans deux langues avec plus de sept cent personnes à travers le monde, ont occasionné des omissions et des erreurs inévitables. La tâche n'a pas été simplifiée par le fait qu'il était absolument nécessaire de publier cet ouvrage avant le Congrès. En d'autres circonstances, nous serions entrés en contact avec un plus grand nombre d'auteurs avant de nous prévaloir de notre rôle de rédacteurs. Compte tenu des problèmes occasionnés par un congrès de cette envergure et de cette complexité, il nous était impossible de publier un recueil de travaux dont le style et la qualité soient homogènes; nous avons plutôt compilé un ouvrage qui se lit bien, du moins nous l'espérons, et qui servira de document de travail utile pour le Congrès; par dessus tout, il constituera une importante source de références après le Congrès. Les principales normes dont nous avons tenu compte dans l'édition de ces volumes sont les suivantes: la publication de courts textes d'environ 1200 mots, y compris les références (la plupart des études ne se conforment pas à cet idéal mais c'est à partir de cet objectif que nous avons dirigé notre travail d'édition) ; la publication en français ou en anglais, selon la langue utilisée; l'inclusion du nom, des initiales, de l'appartenance professionnelle et du pays de tous les collaborateurs; l'élimination de tout sous-titre à l'exception de ceux qui sont absolument nécessaires (y compris 'Remerciements' et 'Références' qu'on trouve dans cet ordre à la suite de chaque communication); le désir

Préface I vii nical language by the event organizers named at the beginning of each section. These persons, their selected referees, and Professor J. Warkentin, Chairman of the Programme Committee, were responsible for screening the papers. Further details of the organization of the volume, numbering of articles, pagination, etc. are given below, and additional notes are included at the beginning of each major division of the books. We gratefully acknowledge the assistance of: The event organizers named at the beginning of each section; Jill Adams; Douglas R. Barr; J.B. Bird, Chairman, Congress Organizing Committee, for consistently firm encouragement; R.I.K. Davidson, Lorraine Ourom, and other members of the staff of the University of Toronto Press; Neil Durford; J.K. Fraser, Congress Executive Secretary, and his staff, especially Janet McDonald and Louise Moore for devotion to duty; Nicole Gaudet; Lois Helleiner; Cathy Hewton; D.P. Kerr; Louise Lettelier; Department of Geography, Laurentian University, Canada; Judy Pack, for almost two years of careful work; Marion Robertson, for assistance with translation from the Russian; Louis Trotier; J. Warkentin. W.P.A., F.M.H. Trent University, Canada June 1972 Juin

de publier des textes 'lisibles' (on n'a pas essayé d'établir un style uniforme). Toutes les communications ont été examinées quant au contenu et au vocabulaire technique par les responsables dont le nom apparaît au début de chaque section de ces volumes. Ces personnes et les arbitres choisis par eux, de même que le professeur J. Warkentin, président du Comité du programme, se sont chargés d'examiner et de sélectionner les textes soumis. On trouvera d'autres renseignements concernant la disposition de l'ouvrage, le numérotage des articles, la pagination, etc., sur une page supplémentaire au début de chaque chapitre. Nous remercions sincèrement: les responsables dont le nom paraît au début de chaque section; Jill Adams; Douglas R. Barr; J. Brian Bird (président du Comité d'organisation, pour son encouragement sincère et constant) ; R.I.K. Davidson, Lorraine Ourom et autre personnel de l'University of Toronto Press; Neil Dunford; J. Keith Fraser (secrétaire exécutif du Congrès, et son personnel, tout spécialement Janet McDonald et Louise Moore pour leur dévouement; Nicole Gaudet; Lois Helleiner; Cathy Hewton; Donald Kerr; Louise Letellier; Département de géographie, Université Laurentienne, Canada; Judy Pack, pour presque deux ans de travail consciencieux; Marion Robertson, pour son aide dans la traduction de travaux en russe; Louis Trotier; J. Warkentin.

Users' Guide Guide de l'usager There are fifteen major divisions of these volumes, one for each of the thirteen 'Sections' of the International Geographical Congress Programme in Montreal, a fourteenth for all Commission meetings held in conjunction with the Congress, and a fifteenth for all Symposia. These major divisions are listed in the Contents. The Symposia and Commission sections are subdivided by individual symposium and commission. Papers are designated by the division, or subdivision, and the number of the paper within it. Because of the exigencies of the processing and printing schedules, the papers are not in alphabetical order within divisions, although they may, in places, appear so. The first thirteen divisions have the prefix p (for Programme) so that, for example, pOl is the first division (Geomorphology) and Pl3 is the thirteenth division (Remote Sensing). Thus paper P0127 is the 27th paper in Programme Section 1. Commission papers bear the prefix c, and Symposium papers the prefix s. Thus, c0113 and c0801 are respectively the 13th paper in Commission 1 and the first paper in Commission 8. The titles of the thirteen Programme divisions are self-explanatory and provide the simplest means of access to the papers in the collection. At the beginning of each of these divisions is a list of the papers which it contains, in numerical order, by short title and first author. Contributors were asked to include key words in their titles so that a scan of these lists should provide a useful indication of the content of the papers. In addition, at the beginning of each major division, we have noted groups of papers located elsewhere in the volumes, of interest to readers of that division. For example, at the beginning of POl (Geomorphology), we draw attention to s05 (Karst Geomorphology), a collection of papers devoted to this one geomorphological topic. To complement the list of titles and authors at the beginning of each division, we have included two indexes at the end of each volume. One is an index to paper number by author and co-author, and the other an index

Ces volumes sont divisés en quinze grands chapitres. A chacune des treize 'Sections' du Programme de Montréal correspond un chapitre; tous les textes soumis en vue des Commissions se trouvent dans le quatorzième chapitre; le quinzième chapitre comprend les textes soumis à tous les Symposiums dans le cadre du Congrès. Une liste de ces chapitres apparaît dans la Table des matières. On a assigné un sigle à chaque texte selon le chapitre du livre où il se trouve et selon son numéro d'ordre à l'intérieur de ce chapitre. A cause de l'énormité du travail et des échéances, les textes ne sont pas en ordre alphabétique dans chaque chapitre de l'ouvrage. Les treize premiers chapitres sont identifiés par le sigle P, pour Programme: pOl, par exemple, réfère au premier chapitre, qui est réservé à la géomorphologie, et P13 réfère au treizième chapitre, qui porte sur la télédétection. La communication P0127 est donc la vingt-septième dans la Section 1 du Programme. Les communications relevant des Commissions sont identifiées par la lettre c, et celles des Symposiums par s. Ainsi, c0113 et c080l sont respectivement la treizième communication de la Commission 1 et la première de la Commission 8. Les titres des treize premiers chapitres ne requièrent aucune explication; ils permettent de trouver sans difficulté les textes qui se trouvent dans l'ouvrage. Au début de chaque chapitre, il y a une liste des textes par numéro d'ordre d'après le titre abrégé et l'auteur. On a demandé aux collaborateurs d'inclure des mots clés dans le titre, ce qui permettra de déterminer plus facilement, en examinant la liste, de quel sujet traite chaque étude. De plus, au début de chaque chapitre, on mentionne les communications qui se trouvent dans une autre section de l'ouvrage, mais qui peuvent intéresser le lecteur de ce chapitre. Au début de p0l de géomorphologie, par exemple, nous attirons l'attention sur s05 (Géomorphologie karstique), où l'on trouve une série d'études traitant spécifiquement de ce sujet de géomorphologie.

Guide de l'usager I ix of the geographical locations to which papers relate. To save space, we have reduced all acknowledgments to a simple list immediately following the text of an article. Individuals' titles and explanations of the contributions of those acknowledged have been deleted. It should be noted that in some cases the 'references,' which follow immediately after the 'acknowledgments,' may be references sensu stricto, whereas in other cases they might be better described as 'general bibliography' in that they may contain more than the references actually cited.

Afin de compléter la liste des titres et des auteurs au début de chaque chapitre, nous avons inclus deux index à la fin de chaque volume. Le premier est un index par auteur et coauteur; l'autre index enumere les lieux géographiques dont traitent les communications. Pour économiser l'espace, tous les remerciements se réduisent à une seule liste qui paraît immédiatement à la suite du texte. On a retranché les titres des personnes et les détails concernant la contribution de ceux qu'on remercie. Dans certains cas, il est fort possible que les références apparaissant à la suite des 'remerciements' soient des références sensu stricto-, dans d'autres cas, on devrait plutôt les qualifier de 'bibliographie générale,' car on y trouve d'autres références que celles qui avaient été citées par l'auteur.

To find a topic or paper of interest to you 1. Scan the Contents page for the most appropriate major division of the volumes. 2. Scan the list of titles and authors at the beginning of that division. 3. Find the paper by its number (e.g. P0101, P0102 are the first and second papers in the first division). To find an author use the author index at the end of each volume, where the papers are cited by number as in item 3 above. To find papers related to a particular geographical location use the locational index at the end of each volume, where papers are cited by number as in item 3 above. Note that the bibliography following a paper may often contain more than references actually cited.

Pour trouver un sujet ou une communication qui vous intéresse 1. Examinez la Table des matières afin de trouver le chapitre approprié. 2. Examinez la liste des titres et des auteurs au début de ce chapitre. 3. Trouvez le texte que vous cherchez en utilisant le sigle (p.e. p0101, P0102 sont les première et deuxième communications dans le premier chapitre). Pour trouver un auteur, consultez l'index par auteur à la fin de chaque volume où les textes sont enumeres comme dans la section 3 susmentionnée. Pour trouver des communications traitant d'un lieu géographique spécifique, consultez l'index des lieux à la fin de chaque volume où les textes sont enumeres comme dans la section 3 susmentionnée. A noter: i/ est fort possible que la bibliographie à la fin d'un texte contienne d'autres références que celles qui avaient été citées par l'auieur.

Contents Table des matières VOLUME 1

1 2 3 4 5 6 7 8 9

Préface / Préface v Users' Guide / Guide de l'usager viii Geomorphology / Géomorphologie 1 Climatology, Hydrology, Glaciology / Climatologie, Hydrologie, Glaciologie 127 Biogeography and Pedology / Biogéographie et Pédologie 245 Regional Geography / Géographie régionale 321 Historical Geography / Géographie historique 393 Cultural Geography / Géographie culturelle 473 Political Geography / Géographie politique 499 Economie Geography / Géographie économique 525 Quality of the Environment / Qualité du milieu 629

VOLUME 2 10 Agricultural Geography and Rural Settlement / Géographie agraire et Peuplement rural 695 11 Urban Geography / Géographie urbaine 789 12 Geographic Theory and Model Building / Théorie géographique et Elaboration des modèles 891 13 Remote Sensing, Data Processing, and Cartographie Présentation / Télédétection, Traitement des données et Représentation cartographique 963 14 Commissions / Commissions 1001 15 Symposia / Symposiums 1231 Addenda 1351 Index of Authors and Co-authors / Index des auteurs et coauteurs xiii Selected Index of Locations / Index des noms de lieux xxi

P01 Géomorphologie Geomorphology CONVENOR/CONVOCATEUR: Derek C. Ford, McMaster University, Hamilton This listing of short titles and first authors' surnames will assist in identifying articles, topics, and places of interest. A complete author and co-author index is located at the end of this volume, as well as a selected index relating papers to geographical locations. Note that the papers are not listed in alphabetical order. The organization of the volumes is described in full in the Preface. Other papers of particular interest to readers of this section are to be found under c03 (Coastal Geomorphology), c04 (Présent Day Geomorphological Processes), c08 (Geomorphological Survey and Mapping), and s05 (Karst Geomorphology). These are in thé second volume.

P0101 P0102 P0103 P0104

Cette liste des titres abrégés et des noms des auteurs principaux permettra d'identifier les communications, les sujets et les lieux qui présentent un intérêt quelconque. Un index complet d'auteurs et de coauteurs se trouve à la fin de ce volume, ainsi qu'un index des lieux géographiques. Prière de noter que les textes ne sont pas classés par ordre alphabétique. On explique en détail le plan de ces volumes dans la Préface. D'autres études qui intéresseront peut-être les lecteurs de cette section se trouvent dans c03 (Géomorphologie côtière), c04 (Les processus géomorphologiques actuels), c08 (Recherche et cartographie géomorphologiques) et s05 (La géomorphologie karstique). Elles font partie du deuxième volume.

Slope process and slope form: a theoretical study AHNERT 3 Chemical weathering of tills and surficial deposits, east Baffin Island ANDREWS Structural geomorphology around Shillong, India BANDYOPADHYAY 7 Repeated ground photogrammetric survey for studying slope dynamics (USSR) BLAGOVOLIN

9

P0105 A process model for barrier evolution CARR 11 P0106 Quaternary geomorphology of Aconcagua Valley (Chile) CAVIEDES 13 P0107 Cryopedimentation: a type of slope development in cold environment DEMEK P0108 Inferring process from form: asymmetry of glaciated mountains (Canada) P0109 pOUO P0lll P0112 pO113 P0114 P0115

EVANS

5

15

17

Objectives and methods in study of glacier depositional landforms (Canada)

FALCONER

19

Dating cave calcite, uranium disequilibrium method, Alberta FORD 21 Morphogenetic interpretation of a mountain section, Mexico GRANIEL-GRANIEL 23 Postglacial rock wall recession, Ogilvie and Wernecke mountains, Yukon GRAY 24 Hillside slope forms and their evolution (Japan) ICHIKAWA 26 Periglacial slopes JAHN 28 Variations in degradation of ice-cored moraines, St Elias range, Yukon JOHNSON 29 P0116 Planetary and hypsometric variation of valley asymmetry KARRASCH 31 P0117 Regularities in cryogenic phenomena development (Arctic) KATASONOV 34 P0118 Hypothesis on paleomorphology in south-central Mexico LOPEZ-SANTOYO 35 P0119 Talus slope debris accumulation, Surprise Valley (Alberta) LUCKMAN 36 P0120 Wind factor in dune formation at Great Whale, Quebec MacFARLANE 38 P0121 Cassures périodiques et modelé éolien dans les grès péri-tibestiens (Tchad) MAINGUET-MICHEL

40

P0122 Magnitude and frequency of processes, arctic beaches, Queen Elizabeth Islands, Canada MCCANN 41 P0123 Quantitative aspects of periglacial slope deposits, southwest England MOTTERSHEAD

43

P0124 Morphology of lunar formations MUKERJEE 45 P0125 Pediment formation in Mojave Desert, California OBERLANDER

47

2 / Geomorphology P0126 P0127

Hillslope evolution on Nsukka Plateau, eastern Nigeria OFOMATA 49 Classification of alluvial plains based on geomorphological characteristics (Southeast Asia) OYA 51 P0128 Tectonic control of morphology, Canadian interior plains OZORAY 51 P0129 Main types of landslides PECSI 54 P0130 Solifluction rates, Ruby Range, Yukon Territory PRICE 56 P0131 Genetic classification, major types of sinkholes and related karst depressions QU1NLAN

58

P0132 Anomalous behaviour of model rivers SCHUMM 60 P0133 Erosion surfaces, lower Chambal Valley, India SHARMA 61 P0134 Rock glaciers, Aquarius Plateau, Utah SHRODER 63 P0135 General regularities, relief development, Ukrainian SSR SOKOLOVSKY 65 P0136 Rock control in coastal erosion, Miura Peninsula, Japan SUZUKI 66 P0137 Landslide damage in mountain areas and its control, Tanzania TEMPLE 68 P0138 Litho-control of morphometry, Upper Gondwana, central India VERMA 70 PO139 Relief evolution of old mountains of central Europe in the Tertiary - the Sudetes WALCZAK 71 P0140 Fluvial cave sediments: description and interpretation (West Virginia) WOLFE 74 P0141 Classification of Martian terrain features on Mariner 6 photos WOODRUFF 76 P0142 Reworking of shores in permafrost zone (USSR) ARE 78 P0143 Coupe transversale de la moraine terminale de Saint-Narcisse, Québec DENIS 79 P0144 Wind in periglacial environments, Banks Island, Canada FRENCH 82 P0145 Beach erosion, Durban (South Africa) NAIDOO 84 P0146 Formes particulières de l'érosion différentielle dans les tillites (Afrique) ROGNON 86 P0147 L'aggravation de l'érosion dans l'Ouarsenis (Algérie) SARI 87 P0148 Problems in correlation of landslide movement and climate (New Zealand) CROZIER 90 P0149 A neglected landform: the meteorite crater (Canada) DOJCSAK 93 P0150 Karsts de type tropical sous climat tempéré (France) FENELON 93 P0151 Geologic and physiographic control of karst landscapes, Cuba PANOS 94 P0152 Les terrasses d'abrasion de la côte du Chili semi-aride PASKOFF 96 P0153 Erosional features due to piping, Venezuela PEETERS 98 P0154 Simulation model for landslide prediction - coast of Normandy, France ROSENFELD

98

P0155 Complex rapids (sula complexes) in tropical rivers (Guyanas) ZONNEVELD 100 P0156 Effects of lithology and time on slope characteristics (USA) BARISS 100 P0157 Principaux résultats géomorphologiques du projet Hudsonie (N. Canada) CAILLEUX 104 P0158 L'exportation de produits en solution par cours d'eau appalachiens, Québec, Canada CLEMENT

105

P0159 Dissection des dépressions fermées : les kaluts du désert du Lut (Iran) COQUE 106 P0160 Le système morphogénétique 'géodynamique' (Amérique du Sud) DEMANGEOT 108 P0161 Les phénomènes cryonivaux et glaciaires, marges du domaine méditerranéen DRESCH 110 P0162 Etude de la denudation chimique des pays montagneux (Caucase) GABRIELIAN 112 P0163 Simulations en laboratoire au centre de géomorphologie du CNRS à Caen (France) JOURNAUX 114 P0164 Surfaces structurales et surfaces pseudo-structurales (Belgique) MACAR 116 P0165 Moraine de poussée valders (dryas supérieur) à Saint-Narcisse, Québec P0166 P0167 P0168

OCCHIETTI

117

Contemporary denudation in Lake Ontario watersheds ONGLEY 119 Le problème des paléo-deltas quaternaires en Roumanie POPP 121 Isostatic readjustments, their rejuvenation morphology, central Himalaya KAUSHIC 122 P0169 Mouvements de sols gelés subissant des variations de température sous 0° PISSART 124

Géomorphologie / 3 P0101 Slope process and slope form: a theoretical study FRANK AHNERT University of Maryland, USA

Many slopes are polygenetic, not only in the sense of having been affected by different processes at different times in their morphological history, but also in the sense of being shaped by combinations of several types of denudational processes at the same time. Because of this complexity, the effects of individual processes upon the slope form are difficult to assess by empirical field investigation alone. However, isolation of a specific process, and of its particular morphological effect, is possible through theoretical abstraction by means of quantitative models (cf. Lehmann 1933; Bakker and Le Heux 1947, 1953; Scheidegger 1961, 1964, 1970; Young 1963; Gossmann 1970; Ahnert 1964, 1971). The present study uses a new general and comprehensive model of slope development for this purpose. The model has been designed as a Fortran computer program called COSLOP (Ahnert 1971) that treats the processes of fluvial downcutting, of waste production by bedrock weathering, of waste removal, and the changing shape of the slope profile itself as functionally interdependent components of a system. For each of the process groups, the model provides several different options that permit its adaptation to a wide range of environmental conditions. Care has been taken to distinguish those aspects that are merely peculiarities of the model design, and hence properties of the model, from the actual properties of the slopes and slope processes which the model seeks to represent. The present study concerns itself specifically with the effects of three types of waste transport upon the slope profile, namely, slope wash, creep by expansion-contraction cycles, and viscous-plastic flow. Of these, slope wash is expressed as a removal of waste from each point on the model profile : (1) R = kV j (1. 0 + dj.0.6) where R = amount removed from point j in a given time unit; k = a coefficient whose value is proportional to the local thickness (and by implication to the comminution) of the waste; Vj = a quasi-velocity at point j, proportional to the sine of the local slope angle; d} = relative distance from the summit to point / (0.0 < d < 1.0). Equation (1),

which was designed mainly after Zingg ( 1940), causes the rate of removal to increase with increasing steepness, increasing distance from the summit, and decreasing particle size. Creep is defined by (2) # = /itana y where h = coefficient representing expansion normal to the slope surface; o^ = local slope angle. In contrast to slope wash, creep is programmed as a point-to-point transfer of waste; while point / loses the waste amount R, an amount (3) A = hümaj + l is supplied to / by creep from the next higher point / + 1. Viscous-plastic flow also transfers waste from point to point, and has been programmed, in accordance with the theory of Souchez (1964), as (4) R = b[Vj-(alCj)] where b = coefficient of fluidity (the inverse of viscosity) ; a = coefficient of cohesion; C¿ = waste cover thickness at point /. In all models used here, the waste supply by bedrock weathering varies as a function of waste cover thickness. The specific rate of weathering has no influence upon the rate of creep as long as there is more waste present at each point than can be removed in one time step. For wash and for viscous-plastic flow, however, the waste thickness, and through it the rate of weathering, affects the magnitude of the waste movement. The effect of these processes upon the profile shape will first be investigated on slope models with fixed base level, and then on steady state models. Fixed base level models of slope development start off with an arbitrary initial profile, the shape of which has, however, little or no influence upon the later phases of profile transformation. From different initial profiles, but under identical process conditions, the profile transformations converge towards the same characteristic shape. The latter is always entirely concave for slope wash, and always convex for creep. The basic profile shape for viscous-plastic flow is convex-concave (XV\ Ahnert 1970), but the relative extent of the upper convex and the lower concave segment,

4 / Geomorphology their curvatures, and their steepness varies with the cohesion coefficient a, if the fluidity coefficient b is held constant. The profile is predominantly convex, and thus similar to a creep slope, when cohesion is low (a — 0.2) ; it is more symmetrical (X5V5) when cohesion is moderate (a = 0.5) ; and it is predominantly concave (X2V8), close to the shape of a wash slope, when cohesion is high (a = 1.0; in all these cases b = 0.5). This last similarity is one of shape only, and does not include the distribution of the waste cover thickness over the profile; for viscous-plastic flow the thickness increases downslope; for wash it does not. The convex shape of creep slopes and the concave shape of wash slopes have also been derived recently in a theory by Kirkby (1971). However, the range of different model profiles obtained with viscous-plastic flow indicates that one particular process type can cause widely different shapes, and, on the other hand, that different processes can lead to very similar shapes. The attainment of a steady state profile requires that the primary energy input, here represented by downcutting at the slope foot, be uniform through time. In the original COSLOP model, this was accomplished by a constant-rate lowering of the lowermost bedrock point on the profile. Since then, it was found that constant-rate lowering of the lowermost waste surface point is preferable, as it more accurately represents an isentropic condition. Penck (1924) postulated that in the case of uniform development the slope profile should be rectilinear. The model shows that this applies of necessity only when the rate of downcutting is too great to be ever equalled by the rate of bedrock weathering, so that the slope stands at the maximum possible angle permitted by the properties of the bedrock, and downslope waste transport consists mainly of landslides (field example: canyon of the Yellowstone R., Wyoming). However, when the uniform rate of downcutting is small enough to be eventually equalled by the rates of weathering and of surface lowering on the slope, the steady-state profile is convex for the point-to-point transport types (creep and viscous-plastic flow), and concave for slope wash. The last of these results appears at first

sight particularly surprising, as the association of concave slope shape with uniform downcutting has only rarely been made in geomorphology. Yet it follows necessarily out of equation (1 ), in which for steady-state conditions (R and k held constant) V^ and with it the slope angle, decrease progressively downslope as a function of d. Sufficiently slight but uniform downcutting and sufficiently intensive wash combine to maintain a very gentle, concave steady-state model profile for which the rate of weathering, of removal, and the regolith thickness all remain constant and equal at every point. The model thus confirms in quantitative terms the theory of Biidel (1957) that tropical erosion surfaces can develop by 'double planation' (i.e., synchronous and equal lowering of a sheet wash surface and of its subterranean weathering front) under conditions of steady state. This paper has shown that there are definite relationships between denudational process type and profile shape. Although emphasis was placed here upon the effect of individual processes, the COSLOP model program lends itself also to the simulation of polygenetic slope development with downcutting in several phases, and with combinations of different types of weathering and of waste transport taking place simultaneously or in succession. The 'real world' of actual field slopes can thus be more closely approximated. Computer Science Center, University of Maryland. Ahnert, F., 1964 Quantitative models of slope development as a function of waste cover thickness, Abstracts of Papers, 20 IGC, London, 118. - 1966 Zur Rolle der elektronischen Rechenmaschine und des mathematischen Modells in der Géomorphologie, Geog. Zeitschr. 54, 118-33. - 1970 An approach towards a descriptive classification of slopes, Zeitschr. f. G com. Suppl. 9, 88-101. - 1971 A general and comprehensive theoretical model of slope profile development. U. Maryland Occ. Papers in Geog. no. 1. Bakker, P., and J.W.N. Le Heux, 1947, 1953 Theory of central rectilinear recession of slopes, Kon. Ned. Akad. van

Géomorphologie I 5 Wetensch. Proc. 50, no. 8 and 9, and Proc. 54, no. 7 and 8. Biidel, J., 1957 Die 'doppelten Einebnungsflachen' in den f euchten Tropen, Zeitschr. f. Geom. 1, 201-28. Gossmann, H., 1970 Theorien zur Hangentwicklung in verschiedenen Klimazonen. Würzburger Geog. Arbeiten no. 31. Kirkby,M., 1971 Hillslope process-response models based on the continuity equation, in D. Brunsden, éd., Slopes: form and process, IBG Spec. Pub. no. 3, 15-30. Lehmann, O., 193 3 Morphologische Théorie der Vertwitterung von Steinschlagwanden, Vierteljschr. d. Naturf. G es. Zurich 78, 83-126. Penck, W., 1924 Die morphologische

Analyse (Stuttgart). Scheidegger, A., 1961 Mathematical models of slope development, Bull. Geol. Soc. Am. 72, 37-50. - 1964 Lithologie variations in slope development theory, uses Circ. no. 485. - 1970 Theoretical geomorphology 2nd ed. (Springer). Souciiez, R., 1964 Viscosité, plasticité et rupture dans l'évolution des versants, Ciel et terre, 389-410. Young, A., 1963 Deductive models of slope evolution, Nachr. d. Akad. d. Wiss. Gottingen, II Math.-Naturw. Kl. 5,45-66. Zingg, A.W., 1940 Degree and length of land slope as it affects soil loss in runoff, Ag. Eng. 21, 59-64.

P0102 Chemical weathering of tills and surficial deposits in east Baffin Island, NWT, Canada J.T. ANDREWS and G.H. MILLER University of Colorado, USA Evidence for chemical weathering in the eastern Canadian Arctic has not been extensively documented, in part because of the overwhelming visual impact of the effects of mechanical weathering. It is generally recognized (Carroll 1970) that chemical weathering will be slow in polar regions because of the combination of low temperatures, limited precipitation, and rapid run-off. However, the very fact of the limited amount and rate of weathering does have some advantages from the viewpoint of our study because our interest in chemical weathering is in large measure concerned with developing calibrated weathering curves that can be used as 'geological clocks' to date glacial deposits. Along the uplifted rim of the eastern Canadian Arctic - from the Kiglipait Mountains of Labrador (ca 56°N) to the fiord country of east Baffin Island (ca 70°N) evidence gathered in the 1950s and 1960s (Ivés 1963) has suggested that multiple glaciations can be recognized on the basis of differences in the weathering of the surficial deposits. Three broad weathering zones have been delimited (Pheasant 1^72) : Zone i is composed of in situ weathering products and is characterized by mountain-top tors; there is no evidence that these mountains have ever been overrun by active glaciers. Zone n is composed of a weathered till; oxidization is

apparent in the form of weathering rinds; micro-pitting has developed on many boulder faces and feldspar crystals project from the surface of boulders. Many boulders disintegrate with a strong hammer blow. Zone ra is on the other hand composed of a fresh till; boulders are smooth and do not disintegrate and there are no weathering rinds. Exceptions to these statements are considered to represent the inclusion of material of Zone n into the till of Zone in. The bedrock is predominantly granite gneiss. Until recently the relative age of these units was not known and even now the age of Zone in is the only one dated with any precision. Uranium series dates on marine shells related to a glacier advance indicate that Zone in spans the interval between 120,000 ± 10,000 and the present. Although weathering differences can be recognized within this zone they are relatively small compared with the change between Zone in and Zone n. Various ratios based on weathering criteria (such as the thickness of weathering rinds, percentage of fresh boulders, etc.) indicate a weathering ratio of 5:1 for Zone ii/Zone in. As these two lower zones are delimited by moraines and glacial features we discount any suggestion that the difference is related to vertical changes in the rates of weathering (e.g., Dahl 1966). In the Colo-

6 / Geomorphology rado Rockies various weathering ratios on Bull Lake and Pinedale deposits (Miller 1971 ) indicate as a first approximation that weathering ratios are a linear product of the age of the deposits. The significant increase in weathering between Zones n and ra prompts the suggestion that Zone n may reflect weathering processes operating over 500,000 years or so. Soil samples from the different zones were collected and have been analysed by a variety of methods, including mechanical grain-size, percentage composition of the major elements, content of free iron, x-ray diffraction of oriented samples of material less than 2/¿, and scanning electron microscope (SEM) examination of sand-size quartz grains. One potentially serious problem was that local sources of bedrock could drastically have altered the percentage contribution of major elements in the surficial samples. Thus changes in percentage composition would reflect changes in bedrock sources rather than in the processes associated with chemical and associated physical weathering. As most tills are composed predominantly of local rock, attention was paid to the local geology, and this is used as a group classification in the study to ascertain if it has any significant impact on the result. Multiple stepwise discriminant analysis (MSDA) was used to enquire whether the chosen classification of the samples (i.e., classified as belonging to Weathering Zone i, n, or in) was statistically appropriate. This technique was chosen over factor analysis because we are not primarily interested in generating a classification but rather in evaluating the correctness, or otherwise, of an existing scheme. The percentage composition of eight major compounds was determined; they were MnO, MgO, CaO, A12O3, SiO2, Fe2O3, TiO2, Na2O. In the analysis six of the results were used in straight percentage form, whereas A12O3 and SiO2 were combined to form a ratio by: (percent SiO2/ molecular weight SiO2)/(percent A12O3/molecular weight A12O3). The percent of silt and clay was also used as a variable in the discriminant analysis. Numbers of samples of indisputable affiliation were, Zone i = 13, Zone n = 6, and Zone ra == 6. Another 17 samples were available and were from Zone i or Zone n, but assignment to one or the other unit was not certain. Part of our aim was to develop a discriminant

function that would assign these samples to one or other zone. MSDA indicates that only 2/25 of the samples in the three zones were not classified correctly. The stepwise part of the program indicated that the most useful variable to discriminate between the three groups was the molecular ratio of SiO2/A12O3. On the young moraines of Zone ra this ratio amounted to an average of 9.34:1, in Zone n it was 7.57:1, and in Zone I it averaged 7.35:1. These results indicate that silica is moderately mobile in these arctic environments - a statement in agreement with the results of Church (1971 ) on the composition of run-off from certain rivers on Baffin Island. The other compounds showed an increase with age; this was especially noticeable in the case of MnO and Fe2O3; the latter increased from an average of 3.7 per cent in Zone in to 4.84 in Zone i. The percentage of silt and clay decreases with age, probably because of the addition of granular disintegration products. Iron oxidization is visually important on Baffin Island and is of potential geochronological importance. Analyses have been undertaken on the percent of free iron in various soil horizons (A, B, and C). The results indicate that on young, neoglacial moraines, the percentage varies between 0.1 and 0.3; on till of late-Wisconsin age (8000 BP) values of 0.4 to 0.6 are more common in the B horizon; in moraines about 25,000 BP the amount has risen to 0.9 and by 45,000 BP the figures can reach 1.2 per cent. A general feature is that the A horizons all have slightly greater percents of free iron than the B or C. Hydrolysis acts very slowly in the area. X-ray diffraction of samples from Zone m indicates little or no development of clay minerals: the only peaks are small and limited to illite/muscovite and hydrobiotite. In Zones n and in the diffraction patterns are more complex and peaks are larger with a greater number of significant peaks. In particular materials from both zones have significant peaks showing the presence of hydrobiotite and mixed layer chlorite /vermiculite at 7.5 to 7.7 and 6.31 20 respectively. These peaks are not present in Zone in soils. Finally, detailed examination of quartz grains with the SEM system at 500 to 50,000 magnification was undertaken. In Zone ra the surface is fresh and exhibits typical glaciated

Géomorphologie I 7 patterns (Krinsley and Margolis 1969). The quartz surfaces in Zone n material still show the glaciated fracture patterns, but pitting of the surface is evident; and by Zone i there is no evidence of glaciated surfaces and the surfaces show considerable pitting and wasting. Chemical weathering clearly occurs in the eastern Canadian Arctic but rates are extremely slow. Only limited leaching and oxidation has taken place in the last 120,000 years, but in deposits older than this there has been considerable removal of silica and a concentration of certain oxides. Hydrolysis is limited and is largely limited to rudimentary clay mineral development occasioned by the weathering of mica and the formation of illite, hydrobiotite, and mixedlayer clays. Montmorillinite does occur in certain Zone i samples. Graphs of rates of increase or decrease of the various aspects of chemical weathering will be presented. Although more control is needed, results to date warrant further research and indicate that chemical changes can be used to distinguish between the ages of surficial deposits in this area.

National Science Foundation, grant no. GA-10992; Geological Survey of Canada; J.D. Ivés. Carroll, D., 1970 Rock Weathering (New York). Church, M., 1970 Baffin Island Sandur: a study of Arctic fluvial environments. PhD thesis, U. British Columbia. Dahl, R., 1966 Block fields and other weathering forms in the Narvik Mountains, Geog. Ann. 48, 224-7. Ivés, J.D., 1963 Field problems in determining the maximum extent of Pleistocene glaciation along the eastern Canadian seaboard, in North Atlantic Biota and Their History (Pergamon), 337-54. Krinsley, D., and S. Margolis, 1969 A study of quartz grain surface textures with the scanning electron microscope, NY Acad. Sci. ser. 2, v. 31,457-77. Miller, C.D., 1971 Quaternary glacial events in the northern Sawatch Range, Colorado. PhD thesis, U. Colorado. Pheasant, D.A., 1972 Glacial history of Narpaing and Quajon fiords, northern Cumberland Peninsula, Baffin Island, NWT, Canada. PhD thesis, U. Colorado.

P0103 Structural geomorphology of the area around Shillong, India M.K. BANDYOPADHYAY and GOURi BANERJEE Calcutta University, India The paper deals with the topography of the area around Shillong (the capital of Meghalaya, India) and its relationship with the underlying structure. The area is located in the central part of the Shillong Plateau. The plateau as a whole is regarded as an autochthone, composed mainly of crystalline and metamorphic rocks of pre-Cambrian age which form a detached part of the Deccan Shield. Geologic structure is a dominant control in the evolution of landf orms in the area. The major portion of the area is covered by the rocks of the Shillong series which includes quartzites with slates, schists and conglomerates, with occasional intrusions of epidiorites, locally known as Khasi greenstone. The Shillong series belongs to the Dharwar system of Archean (Huronian) age. Quartzites predominate. Mylliem granite is intruded into

the Shillong series and covers a wide area to the south of Shillong. It is younger in age than the Shillong series. Diorites occur as xenoliths in some places in Mylliem granite, e.g., on the Shillong-Cherrapunji road about 18km to the south of Shillong. The grains of Mylliem granite are coarser and the colour is pinkish. Fracturing and brecciation are common along the contact zone. The general trend of the crystalline and metamorphic rocks is in a NE-SW direction. As a result, planar structures such as gneissosity, schistosity, etc., are also oriented NESW. The rocks are tightly folded and lineated (Bhattacharyya 1968). Folds in the Shillong series around the Mylliem granite have their axes towards the granite. The trend of the rocks around the margins of granite is nearly parallel to the outer boundary. This structural relationship between the rocks of the Shillong

8 / Geomorphology series and the pluton encircled by it is distinctly visible at several places along the Shillong-Cherrapunji road. Folds near the contact zone are nearly concentric in layout and the rocks are highly metamorphosed, schists being more common than shales and slates. Younger rocks of Tertiary and pre-Tertiary age occur in the southern section. These rocks lie almost horizontally with a gentle inclination to the south. However, near the southern margin of the plateau, they are steeply inclined towards the south and plunge down beneath the alluvium of the Surma valley of East Pakistan. The most important river in the region is the Wah Umiam which flows through the northern section of the area. The direction of flow is from sw to NE which coincides with the general trend of the structure. The river may therefore be regarded as well adjusted to the structure. Several small streams originate in the Shillong Peak region and combine together to form a larger stream called Umban. It flows in a westerly direction and produces a series of waterfalls. The rocks here dip in the upstream direction and hence the stream may be regarded as an anti-dip stream. Near the southern boundary of the Shillong series also the rivers form a large number of waterfalls, Here the pre-Cambrian rocks lie in juxtaposition with the Tertiary sedimentaries. It is evident that the pre-Cambrian rocks are much more resistant than the tertiaries and hence differential erosion has produced differential landf orms. The pre-Cambrian rocks form a gently undulating terrain at an ejevation of about 1800m while the land below the falls is highly dissected where the rivers form abnormally deep trenches separated from each other by flattish interfluves. The Tertiary sedimentarles include sandstones, limestones, and shales which are almost horizontal or are slightly inclined to the south. The famous meteorological station of Cherrapunji is located in one of these interfluves which, with their striking 'accordance of summit levels' misled many geomorphologists in the past who considered them as parts of a dissected peneplain. However, the rock strata are almost horizontal in the area and there is no evidence of truncation of rocks of varying resistance. It is therefore more reasonable to

consider these surfaces as the erosiona! remnants of a once-continuous structural bench. Also it is to be noted that the area receives the highest amount of rainfall in the world and that, too, comes as torrential downpours concentrated during the three summer months only. This, no doubt, has tremendous effects upon the sculpturing of the landf orms. The lOOOm-deep gorges with almost vertical walls resemble those of the Grand Canyon of the western USA. The horizontal structure with alternate hard and soft rocks and the peculiar climate have resulted in more rapid vertical downcutting than horizontal widening. The pre-Cambrian terrain is supposed to be a part of an ancient peneplain which was uplifted and dissected after the deposition of the Tertiary rocks in the south. A new cycle of erosion started in the south and the rivers continued to extend headward, i.e., northerly. The boundary between the highly dissected terrain and the less dissected terrain has been shifting towards the north, i.e., towards Shillong. At present the boundary lies at a distance of about 15km to the south of Shillong where even a non-geographer tourist can note the surprising contrast between the two different types of landform. The river Umbanium originates on the southern slope of the Shillong Peak (1963.2m), the highest point in Shillong Plateau. For about 5km it flows parallel to the Umban and follows exactly the contact between the Mylliem granite and the quartzite. This is the zone of structural weakness. The quartzites lie to the north of the granite and form the higher lands. It appears that the river was initially located on the quartzite upland and gradually, by the process of lateral shifting, moved further south into the zone of weakness. At present the river is well adjusted to the structure and further southward shifting olthe course does not seem to be-likely. It is evident from the above discussion that the structure has distinct influence upon the landform and drainage system of the area. L. Starkel; W.D. Thornbury; B. Zakrzewska. Bhattacharyya, C.C., 1968 Structure and petrology of the Shillong Plateau. Proc. Phys. Geog. Eastern Himalaya, IGU, Gauhati.

Géomorphologie I 9 PU 104

The use of repeated ground photogrammetric survey for studying the dynamics of slopes N.s. BLAGOVOLiN and D.G. TSVETKOV Academy of Sciences of the USSR Stereophotogrammetric methods which compare the results obtained from repeated surveys are at present in wide use for studying the dynamics of natural processes. The choice of a particular method depends on specific features of the object under investigation. Thus, for example, the analytical method which allows full use of the accuracy of stereophotogrammetry is expedient when observing space dislocation at separate contour points or when studying the movement rate of avalanches, landslides, glaciers, etc. With this method, both a common Stereophotogrammetric survey from a 'linear' baseline and in some cases a repeated survey from a 'temporary' point (the so-called method of parallaxes) can be used. The Institute of Geography of the Academy of Sciences of the USSR has been studying the dynamics of slopes of the mountainous Crimea with the help of repeated ground Stereophotogrammetric survey (GSS) since 1965. Elementary forms of relief (separate peaks, ridges, residual mountains, parts of slopes of various kinds) were chosen as objects for GSS; the aim of the work was to obtain the most accurate quantitative evaluations of their development possible. During 1965-69 there were conducted repeated GSS of 12 portions of slopes characterized by different steepness, exposure, lithology, and so on. The present paper discusses methods and the results of repeated GSS of two sites which differed in morphology and in scale and activity of exogenic processes. The first site was an avalanche-fall slope of the Demerdgie Mountain. It is a flattopped massif limited by breaks and belonging to the main (southern) ridge of the

Crimean mountains. It is composed of conglomerates and sandstones of the Jurassic (Oxford-Kimeridglan stage) period. As a result of differential fault block movement there is a sharply contrasted relief with great height amplitude. Slope processes are extremely active. There are grounds to suppose that even at present the faults maintain their activity. Continuing landslides on the slope (the total height of which in this part reaches 500m) are evidence of this. The uppermost part is a sheer precipice with rocky ledges of varied shape; it is a zone of gravitational break-away devoid of friable slope deposits. The middle part of the slope (denudation and transit zone) where only separate outcrops of bedrock occur has a steepness of 40-45°; the lower transit-accumulative part of the slope is equally steep. In the lowest part, the slope gradually converts into a gently sloping, almost horizontal, site at the foot of the massif upon which the surveying baseline was pitched. Slope deposits in the zone of transit and accumulation are primary products of physical weathering of bedrocks, mainly conglomerates. The size of fragments ranges from blocks several metres in diameter to fragments of several centimetres. Fine material is practically absent. GSS from a baseline 91.5m long which was at a distance of 500m away from the slope foot were taken three times: on 01.10.1965, 20.09.1967, and 23.09.1969. The first survey was made a year before a big collapse of rock material occurred. The later surveys provided a possibility to evaluate quantitatively the development of the slope both as a result of the instantaneous fall and in 'still' conditions

TABLE 1

Denudation zone

Accumulation zone

Period

S AjfiTmean (1000m2) (m)

Ajy mean S AF (1000m3) (1000m 2) (m)

1965-67 1967-69 1965-69

3.8 3.8 3.9

-6.0 -3.8 -8.6

-1.6 -1.0 -2.2

2.5 4.8 4.7

+ 0.8 + 1.1 + 1.3

Balance of material

AF AF (1000m3;) (1000m3) + 2.0 + 5.3 + 6.1

-4.0 + 1.5 -2.5

10 / Geomorphology TABLE 2

(B/Y)

my (m)

Pattern scale (1:150

mAH (m)

1/5-1/8 1/5-1/7

+ 0.4 ±0.03

1 :2000 1:200

+ 0.4 ±0.04

Site

B (m)

Range Y

(m)

Basic condition

Demerdgie Badland

91.5

500-700 40-50

7,5

which ensued. Comparing surveys i and n established the landslide zone and also considerable changes in relief associated with it. Survey m showed that active development of the slope continued in the same zone. The surveys provided accurate quantitative deformation values for the slope (AH), selected areas of denudation and accumulation and the volume of the material pulled down and redeposited (AF). The four-year period of observations falls into two subperiods: 1965-67 with the big collapse of rock material in 1966, and 196769 when slow mass advances of the slope material prevailed. During the first period there were three distinct slope zones: (a) transit and denudation zone where the slope surface was lowered everywhere; (b) accumulation zone where there was aggradation; (c) 'neutral' zone where no essential changes occurred. Prevalence of erosion over accumulation was characteristic. From the denudation zone observed by the survey 6000m3 of rock was removed (the average layer 1.6m); extreme values of denudation at separate points were up to 13.5m. Of the total volume of 10,000m3 of rock that fell, only 2000m3 was detained in the accumulation zone. The remaining material displaced into a 'zone of invisibility', hidden behind trees at the massif foot. The second period was characterized by complex distribution and configuration of denudation and accumulation zones. The character of slope processes changed somewhat in comparison with the first period. Though the zone of denudation occupied the same area the average thickness of the pulleddown layer decreased to 1.0m. The area of the accumulation zone almost doubled. In summary, rapid gravitational processes had a decisive influence upon development of the denudation zone. In contrast, the relief of the accumulation zone formed primarily due to slow pulling down, sorting, and gravitational condensation of the material in time intervals between landslides.

The second survey site was a badland formed in loamy schists in the eastern part of the mountainous Crimea. Hilly relief of small amplitude is characteristic of this region. The system of thalwegs and interravine ridges is generally symmetric. The average inclination of thalwegs and ridges is close to 40°. Side slopes of ridges with different exposure have nearly the same steepness (45-50°). Analysis of GSS photographs showed that height changes at different points on slope surfaces reached 0.3—0.4m over two years, sw-facing slopes suffered the greatest change; SB-facing slopes displayed almost no change. There was axial displacement of thalwegs towards the southeast which is associated with intensive destruction of slopes of sw exposure. There was some lowering of inter-ravine ridges (down 0.1-0.2m), though it did not occur everywhere : the highest and the lowest parts of ridges were little changed. Watershed lowering was acompanied by displacement to the southeast. Gully heads receded rapidly. Slope surfaces were lowered at a rate of approximately 10cm per year. Though we deliberately chose the most dynamic site (friable rocks, absence of vegetation, sw exposure), it is however rather typical of the vast region of the eastern Crimea. Our experience of repeated phototheodolite surveys to study present-day slope processes in the Crimea indicates that this method is valuable. We recommend it for other mountainous countries. Results of the instrumental investigation of the dynamics of the surfaces of the slopes (A#) can be characterized by the following mean quadratic errors mAH in Table 2, where Y is distance in metres; B is length of the survey baseline; my is error of determining the distance. A detailed account of the survey methods, the results, and graphical appendices was published in the Geomorphology Journal, I, 1971.

Géomorphologie / 1 1 P0105 A process model for barrier evolution

ALAN p. CARR and JOHN R. HAILS

Unit of Coastal Sedimentation, England

According to Leontev (1965) and Zenkovitch (1967) between 10 and 13 per cent of the world's coastline is of barrier form (following Shepard's 1952 terminology). Therefore, the origin of such depositional features is of importance on the grounds of both their areal extent and the way they provide information about past and present geomorphological/ geological processes. Although several workers have discussed the origin and development of barriers, few publications have considered a barrier model in terms of 'climatico-geomorphic' processes. The purpose of this short paper is to outline some of the principal processes contributing to barrier development (and destruction), and to give examples. The inter-relationships are probably best explained by means of a model (Fig. 1 ), but even this involves some simplification. It may seem obvious that the origin and form of barriers are related to three principal interrelated factors - sediment supply, wave climate, and fluctuations in relative sea level. However, this relationship has more subtle

Fig. 1 (terminology based on Shepard 1952)

implications because of the paramount control effected by climatic change. The latter influences rates of weathering, river discharge, and ultimately sediment yield. Climatic and associated meteorological factors affect offshore wave height, frequency, direction, and profile (such as that imparted by wind-driven as against swell waves) both directly and indirectly through relative sealevel. The wave climate will determine the rate of movement of such material as is available and its direction onshore or offshore. However, variation in wave energy along any coastal sector is largely controlled by the topography of the nearshore zone. In addition, the form of the continental shelf and of earlier coastlines, especially in terms of rock outcrops, are both germane to the problem. The net effects of all these factors may well result in the formation of a barrier of some sort. What particular form such a feature will take is dependent upon the interrelationship of the various factors mentioned. Partial (or complete) destruction of older barrier shorelines may provide material for their successors. The continental shelf is another possible source of sediment if it has not been swept clean of unconsolidated material. Perhaps what is more important is the fact that the processes (and their consequent effects) contributing to the formation of one barrier shoreline are not always repeated at a subsequent date. The majority of the world's barrier shorelines are composed of quartzose sand, with associated tidal flat or lagoonal-salt marsh deposits generally falling within the silt-clay fraction. Barrier environments tend to be located along the margins of continents where swell waves are dominant. There is argument about the relative importance of terrestrial longshore and offshore (shelf) sources of material from which barriers are ultimately built. While heavy mineral analyses suggest an initial landward source for many sand barriers, other sources are equally important, and, especially in the British Isles, examples of analogous pebble structures occur although again with fine-grained deposits to landward. These barriers are frequently derived, in large

12 / Geomorphology measure, from relict deposits or residues of flint or chert. The source of such material must be related to the effects of glacial and inter glacial sea levels which have permitted a 'sweeping up' process over the relatively shallow sea bed close inshore (Hardy 1964). The pebble structures are clearly atypical of their landward environment. The comparatively high rates of flow of tidal currents in areas around the British coast (Stride 1963) would also have a bearing, at least in the case of finer material. Chesil Beach, Dorset, is a pebble structure some 18-28km long, while its evanescent precursor further seaward is thought to have extended as long as 75km. Orford, Suffolk, about 15km long, has the appearance of a cuspate foreland and associated spit developed from the north. However, in both cases there is conclusive evidence of earlier shorelines of barrier form to seaward. The sand and pebble banks which still remain offshore of the Orford site may be, in part, a legacy of such structures. Other relic barriers might be found offshore covered by more recent sediments. The time span over which evidence for existence of shingle barriers is available is much shorter than for many of the sand formations recorded elsewhere. This is probably the result of a combination of factors such as the effect of storm waves and comparatively shallow bedrock and the great mobility of coarse clastic material. The small shingle feature at Cemlyn, Anglesey, is alleged to have been built largely as the result of one storm. Even examples restricted to sand barriers show a multiplicity of form, however. Thus there are fundamental differences between the coastlines of eastern Australia and the southeastern seaboard of the United States. In Georgia there is a gently sloping profile less than 1m per km-and hence broad lagoonal salt marshes occur between the low, widely spaced barriers; but in New South Wales the continental shelf has a gradient of as much as 20m per km and the barriers are closer together. Similarly, erosiona! features in the southeastern United States tend to be mainly washover fans and small dunes compared with active or fossil transgressive dunes in the Australian examples. In Georgia, the coastal plain is drained by fairly large rivers

which have maintained their courses during submergence and barrier development, whereas, in contrast, some rivers have been deflected by the Outer Barrier in New South Wales and Queensland. This is probably a response to the low-medium and high wave energies respectively, while the spit development along the southern side of barrier islands in Georgia is associated with tidal currents which assume a more important role there. There is thus an analogy with the apparent and real longshore development of the Orford pebble beach, while if Bird's (1961) view is accepted, the 'sweeping in' noted in the storm wave-pebble environment may be repeated in the Holocene in respect of Australian sand barriers and, indeed, elsewhere. A further difference between Australian and North American barriers is in their relation to the solid geology. Bedrock promontories and former offshore islands of bedrock form the headlands of zeta-curved or arcuate bays in New South Wales and Queensland, but bedrock does not outcrop on the coastal plain of Georgia. The diversity of form and conditions which are found with coastal barriers in temperate climates is repeated with other environments which can, themselves, provide new factors. The picture of barrier type, composition, and formation is, therefore, one of greater complexity than has previously been suggested. Bird, E.C.F., 1961 The coastal barriers of East Gippsland, Australia, Geog. J. 127, 460-8. Hardy, J.R., 1964 Sources of some beach shingles in England, 20 IGC London. Leontev, O.K., 1965 On the cause of the present-day erosion of barrier bars, Coastal Research Notes 12, 5-7. Shepard, P.P., 1952 Revised nomenclature for depositional coastal features, Bull. Am. Assoc. Pet. Geol. 36, 1902-12. Stride, A.H., 1963 Current-swept sea floors near the southern naif of Great Britain, Quart. J. Geol. Soc. London 119,175-99. Zenkovitch, V.P., 1967 In Processes of coastal development, éd. J.A. Steers assisted by C.A.M. King (Edinburgh); originally published in Russian (Moscow, 1962).

Géomorphologie I 13 PO 106

Quaternary geomorphology of the Aconcagua Valley C.N. CAVIEDES University of Wisconsin, USA In the last decade special attention has been given in Chile to geomorphological investigations of the Andean Cordillera and the coastal plains, designed to clarify the evolution of the landscape during the Quaternary. The main problems to be solved have been the character, number, and extent of the high mountain glaciations, the events which have shaped the coastal regions, and the climatic-hydrological conditions that prevailed during that age. The Aconcagua Valley (lat. 32°50's), located in a climatic transition zone between the arid episodic-humid zone of the so-called 'Norte chico' and the temperate winterhumid part of middle Chile, was chosen as a study area. By means of a profile drawn along the valley from the present Andean glaciation limits at 4200m, across the ranges of the Cordillera de la Costa and the hills of the coastal batholith down to the sea, the traces of glacial and periglacial action in the highlands as well as the influence of sea level changes on the coastal morphology can be distinguished clearly. By interpreting the deposits left by the later glaciations, which consist of enormous lateral moraines and lesser terminal-ablation moraines of coarser material, we conclude that they were comparatively weaker and less humid than the earlier ones. These were responsible for shaping the valley down to 1100m and left sizeable terminal moraines, built of glacial till, deposited along transverse rock benches. The fact that the climatic conditions during the glacial advances were mainly humid is testified by the presence of alluvial or mud fans, colluvial deposits, and Akkumulationsglatthánge (smooth-accumulation-slopes; Weischet 1969), all located in the periglacial zone. In the Andean piedmont well-developed terraces indicate the existence of several periods of river discharge of which the older ones were again the most powerful. Considering the features of glacial erosion, morainic accumulations, and the correlative proglacial deposits, three major Quaternary glaciations are represented in the upper Aconcagua Valley. Their best-preserved traces can be found at heights of 2800m

(Portillo), 1600m (Guardia Vieja), and 1300m (Salto del Soldado). The possibility of a fourth glaciation at 1100m is suggested by some characteristic erosional features in the shaping of the valley. Contrary to older opinions (Briiggen 1950), none of the glacial advances passed through the frame of the Cordillera. Volcanic and seismic-catastrophic incidents, very common in the Andean valleys south of the Aconcagua (Borde 1963, 1966) did not occur in the Upper Aconcagua and consequently the morainic systems are undisturbed. In the middle course of the Aconcagua, where the river crosses chains and intermontane basins, probably as a consequence of local depressions and a very strong alluviation of the valley bottom during the last glacial/pluvial stage, former stream terraces were covered by a thick alluvium so that today the flat surface of the low terrace dominates the whole valley floor. Only on the slopes facing the valley do some red weathering crusts and fine colluvial deposits indicate the alternation of humid-cool and temperate-dry periods. Geomorphological features in the coastal region are directly related to the changing sea level during the Quaternary: stream and marine terraces as well as dune fields were the results. Below a vast Tertiary coastal plain (200m above the sea level) five stream terraces at 120m, 80m, 60m, 40m, and a lower terrace coinciding with the principal valley floor, flank the river on both sides and pass into corresponding marine terraces at the coast. No major tectonic deformations have affected the terrace surfaces. Because of the exceptional tectonic tranquillity of this coastal section of middle Chile, all of the marine and stream terraces are clearly related to periods of stable or even rising sea levels within a general regression which began at the transition of the Pliocene to the Earlier Pleistocene (Herm 1969; Paskoff 1970). Morphometric studies on stream and marine pebbles reveal that they were deposited during rises of the sea level (transgression conglomerates). Furthermore, the studies support the idea that the

14 / Geomorphology transport conditions during the earlier rises of sea level were of an almost torrential character, while later on they became less powerful. The existence of repeated retreats was proved also by means of morphometric measurements. Pedological observations of modern soils and paleosoils on terraces and dunes reveal that the optimal conditions for their formation reigned during cool-humid periods, as can be seen from their similarity to middle-latitude brown soils. During the arid temperate periods the drying conditions led to the rubéfaction of the soils. These ideas are supported by mineralogical analyses of sand from coastal dunes which lie upon the marine terraces. The older dunes were probably accumulated under a warmer and drier climate than the younger ones because they are rich in iron-manganese oxides and resistant minerals and their colour is brown-reddish. The younger dunes exhibit few oxides; the colour is light grey; and alteration minerals are very frequent. Consideration of all the geomorphological features described along the Aconcagua Valley indicates the following conclusions. First, during the Quaternary there was an alternation of temperate-dry climatic periods with hydrological conditions of torrential character and cool-humid periods with more constant flow and carrying power. These periods correspond, without doubt, to inter glacial/ interpluvial and glacial/pluvial periods respectively. Second, during glacial/pluvial periods glaciers advanced down the Aconcagua Valley and left evidence of at least three glacial systems. Meanwhile, in the coastal region, because of a drop of the sea level, the river degraded, leaving free space for the sedimentation during the interglacial transgression that followed. Third, the three glacial advances in the High Cordillera (a fourth and older one is questionable) are obviously correlative with the widespread Quaternary glaciations which affected the southern hemisphere. The glacial systems of Portillo, Guardia Vieja, and Salto del Soldado should be considered as belonging to the three last principal glaciations of the Pleistocene. This conclusion is supported by the palynological studies of Heusser (I960, 1966) in south Chile. They also allow the presumption that the glaciations in the middle Chilean Andes occurred 'in phase' with the northern hemisphere's cold periods.

Accordingly, we can presume that sea level rise during the corresponding interglacials was responsible for the formation of the marine and stream terraces. Whilst dating is not too difficult on the Tertiary coastal plain (because of paleontological proof offered by Herm 1969), or on the Holocene lower marine terrace (archeological findings by Montane 1964), the other terraces (120m, 80m, 60m, and 40m) must be assigned by interpolation. They are attributed to interglacial high stands during the Early and Middle Pleistocene. These results in the Aconcagua Valley parallel those obtained by Herm and Paskoff (1967), Herm (1969), Weischet (1969), and Paskoff (1970) in another quite stable region contiguous to middle Chile. It is desirable that these results be proved by absolute dating methods, so that they could be assigned a firm stand within the world-wide chronology of the Quaternary. Universidad Católica de Valparaiso, Chile; Geographisches Institut I of the University of Freiburg, West Germany; Alexander-vonHumboldt Foundation; W. Weischet. Borde, J., 1963 Las incidencias cataclísmicas en la morfología de los Andes de Santiago, Inf. Geográficas 10, 7-25. Borde, J., 1966 Les Andes de Santiago et leur avant-pays: étude de géomorphologie (Bordeaux). Brüggen, J., 1950 Fundamentos de la geología de Chile. Insto. Geog. Mil., Santiago de Chile. Herm, D., and R. Paskoff, 1967 Vorschlag zur Gliederung des marinen Quartârs in Nord-und Mittel-Chile, N. Jb. Geol. Palâont., Mh., v. 10, 577-88. Herm, D., 1969 Marines Pliozàn und Pleistozàn in Nord- und Mittel-Chile unter besonderer Beriicksichtigung der Entwicklung der Mollusken-Faunen. Zitteliana. Abh. d. Bayer. Staatssammlung f. Palâont. u. hist. Geol. v. 2, 1-159. Heusser, C., 1960 Late-Pleistocene environments of the Laguna de San Rafael area, Chile, Geogr. Rev. v. 50, 555-77. Heusser, C., 1966 Late-Pleistocene pollen diagrams from the province of Llanquihue, southern Chile, Proc. Am. Phil. Soc. 110 (4), 269-305. Montane, J., 1964 Fechamiento tentativo

Géomorphologie I 15 de las ocupaciones humanas en dos terrazas a lo largo del litoral chileno, Arqueo* logia de Chile central y áreas vecinas (Santiago de Chile), 109-24. Paskoff, R., 1970 Le Chili semi-aride:

recherches géomorphologiques (Bordeaux). Weischet, W., 1969 Zur Géomorphologie des Glatthang-Reliefs in der ariden Subtropenzonen des Kleinen Nordens von Chile, Z. /. Geomorph. 13, (1 ), 1-21.

P0107 Cryopedimentation: an important type of slope development in cold environment JAROMIR DEMEK Czechoslovak Academy of Sciences Cryopedimentation is the parallel retreat of scarps and the resulting development of gentle erosion piedmont surfaces (pediments) in cold (periglacial) environments. In the literature the development of pediments is usually linked to arid and semiarid climates. A number of authors (e.g., Baulig 1952; Dylik 1957; etc.) pointed out the considerable similarity of morphogenesis in dry-warm and cold climates. In recent years, recognition of the widespread occurrence of pediments in both regions of present cold climate and in those of former Pleistocene periglacial climate has increased. Pediments that develop owing to the effects of a complex of cryogenic processes are called cryopediments. The extensive occurrence of cryopediments in these regions was confirmed by our investigations in the regions of the present continental variant of the cold climate of eastern Siberia and in the regions of the Pleistocene periglacial zone of central Europe. Marginal pediments at the foot of marginal slopes, valley pediments, and summit pediments (cryoplanation terraces) can be distinguished. Cryopediments are very common in the mountainous regions of eastern Siberia. The marginal pediment borders the foot of the Verchojanskij Chrebet Ridge for a distance of many hundreds of kilometres. Characteristic are the valley pediments. The valleys of the water courses in eastern Siberia have often large valley floors on which underfit streams meander. The wide floors are bordered by a few (usually two or three) low river terraces. Above the terrace a gentle slope (usually of an inclination between 1 and 10° ) rises. The foot surface is usually several hundreds of metres wide, in places even 2 to 3km. In the past it was often considered a fluvial erosion terrace or a talus built of slope deposits, but boreholes and

exposures indicate that this surface is in substance an erosion form. The scarp over the foot surface does not bear any evidence of river erosion (meandering of the water course, etc. ). The thickness of the slope deposits on the foot surface is small, usually reaching only about 1 to 2m. The foot surface passes over into steeper valley sides in a break of slope. In places, there is a sharp knick, but usually the transition is very gentle. The steeper slopes are mostly covered with debris and block fields. On foot surfaces and valley floors there is a cryogenic microrelief (patterned grounds, solifluction tongues, etc.). In the entire area continuous permafrost is developed, its thickness often reaching as much as 1500m. The cryopediments are developed in various rocks - granites, gneisses, sandstones, shales, etc. They occur in rejuvenated mountain ranges built of old rocks of the Siberian Platform as well as in young mountain ranges created by the Pacific orogenesis, but they are typical of the regions in which lateral widening of the valleys prevailed over their incision toward the end of the Pleistocene and in the Holocene. The cryopediments develop as a result of a complex of cryogenic processes. On the scarps, mainly frost weathering and frost creep manifest themselves. In rocks with vertical fissures (granites, limestones, dolomites) the separating of rock columns, their successive tilting due to the development of frost wedges in the fissures, and final toppling over and disintegration take place. This special type of slope development plays a very important role in parallel slope retreat. The feet of the scarps are also undermined by nivation. In places with a thicker active layer at the foot, groundwater appears. The soaking of the slope foot leads to greater activity by cryogenic processes.

16 / Geomorphology Another complex of processes operates on the lower surfaces. The main process is frost sorting and frost heaving. In a smaller way even solifluction comes into play, but its effects are reduced to a short period in spring during the snow melt, because of the dryness of the climate. Numerous shallow delldepressions are also of importance in the modelling of the foot surface. These are lines along which the surface water is concentrated and in which the cryogenic processes are more intensive due to greater soil humidity. In places even traces of thermokarst processes can be found in them. The material slides slowly down the gentle slopes into the dells and is transported by frost creep, solifluction, and running water towards the water course. In some parts of Siberia (such as for instance in the foothills of the Stanovoj Chrebet Ridge) the merging of adjacent cryopediments, the origin of pediment passes, and the development of inselberg landscapes have already taken place. The valley pediments are either directly linked up to the river flood-plain or to the low terraces of young- or middle-Pleistocene age. This points to their relative youth. The geomorphological investigation of cryopediments is of considerable practical importance in the study of the occurrence of precious minerals (gold, tin, etc.). A special problem relates to the Pliocene and older pediments in the Siberian region developing continuously in the conditions of the presentday cold climate. These are mainly the socalled pediments of the Siberian type, according to Timofeev ( 1962). Cryopediments are common even in the zone of the Pleistocene periglacial climate in central Europe. In Czechoslovakia they are also mostly linked to low river terraces of young- or middle-Pleistocene age. They occur in various rocks (sands, clays, sandstones, shales, etc.). In places, they display small declivities (0.5 to 3 ° ). Their width amounts usually to several hundreds of metres up to 2km and their length to as much as 15km. In permeable rocks (sands, sandstones) their development was possible only in the presence of permafrost which formed the impermeable layer; in present climatic conditions they do not go on developing. Even cryopediments buried below loess or slope deposits were found.

The summit pediments are given various names, such as cryoplanation terraces, altiplanation terraces, goletz terraces, etc. They develop on summits and in upper slope sections. They consist of the frost-riven cliff or frost-riven scarp and the terrace flat. They develop in various rocks in which a steeper slope can come into origin. They are due to the parallel retreat of the frost-riven cliff or scarp owing to nivation, frost weathering, and frost creep. At the foot of the scarp a rocky flat develops covered with a thin layer of loose material of an inclination between 1 and 12°. The material developing on the frost-riven cliff or scarp is transported on the flat only by frost heaving, frost creep, and running water, and on the margins also by solifluction. Cryoplanation terraces reach great dimensions in the Siberian mountain ranges. Terraces were found as much as 2-3 km wide and more than 10km long. In the mountain ranges of central Europe their dimensions range between several tens and several hundreds of metres. They occur in groups one above the other. Most often 4 to 5 terraces can be found on the slope. But from Cukotka, slopes with more than 30 terraces one over the other were described. The terraces are rock forms disintegrating the older topographical surface. In the mature stage of development a summit cryoplanation flat with tors develops. Through merging of summit flats a levelled surface called a cryoplain comes into existence which can be compared with the pediplain of dry and warm regions. The observations mentioned point to the considerable importance of cryopedimentation in the development of the relief of regions of cold climate. Baulig, H., 1952 Surfaces d'aplanissement, Annales de Géographie 61, 161-83, 24562. Czudek, T., and J. Demek, 1970 Pleistocene cryopedimentation in Czechoslovakia, Acta Geographica Lodziensia 24, 101-8. Demek, J., 1969 Cryoplanation terraces, their geographical distribution, genesis and development, Rozpravy GSAV, rada MPV, 79141, 1-80. Dylik, J., 1957 Tentative comparison of plaration surfaces occurring under warm and under cold semi-arid climatic condi-

Géomorphologie I 17 tions, Biuletyn Peryglacjalny v. 5,175-86. Timofeev, D.A., 1962 K problème proischoMënija formy recnych dolin (na pri-

mere recnych dolin v Juenoj Jakutii). Izvestija AN SSSR, serija geografiâëskaja, 1962/3/, 82-9.

P0108 Inferring process from form: the asymmetry of glaciated mountains IAN s. EVANS University of Durham, UK A major aim in geomorphology is to relate process to form in the context either of present denudational systems, or of land form development through a demonstrable chronology of process. The difficulty of studying processes of glacial erosion directly makes it particularly important to obtain relevant information through the analysis of form. The present study is an attempt to extract the maximum amount of information from readily available photogrammetric contour maps, in connection with the development of asymmetry between opposing slopes in glaciated mountains. Slope aspect controls glacial hydrology in two principal ways. Firstly, accumulation of snow is greater on sheltered slopes, especially if they are to the lee of exposed areas such as summit plateaux which supply winddrifted snow. Secondly, ablation of snow, firn, or ice is greater on sunny slopes, and less on shaded slopes. Because of time-lag effects the afternoon sun causes stronger ablation than the morning sun, and in the Northern Hemisphere the azimuth of minimum ablation is shifted somewhat east of north. Wind drifting is more effective where summit slopes are gentle and snow falls dry during storms; shade, where topography is rugged and skies are clear. The net result is that in the west-wind belt of the Northern Hemisphere, glaciers form at lower altitudes on NE-facing slopes. Mountains which are only just high enough to nourish glaciers are glaciated asymmetrically, with NE-facing glaciers producing NE-facing cirques. Later a lowering of the snowline may lead to glaciation of all slopes, but by then drainage divides have been pushed southwest, and the scope for cirque development on swfacing slopes has been reduced. The tendencies for valleys tributary to east-west valleys to enter from the south rather than the north was noted by Hanson (1924) in the Kitsumkalum area (128°w 54°N), near Terrace, BC, and by Tuck (1935)

in the Kenai Peninsula and Talkeetna Mountains (149°w, 61°N), Alaska. The valleys discussed by Hanson are 2 to 3km long; Tuck suggested that asymmetric glaciation had pushed divides south by around 1km. The mechanism for development of these valleys would be headward cirque erosion during a series of glaciations. Hanson's (1924) hypothesis that north-facing glaciers 'have cut short northerly-trending valleys on the north side of the higher mountains' is tested here for several ranges in the southern Coast Mountains of British Columbia. Since orientation of valleys due to structural lineation would produce two modes of valley direction, 180° apart, which would cancel each other and have no resultant, it cannot alone account for the presence of significant unimodal asymmetry. Two credible hypotheses remain as alternatives to Hanson's; regional tilting, and azimuthal variation in valley density as related to fluvial hydrology. The analysis of valley asymmetry must attempt to discriminate between the three hypotheses. Is this possible on the evidence of form alone? Asymmetric glaciation is limited to a narrow altitudinal band between levels of no glaciation and levels of glaciation on all slopes. Regional tilting, on the other hand, affects valley direction at all altitudes. Hence altitude-specific analysis of asymmetry may permit discrimination between these hypotheses, and if Hanson's is accepted, the analysis will provide information concerning the altitude of morphogenetically significant snowlines. Variation in valley density is likely to affect only total valley length, whereas cirque recession would also affect the mean length of valley segments within a certain range of altitude. A program was written to analyse, for each altitudinal class, the azimuthal distributions of both total map lengths of valleys, and mean length of valley segments per 45° class.

18 / Geomorphology TABLE 1

Range

Longitude w

Latitude

N

Altitude of greatest asymmetry

Zone of asymmetry (m)

Direction of maximum mean length

Bendor Rex Tyax-Gun Tatlow Custer

122° 31' 122° 22' 122° 52' 123° 52' 121° 21'

50° 45' 50° 51' 50° 59' 51° 20' 49° 05'

1740 1920 2070 & 1650 1680 1340

2070-1620 2290-1800 2130-1980 2190-1680 1520-1010

030 038 336 023 015

TABLE 2

Area

Sample size, N

71 Tatlow Range 38 Custer Range Bridge River area ((a)-(9) total, plus 3 isolated cirques) 532 (a) Bendor Range 178 (b) Mission Range & Nosebag Mountain 22 42 (c) Pex Range 84 (d) Shulaps Range 45 (e) Tyax-Gun block 34 (/) Relay block 124 (g) Warner block

Vector mean, degrees

95% confidence limits, degrees

009 001

±16 ±20

009 012 003 017 008 018 358 005

The complete analysis was based on a single data set, generated semi-automatically on a d-mac pencil follower, from British Columbia interim 1/31,680 maps with a 30.48m contour interval. For a whole mountain range, the points where streams crossed contours were digitized, working downwards and labelling each stream with the altitude of the highest contour crossed. Hence the program analysed the generalized stream traces (valleys) in terms of intercontour segments of known altitude. Streams were used because of the difficulty of defining Valleys' objectively: many of those shown are simply valley-side torrents, but these have correspondingly short segment lengths. For one range (Custer), valleys as defined subjectively were also digitized: the results were fairly similar, but resultant vectors were stronger. It is necessary to analyse ranges circumscribed by the lowest available passes, so as not to violate the null hypotheses that streams can flow equally in all directions. Strong N-NE asymmetry was present in four ranges, but the (lower) Tyax-Gun group showed only weak asymmetry (Table 1).

Resultant vector Strength, %

Standard deviation, degrees

43 23

60 62

62 59

±5 ±9

339 97

64 55

56 64

±17 ±15 + 1.1 ±14 ±19 + 10

18 30 57 33 23 80

81 72 67 73 66 65

38 47 51 48 56 56

Length, cirques

Since azimuthal distributions of mean length at both higher and lower altitudes are more symmetrical, it is suggested that these results support Hanson's hypothesis for the areas analysed. Asymmetry becomes marked at cirque floor altitudes and continues some way below, suggesting the effect of cirque development in earlier glaciations. The valley heads concerned do indeed have a heavily glaciated appearance. The asymmetry observed in these areas of southwestern BC is explicable on the basis of solar radiation; distributions of cirque aspect are also strongly to the north-northeast (Table 2), The circular statistics were defined by Curray (1956). Analysis of cirque aspect is possible for many other areas for which authors have published data, and despite possible differences in definition, interesting comparisons are possible. Cirques in Britain, Sweden, the Vosges, and Schwarzwald, have aspects generally too far to the east to be accounted for purely by differential radiation; westerly winds were probably required also. Vector strengths (azimuthal concentrations) are less for these distributions than for those

Géomorphologie I 19 with north-northeast resultants, and greatest for areas of most marginal glaciation. These results suggest that azimuthal distributions of cirque and valley aspect contain valuable information concerning the nature of morphogenetically significant glaciation. Precise analysis is required if the roles of shade and of wind are to be discriminated.

dimensional orientation data, /. Geol. 64 (2), 117-31. Hanson, G., 1924 Reconnaissance between Skeena R. and Stewart, BC, Can. Geol. Survey Summary Report for 1923, Part A, 29-45. Tuck, R., 1935 Asymmetrical topography in high latitudes resulting from alpine glacial erosion, /. Geol. 43, 530-8.

Curray, J.R., 1956 The analysis of twoPÔ109 Objectives and methods in the study of glacier depositional landforms A. FALCONER and EIJU YATSU University of Guelph, Canada In recent years geomorphology has been turning its attention towards a, more detailed examination of the processes and materials involved in the creation of a landscape. It seems that even more detailed attention must be paid to the properties of material if we are to fully understand the intricacies of landscape development. To this end recent studies at the University of Guelph have been oriented towards an examination of the less studied properties of such materials. In the Guelph area the dominant surficial deposit is the Wentworth till (Karrow 1968 ). It is strange that in a study such as geomorphology so little attention has been paid to the stability of this material and the properties which permit it to form and maintain the slopes which constitute the landscape. Equally strange is the lack of attention given by geomorphologists to the processes which acted to create the till. Mechanisms by which till is produced are surprisingly conjectural and, whilst abrasion appears to be a major contributing factor, it is by no means clearly established that it is the dominant factor. It is therefore the concern of the present study to evaluate the contribution of processes and materials to the landforms existing around Guelph, Ontario, Canada. An experiment has been designed to evaluate the effect of abrasion on the shape and size of till stones. However, within an ice mass, the rocks which are undergoing abrasion also are subjected to the stresses of low temperature conditions. A further experiment has been designed to evaluate the behaviour of similar rock-types under these conditions. Together these experiments produce indica-

tions about the processes which have acted to produce the Wentworth till from which much of the present landscape is formed. Further investigations are in progress to evaluate the stability of the material. Shear strength of the glacial deposits of the Guelph area has, in the past, been measured by various engineers for specific sites. These values, however, are related to the individual site investigations and have no clear relevance to any particular strata. Yatsu (1971 ) has also demonstrated that the clay fraction of deposits must be carefully investigated to determine the stability inherent in its clay mineral composition. A total understanding of the landscape of the Guelph area requires a vast amount of research. The components of this undertaken to date have produced the following results. Work by Stratford (1971) indicates that rock disintegration under freeze-thaw conditions may be a function of the duration of the low temperature conditions. Consideration of the results of other investigators (Wiman 1963; Potts 1970) indicates that a linear relationship appears to exist between percentage weight loss of specimens (D) and the duration of the experiment. This relationship may be stated as D = 0.0072f + 2.65 which is significant at the .005 level for the experimental data of Wiman (1963), Potts (1970), and Stratford (1971) (where D = percentage loss in weight, / = duration of experiment in hours). There can be little doubt that low temperature conditions exist within a glacier. That these conditions were of many decades in duration is similarly undisputed. Thus we may expect that the relative abundance and

20 / Geamorphology size of the rock debris is in part a function of the rocks' resistance to frost weathering processes acting through an extended period of time. The till stones being studied indicate the following relationships between dolomite pebbles and distance from their probable source area (Bardecki, in preparation). All these relationships are linear and the correlation coefficient between each pair of variables is significant at the 0.1 level or better. The independent variable (X) in each case is the distance along the direction of ice movement from the source area in kilometres: Dolomite content of gravels (Y^) may be expressed as: Y1 = 68.403 + 0.598JT % Dolomite pebbles classed as tabular (F2) : Y2 = 34.854 - 0.433Z % Dolomite pebbles classed as pentagonal (r3) : F3 - 30.395 + 0.433Z Average roundness of dolomite pebbles (r4) : F4 = 0.0758 + 0.0017Z The linear relationship between distance and roundness, distance and ovoid, and distance and wedge shape were not significant. Thus with distance from source the tabular pebbles show a continuing decrease in occurrence and pentagonal pebbles an increase, thereby supporting the contention of Wentworth (1936) that the pentagonal shape is the preferred shape of glacial pebbles and opposing the contention of Holmes (1960) that there is a tendency for an ovoid shape to develop. These results are based on samples taken within 35km of the source area. The increase in content of gravel-sized dolomite with distance from the source area indicates the importance of the need to study the large material in till. Presumably the content of dolomite gravel stabilizes at some distance from the source area when fragmenting boulders no longer contribute to the gravel fraction. The till itself has been investigated for stability (Barr 1971). The Wentworth till of the Guelph area has a residual friction angle of approximately 30°, when measured in a small shear (6cm X 6cm) box using that portion of the till specimen which passed a 2mm sieve. However, a large shear box (30.48cm X 30.48cm) which tested the portion of the till specimen passing a 38.1mm sieve, gave values of approximately 40°. No systematic regional trend of shear strength has been revealed by this investigation to

date. Clearly the results of the study of till stones and shear strength underline the need for more detailed study of the nature of the large particles in till. Investigations of the areal distribution reveal that the clay-size content of the till increases along the direction of ice advance whilst sand content decreases, a phenomenon observed by Falconer (1970) and Chorley etal. (1966) in similar studies. Thus by changing the methods in the study of landform it appears that the objectives of geomorphology may be more easily achieved. Geomorphological studies 'are required largely to furnish a working model for the outer part, at least, of ... (the) ... earth* (King 1968). By emphasizing the studies of the properties of Wentworth till it seems that linear models offer some valuable data to this end. The authors suggest that the objective reporting of geotechnical properties of landforming materials offer a more fruitful means of continuing geomorphic studies than the descriptive methods commonly employed. Amassing such geotechnical data in a standardized form could lead to rapid realization of the geomorphologists' goal, namely a working model for the outer part of the earth. M.J. Bardecki; M.V. Barr; J. Bones; BJ. Stratford; L. Myers; National Research Council; National Advisory Committee on Geographical Research. Bardecki, M.J. The evolution of the shape of till stones. MSc thesis, U. Guelph. Barr, M.V., 1971 Preliminary investigation of the shear strength of Wentworth till. MSc thesis, U. Guelph. Chorley, R.J., D.R. Stoddard, P. Haggett, and H.O. Slaymaker, 1966 Regional and local components in the areal distribution of surface sand faciès in the Breckland, eastern England, /. Sed. Pet. 36, 209-20. Falconer, A., 1970 A study of the superficial deposits in Upper Weardale. PhD thesis, U. Durham. Holmes, C.D., 1960 Evolution of till-stone shapes, central New York, Geol. Soc. Am. Bull. 71, 1645-60. Karrow, P.P., 1968 Pleistocene geology of the Guelph area. Ont. Dept. Mines, Geol. Rept. 61. King, L., 1968 Geomorphology, in R.W.

Géomorphologie / 2 l Fairbridge, éd., The Encyclopedia of Geomorphology (New York). Potts, A.S., 1970 Frost action in rocks: some experimental data, Trans. Inst. Br. Geog., 109-24. Stratford, B.J., 1971 A study of the frostweathering of rocks. MSc thesis, U. Guelph. Wentworth, W.C., 1936 An analysis of the

shapes of glacial cobbles, J. Sed. Pet. 6, 85-96. Wiman, S., 1963 A preliminary study of experimental frost weathering, Geogr. Annlr. 45, 113-21. Yatsu, E., 1971 Landform material science: rock control in geomorphology, in Research Methods in Geomorphology (Toronto).

P0110 Dating cave calcite by the uranium disequilibrium method: some preliminary results from Crowsnest Pass, Alberta D.C. FORD, P.L. THOMPSON, and H.p. SCHWARCZ McMaster University, Canada Uranium is present in trace amounts in many rocks. When these are chemically weathered it may be leached as a solute. Should host waters contain other, more common, solutes in such abundance that there is supersaturation and precipitation occurs, some of the uranium will be co-precipitated. Trapped within the new deposit it is a potential clock, recording the time elapsed since precipitation by measurable radioactive decay. The uranium radioactive decay scheme yields four possible methods of dating a host precipitate: 1. the isotope ratio, U-234/U-238. Potentially this is effective within the timespan, 50,000-1,500,000 years BP. 2. the isotope ratio, Th-230/U-234. The dating timespan is 2000-350,000 years BP. 3. the isotope ratio, Th-230/Pa-231. The dating timespan is 2000-200,000 years BP. 4. the isotope ratio, Ra-226/U-234. The dating timespan is 300-7000 years BP. Details of the methods may be found in Rosholt et al. (1961), Thurber (1962), Kaufman and Broecker (1965), and Thurber et al. (1965). To date, their application has been principally to marine or salt lake precipitates. In principle, the first two methods and the fourth are equally applicable to freshwater precipitates. In 1967 the authors commenced an investigation of the stalactite, stalagmite, and flowstone calcite deposits of limestone caves. Two important preliminary points have been established: (a) in the oceans, the ratio U-234/U-238 is universally 1.15 [1.00] (Thurber 1962; Koide & Goldberg 1965). In fresh waters, we have found that it may vary considerably both in

space and in time. This severely limits use of the U-234/U-238 method, with its exciting timespan of 1,500,000 years. (b) the Th-230/U-234 method presumes that all Th-230 measured in a sample precipitate is authigenic (Barnes et al. 1956; Fornaca-Rinaldi 1968; Duplessy et al. 1970). We have found that this is so in clean cave calcite deposits (i.e., those without large amounts of detrital material such as clay). Using the Th-230/U-234 method, dates have been obtained for cave calcite deposits from Britain, West Virginia (USA), Mexico, and the Rocky Mountains of Canada. The remainder of this paper focuses upon an investigation in the southern Rockies which has yielded useful information about the age of the alpine relief there. Like all other ranges that compose the Canadian Rockies, the High Rock Range was glaciated during the last glacial (Wisconsin). It is deeply dissected by cirques. A valley glacier 700m deep flowed through Crowsnest Pass, which is the principal breach in the range. The floor of the pass is at an altitude of 1335m nisi. The mountain summits about it rise to 2550-273Om msl. Thus, local relief in these typically alpine mountains is 12001400m. At the surface, all erosional landforms and deposits may be attributed to the last glacial or to post-glacial action, but in the mountains there are many fragments of phreatic (subwatertable) caves which are now fossil. These caves are scattered throughout the local range of altitude. They serve to drain paleogroundwater streams to the pass (Ford 1971 ), which is evidently an ancient feature. Many of the caves contain stalactites,

22 / Geomorphology TABLE 1

Th-230age (yrsfip)

Site

Material

Mt. Coulthard Cave alt. 2278m Middle Cave of Sentry Mtn. alt. 1810m

broken stalagmite (a) core (b) outers sheath flowstone ledge, 10cm thick (a) base (b) centre (c) top flowstone ledge 6cm thick - base

Eagle Cave alt. 1440m

stalagmites, and flowstones in abundance. These deposits could not have begun to form until the caves were drained. The caves could not have drained until the floor of the pass was entrenched below their level. Therefore, the age of the oldest calcite deposit in a cave is the minimum age for the floor of the pass at the altitude of that cave. Following are dates from the three caves that have been investigated at the time of writing. The following conservative conclusions may be drawn. (a) The floor of Crowsnest Pass was below 1810m (modern altitude) at least 275,000 years ago; 63 per cent of the modern relief at the pass existed at that time. (b) The floor was below 1440m (modern altitude) at least 200,000 years ago; 90 per cent of the modern relief at the pass existed at that time. Therefore, although the modern mountains and valleys of the region display only the effects of last glacial or later erosion, etc., upon their surfaces, it is evident that these effects are retouches of much older features. The pass is a breach 1200m deep today. During the past 200,000 years only the final 100m (at most) has been added to that depth. (This yields a mean rate of < 50cm per 1000 years for the entrenchment of the pass since 200,000 years BP.) By reasonable extrapolation, some of the modern relief at the Pass and in adjacent mountain ranges existed at least a million years ago and the broad configuration of the modern ranges was determined by that time. These findings are preliminary. However, they give a much longer view of the erosional evolution of parts of the Rocky Mountains than has been available hitherto (from C-14 dating). They in-

296,000 + 33,000 235,000± 19,000 ?73,000 + 37,000 160,000+13,000 99,000 ± 6,000 198,000± 13,000

dicate the power of the Th-230/U-234 dating method when it is applied to calcite deposits in ancient caverns. In further work at Crowsnest Pass it is hoped to increase the number and span of calcite dates and add thermal information by measuring the O-16/ O-18 ratio in the deposits. National Research Council of Canada; McMaster University. Barnes, J.W., E.J. Lang, and H.A. Potratz, 1956 Ratio of ionium to uranium in coral limestones, Science 124, 175. Duplessey, J.D., J. Labeyrie, C. Lalou, and H.V. Nguyen, 1970 Continental climatic variations between 130,000 and 90,000 years BP, Nature 226, 631-3. Ford, D.C., 1971 Characteristics of limestone solution in the southern Rocky Mountains and Selkirk Mountains, Alberta & British Columbia, Can. J. Earth Sci. v. 8(6), 585-609. Fornaca-Rinaldi, G., 1968 230Th/234Th dating of cave concretions, Earth and Planetary Sci. Letters 5, 120-2. Kaufman, A., and W.S. Broecker, 1965 Comparison of Th230 and C14 ages for carbonate materials from lakes Lahontan and Bonneville, /. Geophys. Res. 70, 403954. Koide, M., and B.C. Goldberg, 1965 U234/U-238 ratios in seawater, Progress in Oceanography 3, 173-7. Rosholt, J.N., C. Emiliani, J. Geiss, F.F. Koczy, and P.J. Wangersky, 1961 Absolute dating of deep-sea cores by the Pa-231/Th-230 method, /. Geol. 69, 16283. Thurber, D.L., 1962 Anomalous U-23 4 /

Géomorphologie / 23 U-238 in nature, /. Geophys. Res. 67, 4518. Thurber, D.L., W.S. Broecker, R.L. Blan-

chard, and H.A. Potratz, 1965 Uranium séries ages of Pacific atoll coral, Science 149, 55-8.

P0111 Morphogenetic interpretation of a mountain section of the Papaloapan Basin, Mexico H. GRANIEL-GRANIEL Escuela Normal Superior, Mexico The area to be discussed covers parts of the states of Puebla, Veracruz, and Oaxaca. Its geographical position is latitude 17°-19°N, longitude 95°-97°w. It is a very important basin with an area of 47,517 sq km which discharges into the Gulf of Mexico about 50,000,000 cubic m of water per year. For hydrological purposes it may be divided into two parts, the high and the low basin. The high basin contains the principal geomorphic unit of the whole basin, called by different geographical names, such as Complejo Poblano Oaxaqueño (Alcorte Guerrero), Sierra Madre de Oaxaca (Atlas de la Comisión del Papaloapan and J.L. Tamayo), Sierra Madre Oriental (Petróleos Mexicanos and Sapper). These and other names are still a subject of controversy. The principal purpose of the geomorphic study of this basin is to establish the physical continuity of these mountains with the geomorphic unit of the Sierra Madre Oriental. In my opinion, the apparent discontinuity between them is due to enormous accumulations of volcanic material which has been burying the Mesozoic folded sedimentary rocks since the middle Tertiary. I believe that for a geographical analysis it is necessary not only to prove the lithologie and tectonic continuity of the Sierra Madre Oriental and of the mountains of the region, but also that at present both ranges are undergoing the same erosion processes. The stratigraphie section of the range in the Papaloapan Basin is similar to that which has been established in the northern ranges; both have the same Mesozoic rocks. At the end of the Cretaceous, there was a geosyncline around the area studied, filled with marine sediments lying over metamorphic rocks (gneiss and schist), all of them having a lateral support on a massif. The crystalline base was inferred by physical prospecting, but recently has been checked by geologic studies which have determined a pre-Paleozoic age. Consequently it is possible that this

rigid continental block acted, and is still acting, in the orogenesis of this part of the country. In early Eocene, after the uplift of this rigid massive block, a slow slide of the enormous thickness of sedimentary rocks occurred, building a primary relief with anticlines and synclines, trending NNW to SSE. The trend proves that the compression forces acted from ssw to NNE. This orogenesis reached its climax in the middle of the Eocene, not merely bending but overturning the Mesozoic rocks until several thrust faults (Cerro Rabón, Orizaba, and Santa Rosa faults), developed. At the end of this period a relative decrease of compression gave a release to the folded rocks and as a physical compensation, a general normal faulting process took place which has set the existing drainage pattern since then. At that time, due to the general faulting process, several sinking faults which run parallel to the Tehuacan normal fault occurred. This stepped faulting formed a graben, capturing the drainage within a closed basin, and filling the bottom with limestone breccia and later with red conglomerate carried down from the land above by ancient streams. The last compression forces of this orogenic revolution bent the eastern part of this range and influenced other overturning and thrust faults; one of these put Cretaceous sediments over early Eocene sediments. By that time a full marine regression began to affect the area. In the Miocene period, through a fault at the 19th parallel, there were several emissions of lava, forming the base of the Pleistocene volcanoes. These events were the beginning of the apparent discontinuity of the Sierra Madre Oriental. Due to a regional uplift, the sea receded and also all the erosional processes were accelerated. The ancient Papaloapan river increased its erosion, guided by young normal

24 / Geomorphology faults on the Santo Domingo canyon until this river gained by a classic piracy all the waters of the inter-mountain basin. The time of this capture probably was in the Pleistocene period. I think this is so because the graben contains modern sediments and because the Santo Domingo canyon is in a very young stage, lacking a flood plain throughout its length. This suggests that the uplift is still in progress. This geomorphic unit is of outstanding geographical importance for the whole Papaloapan Basin, as it acts as a barrier to physical, biological, and human phenomena. Because of its altitude and parallelism to the shoreline, it receives on its eastern slopes all the humidity of the prevailing winds, producing in its western slopes a rainfall shadow. Consequently, there is a difference in precipitation of up to 4500mm between the two regions. This difference accounts for the dry condition of its western slopes and of the graben behind the ridges. This region is dependent on the resources of the range, and its economic exploitation is restricted to

sugar cane culture, along the flood plains of the main rivers. As seen, these ranges of the Papaloapan Basin are not only sharply delimited but at the same time are subordinated to two regions: the coastal plain which is enriched by their run-off and the so-called Oaxaca Depression which separates them from the old southern mountains of Oaxaca. I consider that there is no reason to continue mixing up these two main geomorphic units of the area, which because of their own characteristics should also have different names. I, therefore, suggest that the Western uplands should be called 'Poblano-Oaxaqueno Massif/ The ranges limited to the east by the coastal plain and to the west by the stepped graben of Tehuacan-Tomellin are properly the 'Sierra Madre Oriental.* Sapper, Karl, 1937 Handbuch der Regîonalen Géologie: Mittelamerika (Heidelberg). Schuchert, Charles, 1935 Historical Geology of the Antillean Caribbean Region (New York).

P0112 Postglacial rock wall recession in the Ogilvie and Wernecke mountains, central Yukon Territory JAMES T. GRAY Lourentian University, Canada Several local rates of erosion by scarp recession have been calculated in recent years (Freise, Poser, in Rapp 1960; King 1956; Rapp 1960). The present study of the rate of postglacial scarp retreat in the central Yukon Territory was undertaken partly to augment the limited evidence of erosion rates in periglacial regions and partly to test the superficial observation that the rate of erosion of igneous intrusives of the area is slower than that of metasedimentaries there. The mountain slopes considered in this paper were divided into a zone of erosion, viz. the bedrock wall, and a zone of deposition, viz. the talus cone. Measurement of the talus cone volumes and the rock wall areas from which the talus debris was eroded enable the mean amount of rock wall recession to be determined. The talus accumulations and their rock wall source areas were usually very easy to

delimit, and they were also small enough to be amenable to measurement by means of field traverses. The talus cones are 15-120m in height; the bedrock walls are 50-500m in height. Erosion of the rock wall is considered to include both weathering of the bedrock and transportation of the weathered products to the major zone of accumulation within the postglacial period, i.e. the talus slope below the rock wall. Thus temporary accumulations of debris in the gullies on the rock walls are not included in calculations of talus volumes. Morphological evidence suggests that preexisting debris was almost completely removed by the passage of ice during the last glaciation and also that there has been little removal of postglacially derived debris beyond the talus cones. Therefore, the present talus volumes represent fairly accurately the products of bedrock wall erosion during the postglacial period.

Géomorphologie I 25 TABLE 1

Amount of recession

Mean annual rate of recession

Cone

Mean (m)

Max. (m)

Min. (m)

Mean mm/1000yr

Max. mm/lOOOyr

Min. mm/lOOOyr

6801 6804 6808 6813 6814 6816 6821 6824 6825 6830

0.636 0.863 0.510 2.060 1.025 0.271 0.085 1.245 0.345 0.358

1.170 1.865 1.161 3.880 1.925 0.620 0.160 2.340 0.650 0.674

0.388 0.450 0.270 1.235 0.615 0.144 0.051 0.748 0.207 0.215

053 072 043 173 086 023 007 104 029 030

117 187 116 388 193 062 016 234 065 067

028 032 019 088 044 010 004 054 015 015

The equations used in the calculations of amounts and rates of rock wall recession are as follows: Amount of recession (in linear measure) = Net volume of debris in talus cone/Source area of rock wall; Rate of recession = Amount of recession/ Postglacial time interval. Calculations of talus volumes and rock wall source areas were made for ten sites by a combination of ground measurements and photogrammetric measurements. The postglacial time interval was estimated at between 10,000 and 14,000 years BP. Maximum and minimum possible empirical errors were accumulated through all stages of the calculations (Gray 1971). The final output consisted of maximum, mean, and minimum values for the amounts and rates of rock wall recession (Table 1 ). The median value for the mean thickness of rock removed from the rock walls during the postglacial period of 10,000 to 14,000 years lies between 0.636m and 0.510m. The range of values is from a possible maximum of 3.88m for the most rapidly eroded rock wall area to a possible minimum of 0.0651m for the least rapidly eroded rock wall area. The median rate of erosion has been 4353mm/1000 years. The largest possible range of values for the ten rock walls is from 388mm/1000 years to 5mm/1000 years. In the igneous intrusives zone in the field area, the low ratio of talus height to rock wall height and the thin talus mantle suggested initially that the average rate of postglacial erosion on syenites and monzonites was much less than in the metasediments which comprise most of the field area. Quantitative

studies tend to support this conclusion, although the fact that only two erosion rates were available from the igneous intrusives prevents it from being couched in mathematical probability terms. For two igneous intrusive rock walls a mean erosion rate of 19mm/1000 years was established. The corresponding mean rate for eight metasedimentary rock walls has been 73mm/1000 years. The erosion rates derived from the Ogilvie and Wernecke mountains are the lowest of four detailed analyses in mountain areas of Europe and northwest North America. They are also very much lower than the rates established by King (1956) and Freise (in Rapp 1960) for scarp recession in tropical regions, and emphatically indicate that rock wall recession in periglacial zones is relatively slow. This is believed to be the largest and most detailed evidence on rock wall recession ever assembled from a mountain region. In view of the fact that the mean rates of erosion agree to within an order of magnitude with those calculated by Rapp (1960) for rock walls in limestones, cherts, and sandstones in Spitsbergen, it can be firmly concluded that postglacial rock wall retreat in metasedimentary and in some sedimentary lithologies in the periglacial zone has not been nearly so rapid as used to be commonly believed. Geological Survey of Canada, O.L. Hughes, J.G. Fyles, E.G. Craig, D. Reimer, A. Gray, T.L. Crow, J. Tibert. Gray, J.T., 1971 Processes and rates of development of talus slopes and protalus

26 / Geomorphology rock glaciers in the Ogilvie and Wernecke mountains, Yukon Territory. PhD thesis, McGill U. King, L.C., 1956 Research on slopes in South Africa. Premier Rapport, Commis-

sion pour l'étude des versants, IGU Amsterdam. Rapp, A., 1960 Talus slopes and mountain walls at Tempelf jorden, Spitsbergen, Norsk Geol. Tiddsk, no 119, 1-96.

P0113

Hillside slope forms and their evolution MASAMI ICHIKAWA Tokyo Kyoiku University of Education, Japan

There are many previous works on the evolution of slopes (e.g. IGU Commission on Evolution of Slopes, 1955a, b, 1956, 1960, 1964). Problems of slope evolution are complicated, and there are some controversies concerning processes and types of slope evolution. For example, the uniformitarian theory - backwearing - on retreat of slopes (Penck 1924; Bryan 1940; King 1953; Tanner 1956) is in striking contrast to the declining declivity theory - downwearing.— on slope retreat (Davis 1909, 1932, 1938; Baulig 1940; Horton 1945; Strahler 1950; Derrau 1956; Yoshikawa 1956; Baulig 1957). There is also an opinion that slopes develop by backwearing until a certain stage, and subsequently by downwearing (Lawson 1932; Wood 1942; Mino 1942; Berry and Ruxton 1960). However, there are other views on slope evolution which suggest that general rules for slope evolution do not exist; it depends on the process which prevails in slope formation (Schumm 1956; Tricart 1957; Dumano wski 19 64 ). The purpose of this study is to clarify the processes and types of slope evolution on the basis of field observations and field experiments during periods of heavy rainfall on the slopes of isolated hills in the Atsumi peninsula of Japan. The hills consist of Paleozoic formations and are inliers surrounded by Plio-Pleistocene formations upon which terraces have been formed. Almost all of the longitudinal profiles of hillsides show concavo-convex slopes with round tops and gentle concave foot slopes. The most common slope form in the study area can be divided into five parts from the viewpoint of the changing state of the longitudinal profile from top to foot. They are: convex slope at the top (A-segment), convex inflection slope (B-segment), hillside steep

slope (c-segment), concave inflection slope (o-segment), and concave gentle slope (Esegment). By conducting field experiments, the writer established how erosion, transportation, and sedimentation took place on each segment of the slope during periods of heavy rainfall. Three examples were selected from 63 isolated hills in the Atsumi peninsula, and on each one slope was chosen as the experimental slope. These experimental slopes can be classified into two types; Slope I was a slope having a very steep side slope (c-segment) with small convexity at the top and also small concavity at the foot; Slopes li and in had gentle side slopes with large convexity at the top and also large concavity at the foot. Gradient and length of slopes, and the grain-size distribution of surface debris on the experimental slopes have been measured. Two types of experimental apparatus were set up on each slope segment mentioned above. A-type measured the amount of debris losses from 0.5m2 plots, while B-type measured the quantity of debris flow from the upper part of the slopes to experimental stations during rainfall. The relationships between maximum rainfall (mm) per 10 minutes and average debris losses (gram) per 10 minutes from A-type apparatus during the experimental period are as follows: the debris loss is larger from csegment of Slope i than from other segments of Slopes i, u, and m. The debris loss from A-segment of Slope I is larger than that from B, D, and E-segment of Slopes i, n, and m. Using the B-type apparatus comparison of debris flow between neighbouring experiment stations down each slope was possible and increase and decrease in debris flow between stations was calculated. Increase in debris flow at a lower station means an increment

Géomorphologie I 27 of erosion on the slope segment between the neighbouring experiment stations, whilst decrease means that sedimentation is occurring. CONCLUSIONS 1. For Slope i, erosional retreat of the steepest hillside slope was considerably greater than the lowering of the convex slope at the top. On the other hand, the lowering of this convex slope was slightly greater than the erosional retreat of the hillside steep slope for Slopes n and in. 2. Debris scoured from the hillside steep slope is transported by surface runoff and is deposited largely on the concave inflection slope and gradually decreases in quantity towards the concave gentle slope on Slope i. 3. Erosional retreat of the hillside steep slope of Slope i is remarkably greater than that of Slopes ii and in. 4. The quantity of sediments on the concave inflection slope is small and the lowering of the convex slope at the top is slightly greater than that of the hillside steep slope on Slopes u and in. Therefore, the radius of curvature of the convexity at the top and of the concavity at the foot would be increased in length. The following conclusions may be drawn. Slope segments that are very steep such as Slope i (c-segment) may retreat in parallel until a certain time. The concave inflection slope indicates the local base level of erosion for scouring the upper slope or hillside steep slope. The length of the steep segment may be shortened by debris deposition on the concave inflection slope. Therefore, erosional forces acting on the hillside steep slope will be gradually reduced. After a certain time, wearing down of the segment of convex slope at the top would become relatively great compared with the retreat of the hillside slope, as is the case at Slopes n and m. These present agencies acting on the slopes would have repeatedly occurred in past geological times, so that the accumulated effects of their activities are indicated in the faciès of debris layers on the lowest slopes. Baulig, H.C., 1940 Le profil d'équilibre de versants, Ann. de Géog. XLix, 81-97. - 1957 Peneplains and pediplains (translated from the French by C.A. Cotton), Bull. Geol. Soc. Am. 68, 913-30. Berry, L., and B.P. Ruxton, 1960 The evolu-

tion of Hong Kong harbour basin, Zeitschr. f. Geomorph. N.F. 4, 97-115. Birot, P., et al, 1960 Contributions internationales à la morphologie des versants, Zeitschr. f. Geomorph. Suppl. 1, 1-261. Birot, P., P. Macar, and H. Mortensen, eds., 1964 F'ortschritte der internationalen Hangforschung, Zeitschr. f. Geomorph. Suppl. 55 1-238. Bryan, K., 1940 The retreat of slopes, Ann. Assoc. Am. Geog. 30, 254-68. Davis, W.M., 1909 The geographical cycle, Geog. Essays (New York), 245-95. - 1932 Piedmont Benchlands and Primârrumpfe, Bull. Geol. Soc. Am. 43, 384-440. - 1938 Sheetflood and Stream Flood, Bull. Geol. Soc. Am. 49, 1337-416. Derrau, M., 1956 Précis de géomorphologie (Paris). Dumanowski, B., 1964 Problem of the development of slopes in granitoids, Zeitschr. f. Geomorph. Suppl. 5, 30-40. Frye, J.C., 1959 Climate and Lester King's 'Uniformitarian Nature of Hillslopes,' J. Geol. 67, 111-13. Horton, R.E., 1945 Erosional development of streams and their drainage basins: hydrophysical approach to quantitative morphology, Bull. Geol. Soc. Am. 56, 275370. IGU, 1955a Projet de programme de la commission pour l'étude des versants, 1-12. IGU, 1955b Commission on thé Evolution of Slopes, IGU News Letter, vi, 14-18. IGU, 1956 Report of the Commission on thé Evolution of Slopes, IGU. King, L.C., 1953 Canons of landscape evolution, Bull. Geol. Soc. Am. 64, 721-52. Lawson, A.C., 1932 Rain-wash erosion in humid regions, Bull. Geol. Soc. Am. 43, 702-22. Mino, Y., 1942 Principle of Geomorphology (in Japanese) (Tokyo). Penck, W., 1924 Die morphologische Analyse (Stuttgart). Schumm, S.A., 1956 Evolution of drainage system and slopes in badlands at Perth Amboy, New Jersey, Bull. Geol. Soc. Am. 67, 597-646. Strahler, A.N., 1950 Equilibrium theory of erosional slopes approached by frequency distribution analysis, Am. J. Sci. 248, 67396, 800-14. Tanner, W.F., 1956 Parallel slope retreat

28 / Geomorphology in humid climate, Trans. Am. Geophys. Union 37, 605-7. Tricart, J., 1957 L'évolution des versants, L'information géographique 3. Wood, A., 1942 The development of hill-

side slopes, Proc. Geol. Assoc. 53, 124-40. Yoshikawa, T., 1956 Hypsometric analysis of late mature and old mountains in central Korea, Jap. L Geol. Geog. xxvn, 67-78.

P0114 Periglacial slopes ALFRED JAHN University of Wroclaw, Poland Periglacial slopes reflect processes acting in this zone - weathering, solifluction, wash out - while their formation is conditioned by permafrost. Though permafrost represents a certain homogeneity of substratum, we can still see that the rate of periglacial slope formation is different on consolidated rocks from that on loose rocks. Periglacial processes acting on consolidated rocks tend to smooth the slope. Consequently, the action of these processes results in a concave slope, the steepness of the upper section of gravitational slope amounting to 40°. Smoothing occurs through removal of curves (e.g. cryoplanation terraces), slope truncation, and slope retreat (see Demek 1969 and Fig. 1 ). The process of smoothing acts both downwards and upwards. Moreover, there is lateral smoothing which is responsible for the removal of rocky ribs in connection with the troughs of gravitational slopes. Slopes of loose rocks subject to permafrost are degraded chiefly through the action of solifluction. The slope transformation therefore depends on thermokarst processes, i.e. on the balance of denudation, Bd. The important point is the relation between accumulation processes, including weathering, and the removal of the material at every point on the slope surface. The predominance of the removal, i.e. of destructive processes, indicates a positive balance Bd(+) (Jahn 1968). If the slope happens to be in the state of equilibrium (Fig. 2/A), i.e. if the location and thickness of the active layer remains unaltered, no thermokarst phenomena will occur. Such stability may result from the presence of continuous vegetation which counteracts denudation processes. All sections of the slope will then be characterized by a negative denudational balance, Bd(—). If the rate of denudation is increased, the characteristic features of the positive denu-

Fig. 1. Three types of the periglacial slopes: A-concave; B - convex-concave; c- steplike /cryoplanation terraces; SI - solifluction; Sw - slope-wash; Sg - gravitational slope; Sr - rock wall. dational balance, Bd+, will become apparent principally in the central section of the slope (Fig. 2/B). Accelerated or intensified denudation may be ascribed to processes of surficial washdown, but in periglacial conditions it is caused first and foremost by solifluction. Ground ice will be melted in the central sec-

Géomorphologie I 29 Linear erosion, and backward stream erosion in particular (Fig. 2/C), represents another factor responsible for destruction of the slope surface. A flat upland shows no thermokarst phenomena when heat and denudational balances are negative, Bd(—). Stream valleys dissecting the upland help to bring about a change of denudational balance, evidently only in the valley area, since the ground ice which comes to be exposed along the valleys will be melted and, consequently, thermokarst phenomena will develop there. The backward front of the valleys dissecting the upland represents the thermokarst limit.

Fig. 2. The slope and thermokarst processes. For detailed explanations see text. tion of the slope despite the fact that all of the slope area is subject to conditions of a negative heat balance. The importance of solifluction in thermokarst processes, as exemplified in Siberia, has been emphasized in more recent publications by Gravis (1966) and Dylik (1970). See also Mackay (1966).

Demek, J., 1969 Cryoplanation terraces, their geographical distribution, genesis and development, Rozpravy Ceskoslovenske Akademie Ved. Rada Mat. - Prír, R.79, S.4. Dylik, J., 1970 Erozja termiczna (Societas Scientiarum Lodziensis), R.xxiv, 8. Gravis, G.F., 1966 Rol sklonovykh processov v erosynonno-termokarstovom raschleneyi aluvialnykh ravnin Yakutii, Materialy VIII Vsesoyuznogo Mezhduvedomstvennogo Sovieshchaniya po Geokriologii/ merzlotovedeniyul (Yakutskoye kniznoye izdatelstvo, Yakutsk), w.7. Jahn, A., 1968 Denudational balance of slopes, Geographia Polonica, 13. Mackay, J.R., 1966 Segregated epigenetic ice and slumps in permafrost, Mackenzie Delta area, NWT, Geog. Bull. 8.

P0115 Variations in the degradation of Neoglacial ice-cored moraines, St Elias Mountains, Yukon Territory p.G. JOHNSON University of Ottawa, Canada

Most of the Neoglacial moraines of the valley glaciers of the eastern St Elias Mountains are ice-cored. Four examples are the Klutlan glacier (Rampton 1970), the Donjek glacier (Dentón and Stuiver 1966; author's observations 1970), the Kluane glacier, and the Kaskawulsh glacier (author's observations 1971). The distinct sedimentary structure and crystalline form of the ice cores indicate that it is stagnant glacier ice. This contrasts with the observations of Gunnar 0strem (1959; 1962; 1964) in Scandinavia which indicate that ice coring is of snow bank ice. The

ice core is overlain by up to 5m of till, measured on the Donjek moraine, and the processes of degradation are similar on the moraines of the four glaciers. The degree of degradation varies considerably and investigations on the Donjek and Kaskawulsh moraines were carried out to determine the controls on the degradational processes and the variations in the degree of degradation. The processes of degradation of the moraines are controlled by the ice core which exists, or existed, extensively in each of the moraine systems. These processes form a

30 / Geomorphology continuum from cracking of the till, due to its movement as a solid material, to flow of the material in a very liquid state where the ice core has been exposed. The cracking, caused by movement along the saturated ice till interface in the moraine, usually occurs parallel to and along the ridge crests of the moraines. This is the result of the close relationship between the ice surface morphology and the moraine surface morphology. An increase in the degree of saturation at the ice/ till interface results in the occurrence of mass movements of material. The plane of movement of slips on the Donjek moraine is partly composed of the ice/till interface zone. An absence of structure in the mass movement deposits is thought to be due to initiation of the movement along the highly saturated interface before the water content of the rest of the till would be sufficient to initiate the movement. The existence of lakes and lacustrine deposits on the moraine has reduced the erosion because of the poorer heat flux properties of the sands and clays. The effect is not so apparent at the Donjek moraines where the form is still relatively little denuded and numerous lakes still exist in the interridge depressions. At the Kaskawulsh glacier degradation has proceeded to a much greater extent and lacustrine deposits now occur on the high points of the moraines where they initially protected the ice core. These areas are now being eroded by slumping along the edges as the ice core continues to melt out. The considerable variation in the stage of degradation of the Neoglacial moraines in the St Elias Mountains is unlikely to be accounted for by regional climatic variations. There are, however, considerable variations in the microclimatic responses in the valleys due to local influences and it is thought that differences have probably existed since the deposition of the moraines at maximum of the Neoglacial period. The vegetation colonization of the moraines has varied due to these microclimatic variations. The contrast between the well developed vegetation complex on the Klutlan moraine and the poor colonization of the Kluane moraine is very marked. With a poor vegetation cover there is a greater possible heat input to the till and, therefore, greater potential moraine degradation due to ice core melt. Measurements show that variations in

moraine materials and the altitudes of the moraines (except in the case of the Kluane moraine) are insufficient to explain the different degrees of vegetation cover. It is apparent that the main variant in the microclimate of the Donjek and Kaskawulsh areas is the local wind pattern. The wind patterns are dominated by the katabatic effect which is stronger and more consistent in the area of the Kaskawulsh glacier terminus than in the area of the Donjek glacier terminus. On the Donjek moraine there is, consequently, a more highly developed low vegetation cover which is partly responsible for the lesser degree of degradation of the Donjek maximum Neoglacial moraine. The topographic effect on solar radiation input is also of considerable importance. Basic data on exposure to sunlight measured in 1970 and 1971 show consistently longer values for the Kaskawulsh than for the Donjek moraine thus causing differences in heat input and consequently in ice core melt. The role of surges of the glaciers in the degradation of the moraines is of considerable importance. The evaluation of their role is difficult as the effects are superimposed on those of the microclimate. The Klutlan, Kluane, and Donjek glaciers are all known to have surged or show indicative morphological and glaciological characteristics (Meier and Post 1969 ). The surge of the Donjek glacier, which started in 1969, was causing erosion of the moraine where the advancing ice had dug into the moraine and exposed the ice core. This erosion was producing a low relief ridge, with numerous mudflow structures, which is very similar to the highly denuded form of the Kluane Neoglacial moraine. The advanced stage of degradation of the Kluane moraine may be due to the effects of surges in the post Neoglacial maximum period continuously exposing the ice core. Despite the advanced degradation of the moraines of the Kluane glacier there still exists some ice coring. The Kaskawulsh glacier is one of the few glaciers in the area which has not been known to surge. The moraines of the glacier show a fairly advanced stage of degradation but there is little evidence for the very rapid rates of erosion associated with the exposure of the ice core, as observed on the Donjek moraine. Most of the morphological signs of degradational processes are cracks or stabi-

Géomorphologie / 3 l lised cracks which are indicative of ice core melt beneath the till cover. The surges of the Klutlan glacier have not produced the amount of degradation that those of the Kluane glacier appear to have done. Variations in the microclimates of the valleys of the St Elias Mountains are responsible for the varying stages of degradation of the ice-cored moraines. This is due to effects on vegetation colonization and on heat input to the till. The effect of microclimatic factors is partly obscured by the superimposed effects of surges of some of the glaciers which has rapidly accelerated the rate of erosion by exposing the ice cores of the moraines. National Research Council; K. Lowndes; B. Ross. Borns, H.W., and R.P. Goldthwait, 1966 Late Pleistocene fluctuations of the Kaska-

wulsh glacier, Am. J. Sci. 269, 600-19. Dentón, G.H., and M. Stuiver, 1966 Neoglacial chronology, northeastern St Elias Mountains, Canada, Am. J. Sci. 264, 57799. Meier, M.F., and A. Post, 1969 What are glacier surges? Can. J. Earth Sci. 6 (4), 807-18. 0strem, G., 1959 Ice melting under a thin layer of moraine and the existence of ice cores in moraine ridges, Geog. Ann. 41 (4), 228-30. - 1962 Ice cored moraines in the Kebnekajse area, Sweden, Biuletyn Peryglacjalny 11, 271-8. - 1964 Ice cored moraines in Scandinavia, Geog. Ann. 46 (3), 282-337. Rampton, V.N., 1970 Neoglacial fluctuations of the Natazhat and Klutlan glaciers, Yukon Territory, Canada, Can. J. Earth Sci. 7'(5), 1236-63.

P0116 The planetary and hypsometric variation of valley asymmetry H. KARRASCH University of Gottingen, West Germany

Valley asymmetry is one of the more equivocal periglacial phenomena. Its characteristic is a relationship of exposure to the occurrence of steep and flat slopes. Thereby, certain directions are not to be excluded as several authors have stated (especially Budel 1944). Principally all directions can function both as steep and flat slope exposures but not without complying with distinct rules. As Poser and Miiller (1951) have shown and as the author himself (Karrasch 1970) has confirmed by new observations, it is necessary to differentiate between valley segments where the slopes are formed dominantly by indirect denudation (i.e. nearly independent of river or creek erosion) and others where direct denudation (i.e. strongly dependent on river or creek erosion) is prevailing. The peculiar feature of the valley segments of indirect denudation is an asymmetry with steep slopes of NE exposure whereas the opposite asymmetry type - namely steep slopes in sw exposure - can be found in valley sections with recent, or formerly active, lateral erosion. In short, they may be termed as NEf acing asymmetry and sw-f acing asymmetry. It is interesting to observe that, concerning the distribution of sw-f acing asymmetry,

there seems to exist a polar limit. On the other hand, the majority of occurrences of NE-facing asymmetry is reported from arctic North America as well as from northern Siberia. Reference may be made to the studies of Currey (1964) and Hopkins (1962) for northwest and central Alaska, of Cook and Raiche (1962) for Cornwallis Island, of Ermilov (1934) for Gydanskij Poluostrov, and of Presnjakov (1955) for northern Yakutia. This enumeration could be completed by analogous observations from Greenland (Kosiba 1937; Malaurie 1952, 1968) and West Spitsbergen (Klimaszewski 1960). The dominant distribution of NE-facing asymmetry in the present northern permafrost zone, together with the adverse circumstances in the more southern latitudes, justifies regarding it as a planetary variation of the asymmetry phenomenon. The planetary variation corresponds to a hypsometric one. The above observations, however, have not been made in active periglacial regions but instead in those which have been affected by the Pleistocene glacial climates, namely in various Middle European mountains: in the Schwarzwald, in the Schwábische Alb, in the Bayerischer Wald/

32 / Geomorphology

The triangles mark the altitude of the valley floor s a t the end of the N-3E-asymmetric segment (black) respectively at the beginning of the S -M-asymm e t r i c s egment ( whi te) .

Fig. 1. The hypsometric variation in valley asymmetry in selected mountains. Bôhmer Wald, in the Fichtelgebirge, in the Harz, in the Krkonose, and Hruby Jesenik. In these mountains the prevalence of swf acing asymmetry ceases at a certain altitude and is succeeded above that point by the opposite asymmetry type. The separating line between them can be determined exactly; this has been done for the majority of the mountains. An easily practicable method is to register the altitude level of the valley floor at which NE-f acing asymmetry ceases and sw-facing asymmetry begins. The resulting data can be dotted upon a straight line. In case of sufficient dots the threshold value is directly extractable. Mathematically this procedure resembles a nest of intervals. For the sake of clarity in Fig. 1 belts are used instead of straight lines and little isosceles triangles inserted instead of dots. A disadvantage of this method is that the asymmetric valleys are analyzed irrespective of their totally different lengths. This procedure may affect the exactness of the determined altitudinal limit because NE-facing asymmetry also appears in lower parts of the mountains. As a rule, however, such cases are confined to the uppermost valley segments and are of short length. In order to eliminate

this source of error as much as possible, only those valleys of which the NE-facing or swfacing asymmetric segments amounted to at least 0.5km in length were considered. The changing density of the triangles, which is especially striking in the spectrum of the Bayerischer Wald, results from a specific selection of the valleys. On principle, one or more greater valley systems were systematically investigated, including all their tributaries. As soon as the altitudinal level of the threshold value became evident other valleys belonging to this altitudinal level were added in order to render the value additional reliability and precision. The lower limit of the prevailing NE-facing asymmetry ranges between 740m and 600m. As had been expected, it declines from south to north. The difference amounts to 50m140m according to whether the Harz is compared with the Bayerischer Wald or with the Schwábische Alb. Beyond that a decrease is also noticeable from west to east (Schwarzwald/Schwábische Alb 720m-740m, Bayerischer Wald 650m). The question is whether there is a connection between the change of asymmetry, and especially the occurrence of NE-facing asymmetry on the one hand, and a

Géomorphologie I 33 local glaciation of the mountains on the other. This question is motivated by observations which have been made by Fránzle (1959) in the Spanish Cordillera Central and by Rapp (1960) in the Scandinavian mountains. Here corries and valley segments occur with a pronounced E-facing asymmetry. This E-f acing asymmetry is apparently the result of unilateral glacial erosion caused by an increased leeward snow accumulation and a growing ablation in the exposure favoured by radiation. Two arguments make it evident that the interpretation proposed above does not have a comprehensive applicability. 1. Only some, and not all, of the investigated mountains underwent local glaciation. It did not exist in the Schwabische Alb and in the Fichtelgebirge, and in the Hruby Jesenik it was confined to a single corrie. 2. The snow line rises towards the interior of the continent (in the Wurm-Glacial: Northern Schwarzwald 850m, Bayerischer Wald > 1000m), whereas the lower limit of the prevailing NE asymmetry shows exactly the opposite tendency (see Fig. 1 ). There is, however, no doubt that several of the formerly glaciated valleys belong also to the group of valleys with NE-Íacing asymmetry. But even in these cases a genetic interrelation can only be assumed to be exceptional, because there is no congruence of the valley segments in question. NE-f acing asymmetry often extends downvalley beyond the lowermost moraines, and on the other hand there are some valley heads in high altitudes - the former collecting nevé basins — which possess a symmetrical profile. Thus in the Harz valley asymmetry is absent above the 900m contour line. On the contrary, in the Krkonose some NE-facing asymmetric corries exist. Their genesis may possibly be of the kind which Fránzle (1959) and Rapp (1960) described. But these examples are exceptions which do not interfere with the occurrence of the NE-facing asymmetry in general. Obviously the hypsometric and planetary variations have their origin in the same reason. The final question is whether a planetary asymmetry variation is also characteristic of the periglacial zone in the Pleistocene glacial periods. The question can probably be answered in the affirmative for northern Asia and Alaska (Karrasch 1970, pp. 208-11 ). The answer for Middle Europe, however, is

negative, although some contrary opinions have been brought forward (e.g. Dylik 1956). The reason for the lack of the variation is simply the far-reaching continental glaciation. Frorn a linear extrapolation yielded by the hypsometric threshold values it follows that a zonal prevalence of NE-facing asymmetry could not have been expected south of lat. 64° or 65 °N. Biidel, J., 1944 Die morphologischen Kirkungen des Eiszeitklimas im gletscherfreien Gebiet, Geologische Rundschau 34, 482-519. Cook, F.A., and V.G. Raiche, 1962 Simple transverse nivation hollows at Resolute, NWT, Geog. Bull. 18, 79-85. Czudek, T., 1964 Periglacial slope development in the area of the Bohemian Massif in Northern Moravia, Biuletyn Peryglacjalny 14, 169-93. Demek, J., 1964 Pleistozâne deluviale Ablagerungen und die Hangentwicklung in einigen Gebieten der Tschechoslowakei, Sbornik Geologickych Ved - antropozoikum A2, 7-24. Dylik, J., 1956 Coup d'oeil sur la Pologne périglaciaire, Biuletyn Peryglacjalny 4, 195-238. Fránzle, O., 1959 Glaziale und periglaziale Formbildung im ôstlichen Kastilischen Scheidegebirge (Zentralspanien), Bonner Geographische Abhandlungen 26. Herz, K., and G. Andréas, 1966 Ôkologie der periglazialen Auftauschicht ini Kongsfjordgebiet, Petermanns Geographische Mitteilungen 110, 260-72. Karrasch, H., 1970 Das Phánomen der klimabedingten Reliefasymmetrie in Mitteleuropa, Gottinger Geographische Ahhandlungen 56. Maruszczak, H., 1958 Glówne cechy klimatycznej asymetrii stoków w obszarach peryglacjalnych i umiarkowanych, Annales Universitatis Mariae Curie Sklodowska Lublin Sect. B, 11 (1956), 161-237. Poser, H., and T. Millier, 1951 Studien an den asymmetrischen Talern des Niederbayerischen Hügellandes, Nachrichten der Akademie der Wissenschaften in Gôttingen math.-phys. KL, biol.-physio-log.-chem. Abt. Nr. 1. Presnjakov, E.A., 1955 Ob asimmetrii dolin v Sibiri, Voprosy Geologii Azii 2, 391-6.

34 / Geomorphology Rapp, A., 1960 Recent development of mountain slopes in Kàrkevagge and surroundings, Northern Scandinavia, Geografiska Annaler 42, 65-200,

Tricart, J., 1950 Cours de géomorphologie. 2e. Partie: Géomorphologie climatique. Fase. I: Le modelé des pays froids. 1. Le modelé périglaciaire (Paris).

P0117 Regularities in cryogenic phenomena development

E. KATASONOV Academy of Sciences of the USSR

Cryogenic formations in Quaternary deposits ('traces of permafrost') are considered to be signs of former temperature coolings. Their occurrence as regular beds in various parts of sections is a proof of repeated changes in climate. The basic cryogenic phenomena are undoubtedly associated with deep freezing of the earth crust. Yet in some deposits they do not form even in coldest regions of the Arctic. It is necessary to single out one more important circumstance. In conditions of shallow (0.2-4m) bedding of the perennially frozen substratum, development of cryogenic phenomena does not practically depend on an increase or decrease in the mean annual values of the air temperature: bulgunyakhs, ice veins, and various soil veins at present occur in all climatic zones off the coast of the Arctic Ocean up to the Rear Baikal Region. The study of recent cryogenic forms shows that not only low winter temperatures which provide for existence of the perennial substratum are required for their build-up. Development of these or some other concrete phenomena is governed by geological regularities and it is predetermined mainly by genesis, that is by the conditions of accumulation and freezing of the sediments (Katasonov 1964; Katasonov and Soloviev 1968). Geological regularities of cryogenic phenomena development are dealt with in cryolithology - a branch of science concerned with the thickness of permanently frozen sedimentary rocks, their composition, structure, conditions of accumulation, and measurement of freezing. The principal method of cryolithology - freeze-facial analysis - is the singling out of lithogenetic types and faciès of deposits in sections with simultaneous determination of their cryogenic structure. Recent perennially frozen deposits of different genesis include the following major cryogenic formations:

1. Even ice interlayers, smoothly shaped, with a thickness of 0.5-5cm, forming characteristic stratified cryostructures: concave in flood-plain ancj delta deposits and wavy in slope deposits. They are formed at the lower boundary of the seasonal thawed layer but only in very wet and boggy deposits. With ice thawed out there remain relics of cryogenic stratification. Even ice interlayers are the most reliable permafrost genetic signs indicating that the deposits composing them were being formed under permafrost conditions, that the thawing depth did not exceed one meter, etc. 2. Inclined and vertically oriented ice lenses - schlieren - wñich create oblique cryostructures, are formed with a gradual freezing of laterally closed taliks, and are characteristic only of bottom sediments, such as lacustrine and sea coastal sediments. 3. Ice veins. Their biggest forms, creating clear polygons in the plan, are characteristic of very wet deltas, flood-plains and alass depressions. In peatbog deposits they develop everywhere where the climatic conditions permit the existence of the perennially frozen substratum. Slope deposits contain only solitary ice veins which are very often bent as a result of solifluction. In lacustrine sediments and deposits formed under conditions of insignificant humidity (faciès of a dry floodplain, non-boggy slopes), ice veins do not form. 4. Soil veins different in composition, form, and size. Their formation is associated with frost cracking and other cryogenic and physico-geological processes in moisturedeficient but comparatively thick (from 1 to 4m) seasonally thawed layers. Depending on genesis and facial peculiarities of deposits which compose this layer, there occur various soil veins: e.g. outbending soil veins —in deposits of riverside shoals; crack filling soil veins - in slope loams; ditch crack filling soil veins - in deposits of a dryflood-plain,etc.

Géomorphologie I 35 (Katasonov and Soloviev 1968). 5. Ice cores of large heaving mounds - bulgunyakhs. They are associated in the main with lacustrine deposits. Phenomena under consideration in corresponding deposits occur both in northern and southern districts of the permafrost region. They develop simultaneously, side by side, but still independently. This fact is explained by conditions of sedimentation regulating moistening of rocks and depth of their thawing, determining the processes of ice segregation, the character of frost cracking (in perennially frozen substratum or in a seasonally thawed layer), the mechanism of crack filling, etc. The temperature factor, as was already noted, does not play a significant role. Climate, dry or humid, stimulates only irregular distribution of cryogenic phenomena. In the north, where because of negligible evaporation primarily boggy deposits are formed, ice veins prevail-soil veins occur rarely; on the contrary, in droughty regions of central Yakutia soil wedges are observed more frequently. During the Pleistocene, development of cryogenic phenomena was evidently dependent on the same regularities. Knowledge of these regularities makes possible more accurate ideas about the conditions of periglacial deposit formation in Europe. Judging by

numerous publications there are no typical faciès of boggy plains in their sections of which ice veins and stratified cryostructures are characteristic. There prevail alluvial floodglacial deposits and, mainly, slope deposits with originally soil veins — similar to the veins which are formed at present in the seasonally thawed layer. In our opinion, in the periglacial zone of Europe mainly soil veins were syngenetically formed. Their appearance in various parts of sections is rightly to be associated with changes not in climate but in the conditions of sedimentation. In conclusion it is worthwhile to note that investigation of cryonegic phenomena, ground ice in particular, is of great practical importance. Inferences concerning the distribution of ground ice are necessary for builders, miners, and other specialists who master northern regions of the USSR. Katasonov, E.M., 1962 Cryogenic structures, ice and soil veins as genetical signs of perennially frozen Quarternary deposits, Questions of Cryology when studying Quarternary Deposits (Moscow). Katasonov, E.M., and P.A. Soloviev, 1969 Guide to trip round Central Yakutia: paleogeography and periglacial phenomena (Yakutsk).

P0118 An hypothesis on the paleomorphology in the south of central Mexico A. LOPEZ-SANTOYO National University of Mexico Tectonism and associated volcanic phenomena are undoubtedly some of the most important factors in the development of landforms. In central Mexico the ancient relief has been altered to a great extent and today the rocks , in a large area are of volcanic material that has been coming to the surface since the Tertiary. One of the main tectonic features is the long fracture that crosses the country from west to east at approximately 19°N. This fracture affected the southern part of the central plateau and influenced the development of the depression of the Balsas River. Erosiona! processes were accelerated because of the great difference in height of the various blocks and the steep general slope along the south side of the fracture zone. These condi-

tions have caused important captures of the inner basins; e.g. in the basin of the Balsas River the watershed of the northern slope is likely the one that presents less stability. It changes relatively fast and is always incorporating new areas from other basins. In south central Mexico the Puebla-Tlaxcala Valley, the Valley of Mexico, and the Valley of Toluca are three geomorphic regions aligned from east to west, and clearly separated by two mountainous systems of volcanic origin. The main element of the Puebla-Tlaxcala Valley is a big central plain with a gentle slope from northwest to southeast. The topographic surface is flat, interrupted only by some elevations. The Atoyac River drains this valley and flows into the

36 / Geomorphology Balsas River. There is evidence that the Valley of Puebla-Tlaxcala was an interior basin that was captured later. The Valley of Mexico (also known as the Basin of Mexico) has been drained by artificial channels in the north to prevent floods in the city. Some lakes were found in historical times in the region and it is known that these lakes were not only bigger in the past but also constituted one single lacustrine unit that suffered later desiccation. The Valley of Toluca is opened to the west and is part of the high basin of the Lerma River. It has also as a principal element a large flat plain with a very gentle slope. This region was also an interior basin which afterwards became captured. Thus, it is possible that before the formation of the Sierra Nevada de Mexico, in which the Popocatepetl and Iztaccihuatl volcanoes are found, the Basin of Mexico and the one corresponding to Puebla-Tlaxcala Valley formed a lacustrine unit which in the western part maybe enclosed also the Valley of Toluca. The geomorphological factors which lead us to that conclusion are the following: 1. The Puebla-Tlaxcala Valley has the characteristics of an ancient lacustrine basin, shown by the bulky horizontal sedimentary strata and its great influence on the topographic relief. In the northeastern section of this valley the strata reach a thickness of more than 100m. They are composed of diatomite, that is, silicate remainders of microscopic algae. These deposits of organic origin could not have been formed in a small lake. The region where they are found forms the limit of the Valley of Puebla-Tlaxeala dividing it from the small basin of Lake Xalnene. The region of diatomite was surely lifted while the small sierras of volcanic origin, aligned with it, were being formed.

2. Other evidences are found along the new road that goes from Mexico City to Puebla. Lacustrine deposits are alternated in an uneven way with volcanic material. The thick sedimentary strata and their folding concordant with the superficial forms of the land by the effect of internal pressures of volcanic origin are particularly noticeable in the hills that limit the flatlands of the Basin of Mexico. 3. It is also possible to see lacustrine sediments mixed with the volcanic material hi the sierra that separates the Valley of Toluca from the Valley of Mexico. In the Valley of Toluca there is another important deposit of diatomite. 4. Finally the calcareous platform that forms the base on which the volcanic material is found has been deformed into a deep depression. This is evidenced by the fact that limestone outcrops encircle the region while at its centre the calcareous rock is found at a depth of 2000m under the ground level. The lacustrine conditions existed for long periods of time, possibly from the end of the Mesozoic; they also dominated after the volcanic activity of the Tertiary and Quaternary and consequently are of great importance in the study of the paleogeomorphology of Mexico. The knowledge of those conditions also constitutes the basis for the interpretation of many of the actual processes that modify the present morphology. Arellano, A.R.V., 1948 La composición de las rocas volcánicas en la parte sur de la cuenca de México, Boletín de la Sociedad Geológica Mexicana. Comisión Hidrológica de la Cuenca de México, 1961 Informe sobre la geología de la cuenca del Valle de México y zonas colindantes, Publicación no 6 (México DF).

P0119 Debris accumulation an talus slopes in Surprise Valley B.H. LUCKMAN University of Western Ontario, Canada

The measurement of debris accumulation on talus slopes has been limited to a few experiments using small areas of sacking, netting, or tarpaulin (Rapp 1960; Stock 1969; Gardner 1970a; Prior et al. 1971 ), usually placed near the top of the slope, or inventories of debris on spring snow (Rapp 1960; Gardner

1970b). Results from Surprise Valley in Jasper National Park, Alberta, indicate that accumulation can be successfully measured anywhere on the slopes by recording the deposition on a network of anchored polyethylene squares and cleaned boulders. Initial sampling networks were set up on

Géomorphologie I 37 TABLE 1. Accumulation on the sample sites (all volumes are in eu cm/sq m) All boulders

All squares

Surprise n

EVS

Volume

1969 1970 1971

1969 1970 1971

1970 1971

1970 1971

Less than 1 cc 1-10 10-100 100-1000 10M054 10M0 10M076 10M0

55.9 11.6 11.1 10.0 8.0 2.9 0.4

39.8 6.8 16.2 18.7 12.4 4.9 0.2

38.1 17.4 11.1 15.8 15.8 1.6

31.9 8,.7 9.9 4,.3 18.7 21,.7 23.1 34,,8 12.1 26..1 4.4 4,,4

9.0 10.2 37.5 31.5 10.2 1.7

35.1 22.4 23.1 10.5 6.1 3.3

82.4 2.2 3.7 6.3 4.4 0.7

27.8 4.0 15.0 18.2 17.5 16.7 1.6

Sites sampled Sites set

708 730

951 885 1029 1029

63 72

91 99

235 265

214 265

136 138

126 138

36.2 11.2 16.9 14.7 12.0 6.9 0.8 0.1

seven widely differing scree slopes in August 1968 and supplemented in September 1969. The boulder sites consisted of relatively large boulders (sampled areas mainly between 0.2 and 1.0 sq m) with smooth surfaces rising above the surrounding debris and sloping at less than 20°. The polyethylene squares were 2.32 sq m in area and anchored by painted marker stones. The sampling network was arranged as a series of profiles and transects of each slope with the large majority of the sites within the basal third of the slope. Triaxial measurements were made of freshly accumulated debris during the summers of 1969, 1970, and 1971. These figures were converted to volumes by the use of empirically derived coefficients based on particle shape and the accumulation expressed as cubic centimetres per square metre of sampled area. Further details of the technique and some preliminary results are given in Luckman (1971). The boulder sites require little attention after the initial cleaning although 1-2 per cent of the sites were moved by snow avalanche activity in each of the three years. If those sites covered by late-lying snow and high lake levels at the time of sampling are excluded, between 90 and 97 per cent of the sites produced results in each of the three years (Table 1 ). The squares require more attention since the marker boulders are more easily moved. The upper squares at some sites were completely destroyed by avalanches (20 per cent of all squares in 1971) and their marker boulders swept downslope. However, mapping the distribution of these boulders yields valuable information about the range and character of avalanche erosion (Luckman 1971, 1972) and compensates for

46 69

the greater amount of maintenance needed. The point estimates for the volume of deposition given by both methods are similar except that, because of their smaller area and the nature of the distribution being sampled, the boulders have more observations in the tails of the distribution. If those sites with no accumulation are excluded, the distribution of accumulated volumes is approximately log-normal in form, strongly reflecting the size distribution of individual particles (Table 1). The intensity, distribution, and character of deposition varies from year to year at each site and between individual sites in any one year. Also there is no consistent ranking of the amount of activity at different sites in consecutive years. These variations are masked in the aggregate data but can clearly be seen in the results from the sites at Surprise n and the Eastern Valley Side (EVS) . They reflect the many complex controls of the number, type, and geomorphic activity of avalanches at a particular site in a particular year. The variations, together with the bias in the location of sample sites, preclude any attempt to derive meaningful average figures from the present data, but the maximum debris shift observed at any site can be estimated at 2030 eu m of material. Although these results indicate widespread deposition on the sampled talus slopes, these data cannot be used to make inferences about the erosion of the cliffs. The sedimentary characteristics indicate that most of the deposition is from snow avalanches and the role of rockfall is underestimated, since the boulders only record debris which ablates onto them from the snow and therefore there is no record of the deposition when the scree is

38 / Geomorphology snowfree. Also the bulk of the material incorporated in the avalanches appears to be eroded from the higher parts of the talus (Luckman 1972). The figures are therefore more representative of the amount of debris moved annually by avalanches (with a small rockfall component) and indicate their great importance in the development of these talus slopes. All of the talus slopes studied showed evidence of avalanche deposition in at least one of the three years and on the six sites where avalanche erosion occurred it was clearly the most significant agent of debris movement. National Research Council; McMaster University. Gardner, J., 1970a A note on the supply of material to debris slopes, Can. Geog. 14, 369-72. - 1970b Geomorphic significance of avalanches in the Lake Louise area, Alberta,

Canada, /. Arctic Alp. Res. 2,135-44. Luckman, B.H., 1971 The role of snow avalanches in the evolution of alpine talus slopes, Inst. Br. Geog. Spec. Pub. 3, 93110. - 1972 Some observations on the erosion of talus slopes by snow avalanches in Surprise Valley, Jasper National Park, Alberta, Guidebook for Commission Meeting Ca6, IGU. 1972, in press. Prior, D.B., N. Stephens, and G.R. Douglas, 1971 Some examples of mudflow and rockfall activity in north-east Ireland, Inst. Br. Geog. Spec. Pub. 3, 129-40. Rapp, A., 1960 Recent development of mountain slopes in Karkevagge and surroundings, Northern Scandinavia, Geog. Annlr. 42, 72-200. Stock, R., 1969 Morphology and development of talus slopes at Ekalugad Fjord, Baffin Island, NWT. BA thesis, U. Western Ontario.

P0120 Wind factor in dune formation at Great Whale, Quebec M.A. MacFARLANE Sir George Williams University, Canada In studying the coastal dunes at Great Whale, on the eastern side of Hudson Bay, Cailleux found a discrepancy between their form and the present day winds, although from the grain size distribution they appear to be modern (cf. Cailleux and Hamelin 1969-70). It seemed that a partial explanation might be found by analysing the wind distribution in relation to the regional climatic conditions. This exploratory study is based on the 1969 hourly data of the Great Whale weather station. The approach taken is that of Bagnold: i.e. computing the transporting power (T) of the wind from different directions using a relation of the form T = (wind speed — threshold speed)« (Bagnold 1954). It is assumed that for dry sand the threshold speed is 4.5m/sec (lOmph) ; however only winds over 5.4m/sec (12mph) have been counted (Simonett 1960). This method of estimating transport is commonly believed to be useful only when the sand is fairly dry. Great Whale, however, has a subarctic maritime climate so data were limited to periods when the sand would likely

be dry. Therefore hours with the following conditions were excluded: 1. When there was a stable snow cover. From examination of the data and comparison with climatological averages it was decided that in January, February, and March the dunes were almost certainly not blown free of snow (Potter 1965). In 1969 December and April were transition months and are treated separately. 2. During periods of precipitation of any form or intensity. This restriction might sometimes be too severe, e.g. some forms of frozen precipitation may not wet the ground. 3. When fog was reported. 4. While the surface was drying. Although none of the above conditions might apply at the hour, the surface could remain wet for some time. It was assumed that the sand was wet as long as the relative humidity observations were higher than 90 per cent (cf. Monteithl957). The analysis of the 1969 winds is presented in Figs. 1 and 2 in the form of roses of assumed sand movement, where each arrow

Géomorphologie I 39

Fig. 1. Rose of potential sand movement in snow-free period, May to Nov. 1969. Sand surface assumed dry.

Fig. 3. Wind transporting power in the snowfree period, May to Nov. 1969. No assumptions made about surface wetness.

shows the percentage of transporting power available to move sand in that direction. For comparison Fig. 3 shows, in a similar manner, analysis of winds in the snow-free period using all winds greater than 5.4m/sec. In interpreting these diagrams allowance must be made for the many rough rocky surfaces and considerable amount of grass and shrubs that grow on parts of the dunes. Local topography can be expected to have a major influence on the wind at actual dune sites. The alignment of the coastline, an important consideration in this study, has been shown. In spite of simple analysis a comparison of Figs. 1 and 3 for the snow-free period shows that a partial explanation of present sand

movement, in relation to the winds, seems to have been found. 1. The power of the onshore westerly wind is less important when surface wetness is considered, i.e. 27 per cent when using all weather, 20 per cent for dry sand, and in transition months only 12 per cent. 2. On the other hand the increased percentage of power from the south winds ( 12 to 17 per cent) and southeast winds (21 to 25 per cent) does seem to be correlated with the presently active sand faces in the dunes north of the base (the blowouts on the south side of these dunes and the movement of the dunes to the north). 3. There are some interesting seasonal effects such as the high transport capacity towards the south and west in the transition months (23 and 29 per cent). These and other monthby-month variations reflect the different frequency distributions of synoptic weather situations. The transport capacity is also sensitive to changes in the regional circulation from one year to another. For instance, in 1969 the winter was somewhat warmer than normal while the snow-free period was generally cooler and drier. These results indicated that further study would be worthwhile and analysis is now underway on such questions as the monthly breakdown, the influence of local topography and roughness, and the mass transport.

Fig. 2. Rose of potential sand movement in the transition months, April and Dec. 1969, Sand surface assumed dry.

Canadian Atmospheric Environment Service, Centre d'Etudes Nordiques, Université Laval.

40 / Geomorphology Bagnold, R.A., 1954 The Physics of Blown Sand and Desert Dunes (London). Cailleux, A., et L.-E. Hamelin, 1969-70 Poste-de-la-Baleine (Nouveau-Québec): exemple de géomorphologie complexe, Rev. Géomorphologie dynamique, no 3, 131-50.

Monteith, J.L., 1957 Dew, Quart. J. Roy. Met. Soc. 83, 322-41. Potter, J.G., 1965 Snow cover, Climatological Studies, no 3 (Atmospheric Environment Service, Toronto). Simonett, D.S., 1960 Development and grading of dunes in western Kansas, Ann, Assoc. Am. Geog. 50, 216-41.

P0121 Cassures périodiques et modelé éolien dans les grès péri-tibestiens (Tchad) MONIQUE MAiNGUET-MiCHEL Université de Reims, France

Aux confins septentrionaux du Tchad entre 16 à 23°N et 14 à 23°E s'étendent 650.000 km2 de grès curieusement façonnés en un système de crêtes et de couloirs concentriques au massif volcanique du Tibesti. Ces grès constituent au pied occidental de ce massif la bordure est, cambrienne* de la cuvette du Djado; à son pied oriental ce sont ceux de la cuvette des Erdi, succession d'auréoles paléozoïques et mésozoïques qui plongent du sw au NE vers le bassin de Koufra et s'appuyent au sud, en Ennedi, sur le massif du Ouaddai. Le climat est désertique, les précipitations sont inférieures à 100mm (l'isohyète 100 passe au Nord de Fada, 17° 13')L'alizé monodirectionnel soufflant huit mois sur douze du NE au sw, chenalisé entre le Tibesti et l'Ennedi, est responsable de violents et fréquents vents de sable. La morphologie, tripartite, comprend trois étages : 1. périglaciaire, au-dessus de 1000m (cœur du massif du Tibesti) 2. d'action hydrique, entre 600 et 1000m 3. éolien, au-dessous de 600m Ce dernier étage est l'objet de cette étude. Il comporte à partir du Tibesti : Une bande de crêtes et de couloirs en alternances périodiques. Le phénomène atteint son développement optimum au NW du massif et selon un arc de cercle qui le longe de l'est au SE puis au sud. A 200km de la basé du Tibesti, une aire où domine l'accumulation sous forme d'un nappage de sable et de trains de barkhanes. Le sable venu de Libye par le NE contourne le massif par le SE et le sud puis remonte avant de retrouver le flux éolien venu du NW avec lequel il forme l'Erg de Bilma. Celui-ci se situe au point d'essoufflement du vent qui a tourné à la périphérie du Tibesti.

Une aire qui s'étend jusqu'au lac Tchad ; le phénomène d'accumulation sableuse y reste sous l'influence de la déflexion éolienne périTibesti. L'ensemble constitue une vaste unité attribuée au vent à cause du parallélisme entre : les axes de migration barkhanique - qui matérialisent la direction du vent moyen résultant annuel (Clos Arceduc 1868) -etles axes des crêtes et couloirs, les deux étant conformes à la déflexion de l'alizé autour du Tibesti. Les traînées souvent plusieurs fois kilométriques laissées par le vent au sol et le système de crêtes et de couloirs, traînées résultant de la combinaison derrière un obstacle topographique d'une flèche sableuse et/ou d'une strie de corrasion peu profonde réduite parfois à un simple grattage de la patine sans véritable sillon. Ces arguments suffisent à prouver que le mécanisme de façonnement est éolien, mais l'ampleur du phénomène, la périodicité des crêtes et des couloirs, les changements brutaux de périodicité, l'homogénéité locale de la dimension des crêtes sont expliqués par les données structurales. Le grès, en effet, roche privilégiée car la plus apte à la désagrégation granulaire, explique l'ampleur du phénomène, apportant à la fois l'outil et le substrat nécessaire à la corrasion. La striation géante n'apparaît pas dans les basaltes, ni les roches cristallines ou métamorphiques du socle, exceptionnellement dans les schistes, lorsque ceux-ci, redressés, offrent à la corrasion les plans de schistosité alignés selon la direction éolienne. Des cassures périodiques en nombreux faisceaux, réponse du grès (roche cassante et peu ployante) à la surcharge isostatique engendrée par le volcanisme du Tibesti, ex-

Géomorphologie / 4 l pliquent la périodicité. Les cassures sont très denses, les photographies aériennes permettent d'en dénombrer 40 par km2. L'élargissement des couloirs croît lorsque l'écart primitif entre la direction des diaclases et celle du vent moyen annuel croît, jusqu'à un angle de tolérance de cinq grades, verifiable sur les photographies aériennes ; au-delà de cette valeur, le vent emprunte un autre cheminement. Lorsque la coïncidence vent-cassure est imparfaite dans un premier temps, le vent s'adapte à la cassure, puis dans une étape suivante, il l'adapte à sa direction en l'élargissant en couloir. Les changements brutaux de périodicité correspondent indirectement aux variantes topographiques, elles-mêmes reflets des changements de bancs; chaque banc apportant selon ses caractères physiques sa périodicité spécifique de diaclasage. Les photographies aériennes (couverture au l/50.000e IGN France) apportent la preuve qu'au passage d'une couche rocheuse à une autre la direction des diaclases se conserve tandis que change leur fréquence. Or, les diaclases sont d'autant plus fréquentes que les couches affectées sont minces.

Autour du massif du Tibesti les vents de sable exploitent un système de cassures périodiques, qu'ils accentuent, réalisant un remarquable exemple de soufflerie naturelle, responsable d'un paysage purement éolien parmi les plus grandioses du globe. Le CNRS, l'Institut Géographique National (France). Clos-Arceduc, A., 1969 Essai d'explication des formes dunaires sahariennes. Etudes de photo-interprétation 4, Inst. Géog. Nat., Paris. Durand de Corbiac, H., 1958 Autant en emporte le vent, ou L'érosion et l'accumulation éoliennes aux alentours du Tibesti, Bulletin d'Information de l'Association des Ingénieurs Géographes 11, juillet, 147—55. Mainguet-Michel, M., 1971 Le modelé des grès : problèmes généraux. Thèse de Doctorat d'Etat, Inst. Géogr. Nat., Paris. Verstappen, H. Th., et R.A. van Zuidam, 1970 Orbital photography and the géosciences: a geomorphological example from the central Sahara, Geoforum 2, 33— 47.

P0122 Magnitude and frequency of processes operating on arctic beaches, Queen Elizabeth Islands, NWT, Canada S.B. MCCANN McMaster University, Canada Within the broad group of arctic beaches, which on the whole may be classed as low energy beaches, there are considerable differences in environmental conditions as regards ice cover, wave climate, and tidal range. The Alaskan beaches (Rex 1964; Hulme and Schalk 1967), for instance, are quite different from those of the Canadian arctic archipelago (McCann and Owens 1969; Owens and McCann 1970), which provide the subject matter of this paper. However, the prime characteristic of all arctic beach regimes is the long closed season when beaches are locked in ice, and the short - often very short — open season during which the beaches are active. Though ice on the beach produces certain distinctive features along arctic shorelines, it is the inhibiting rather than the positive role of ice which is important in considering beach regimes. This inhibitory role has two aspects: ice at sea, either as pack ice in summer or a

more solid cover in winter, prevents or severely restricts wave generation; ice on the beach prevents the action of breaking waves and swash from working the beach sediments. The combination of short open water season, pack ice at sea, and fast ice on the beach means that in certain years very little significant wave action may take place. There may be little change in beach profile and little redistribution of beach material. Annual rates of sediment transport are likely to be low and long-term changes in the coastline somewhat slow. There are considerable variations in activity from year to year. In this context it becomes necessary to pay particular attention to problems associated with the magnitude and frequency of periods of significant wave action. The principal question concerns the probability of occurrence of winds of sufficient strength and duration, from relevant compass directions,

42 / Geomorphology at such times during the open water season as the sea is clear of pack ice and the beach is clear of fast ice. In the case of the McMaster investigations in the Radstock Bay-Cape Ricketts area of sw Devon Island this question became highly relevant after three seasons of field investigation, during which beach changes had been monitored during the first half of each open water or active beach season (i.e. late July and August). Only one major storm, producing waves capable of really significant action on the beach, had occurred during the periods when the beaches were under surveillance - on 11-12 August 1969. In order to comment yalidly on the field results concerning this storm and their relation to more normal conditions, and also in order to plan further work, it was necessary to place it in context and know something of the frequency of such events. There follows a brief account of a first attempt to carry out this task, using wind data for Resolute Bay on Cornwallis Island some 105km west of the study site and ice reconnaissance data for the Barrow Strait-Lancaster Sound sea area. The wind data was taken from the Arctic Summaries, which give wind speed and direction (for sixteen points of the compass) on a three-hourly basis for the ten-year period 1959-68; the ice data is irregular and incomplete. In order to make use of the wind data to obtain information about possible wave conditions which may have affected the study beaches on Radstock Bay the following assumptions were made: 1. that the important winds blow from the SE quadrant: this is the direction of maximum fetch for the beaches and includes the sea area most likely to be free of ice in the open water period. 2. that to generate significant waves the wind must blow from the SE quadrant at a speed of at least 20mph for at least six hours. 3. that the three months July, August, and September constitute the period of potential wave action on the beaches. A summary of the wind data organized according to these assumptions shows that August is the stormiest month, containing a half or more of the total of significant wave generating winds in each of the duration categories. For winds over 20rnph which blew for 24 hours or more continuously, August accounts for one-third more again

than the other two months combined. Only three out of the ten years did not have a 24hour period of 20mph winds in August, whereas six of the Julys and seven of the Septembers had no such storms. In the two stormiest years in the decade, 1960 and 1966, 297 out of the 360, and 204 out of the 366, hours of wind over 20mph for at least six hours occurred in the month of August. Not all of the above noted occurrences of suitable winds for wave generation coincided with suitable ice-free conditions, but in this regard also August proved to be the best month. In order to rank the storm of 11-12 August 1969 in relation to events in the previous ten years it becomes necessary, first, to define a 'storm' (or period of significant wave action) and then to isolate the largest storm in each year. Straightforward computing and plotting of 'storm' frequency curves in a manner similar to the calculation of flood frequency curves for river discharge may then be carried out. Two methods were used to define thé largest storm in each year. In the first, the duration of winds over 20mph is important-the longest continuous period of winds over this speed being considered as the maximum storm of the year: in the second, wind speed is taken into account and the largest value for miles of wind, tallied for consecutive three-hourly observations over 20mph, is considered as the maximum storm. In both cases, of course, the directional constraints (within the SE quadrant) are applied. Slight differences in ranking between years occur with the different methods, but the event in each year remains the same. For only one year of the ten is the maximum storm less than that of 11-12 August 1969, and in certain years there are two storms which are greater in terms of wind duration and speed. If only wind strength and duration are considered, then a storm with wave action and effects on the beach similar to those documented in 1969 might be expected to occur every year. However, in at least five of years 1959-68 the maximum storm as defined above, on the basis of wind strength and duration, would have had little effect on the beaches of Radstock Bay because of the inhibiting presence of ice either on the beach or at sea. It is difficult to consider the dual probabilities of both wind and ice conditions, especially as the ice cover data is not complete and is based on an irregular schedule of

Géomorphologie I 43 reconnaissance flights. None the less, the calculations outlined above, considerable assumptions notwithstanding, do provide some indication of the magnitude and frequency of expected events, and enable the one major event recorded in the field to be placed in context as a once in two- or threeyear occurrence. During this event, the 'storm' of 11-12 August 1969, the whole of the beach along the outer western shore of Radstock Bay was combed down and the profile lowered by 0.3 to 0.5m in the mid-to-high tide zone; erosion of the low bluff in raised marine deposits, which backs the modern beach, was considerable; and great amounts of material were transported alongshore. Several marked pebbles travelled over 300m in the one high tide period. Estimated wave height was 11.3m at a frequency of 9-11 per minute: the waves broke some 10m offshore and had a powerful swash which reached over a metre above normal higher high water level. The effectiveness of these waves in eroding the beach and in longshore transport was not impeded at all by the presence of ice on the

beach for an earlier period of large waves on 9-10 August had broken up the remaining beach fast ice and whatever ice floes had remained stranded. This point illustrates the difficulties of isolating and defining single events such as a 'storm,' for part of the effectiveness, in geomorphic terms, of the wave action on 11-12 August was related to events which had occurred previously. Hulme, J.D., and M. Schalk, 1967 Shoreline processes near Barrow, Alaska: a comparison of the normal with the catastrophic, Arctic 20, 86-103. McCann, S.B., and E.H. Owens, 1969 The size and shape of sediments in three arctic beaches, sw Devon Island, NWT, Canada, Arctic Alp. Res. 1, 267-78. Owens, E.H., and S.B. McCann, 1970 The role of ice in the arctic beach environment, with special reference to Cape Ricketts, sw Devon Island, NWT, Canada, Am.J. Sci. 268, 397-414. Rex, R.W., 1964 Arctic beaches, Barrow, Alaska. In R.L. Miller, éd., Papers in Marine Geology (New York), 384-400.

P0123 Some quantitative aspects of periglacial slope deposits in southwest England D.N. MOTTERSHEAD Portsmouth Polytechnic, England This paper attempts to identify in quantitative terms, the sedimentary properties of some fossil periglacial deposits in southwest England. The study area is a coastal strip in south Devon some 5km in length. Geologically the area falls within the Start-Bolt thrust zone and the local bedrock consists of quartz-mica schist and hornblende-chlorite schist. The periglacial deposits form an apron beneath a fossil cliffline cut in solid bedrock. Contemporary marine erosion has provided good exposures in the periglacial deposits. The deposits consist of the whole range of particle size from clay to boulders. They are comprised basically of stones and boulders set in a matrix of red or red/brown sand and clay. The poorly sorted character suggests that no sorting agent, such as running water or wind, has been active, and the material is therefore the result of simple downslope mass movement. The most striking characteristic is the

coarseness. Particle size analysis was carried out on eighteen samples. Of these, thirteen had a maximum particle diameter larger than —60 (65mm), whilst four were coarser than —80 (260mm), and even coarser material exists elsewhere. Sediment characteristics were computed for the samples using 0 units, as described in King (1966). The following formulae were employed: Mean particle size = (010 + 030 + 050 -f 070 + 090)/5 Skewness = (016 + 084 - 2050) /[2 (084 016)] + (05 + 095 -2050)/[2 (095 - 05)] Sorting = (085 + 095 - 05 - 015) /5.4 The mean particle size over eighteen samples is shown to be —1.280 with a standard deviation of 2.210. The coarseness of the deposits is demonstrated, as also is their variability. These characteristics are emphasized by the stone content, defining a stone as a particle greater than —20 (4mm).

44 / Geomorphology The mean value of 54 per cent, and standard deviation of 19 per cent, further underline the variability of the deposits. Skewness shows a mean value of +0.23, indicating a dominance of positively skewed samples, although there is a significant tendency towards negative skewness. The sorting values show a near normal distribution with a mean value of 4.300. The well-defined distribution, together with the relatively small standard deviation of 0.79, indicate that the material is typified by a fairly narrow range of sorting values in contrast to the wide variability of the characteristics described above. All the samples fall within the very poorly sorted, and extremely poorly sorted, categories (King 1966). In order to investigate the effects of distance of transport, three characteristics were plotted against downslope distance of flow. This latter variable was arbitrarily estimated as equivalent to the distance downslope in a seaward direction from the 100-ft (33m) contour to the point of each exposure. It is noticeable that the coarsest material tends to be present in the coves, where exposures in the periglacial material are close to the old cliffline. At Langerstone Point some 270m out from the slope the deposit is comparatively fine, the two samples taken yielding maximum particles of —5.90 and —6.30 respectively. These observations are supported by the correlations. Mean particle size plotted against distance shows a correlation of +0.59, significant at the 99 per cent level. The same relationship is shown by stone content (r = —0.57, also significant at the 99 per cent level). Thus the deposits are shown to be progressively finer with increasing distance of transport. The characteristic of skewness also shows a correlation with distance. The correlation coefficient of —0.55, significant at well above the 95 per cent level, denotes a change from positive to negative skewness during transport. The decrease in particle size with distance can be explained in two ways. It is possible that the finer material is selectively transported with the coarser material being left behind. A second possibility is that the material becomes comminuted by abrasion and frost shattering during transport. This latter contention is supported by the skewness correlation. It is suggested here that the shift

from positive to negative skewness is caused by abrasion of adjacent stones during transport. Thus a portion of fine material is produced directly from the coarse, leading to negative skewness. This shift in skewness is similar to that described by Krumbein and Tisdel for fluvial sediments. They postulate that regolith materials are characterized by positive skewness, whilst transported sediments become negatively skewed. It is generally held (Embleton and King 1968; Benedict 1970), that two processes are responsible for the mass movement of slope material in a periglacial environment - frost heaving and saturated flowage. These will be dealt with in turn. A potent mechanism of heaving is provided by the formation of segregated ice lenses within slope forming materials. Water is drawn upwards towards the freezing zone by molecular pressure through the pore spaces, and soil particles are excluded from the growing ice mass as the water crystallises to ice. The formation of segregated ice depends upon the availability of sufficient moisture, and on the size of pore spaces within the soil, which are largely dependent on the particle size distribution. Beskow (1935) and Cailleux and Taylor (1954) have defined the particle size distributions of soils which are susceptible to the formation of segregated ice. When the particle size curves derived from the less than 2mm fraction of the slope deposits are compared, it is seen that all the samples fall within the range of frost heaving materials. Thus in environmental conditions which produce deep perennial or seasonal frosts and available soil moisture^ these sediments would be prone to the formation of segregated ice. Where the sediments lay on a slope, the frost heaving would result in mass movement downslope. The response of a sediment to varying water content can be defined with reference to the liquid and plastic limits, and to the plasticity index, together known as the index properties. The most important factors influencing index properties are the type and proportions of clay minerals present. The mean value of liquid limit for six samples is 26 per cent, the plastic limit 17,7 per cent, and the plasticity index mean is 8.3 per cent. The plasticity index is in all cases low (in many clays it attains values of 50 or more). The low plasticity index can be attributed to

Géomorphologie I 45 the fact that clay never forms a large proportion of the sediment, reaching a maximum of 14 per cent in the samples tested. The low values of the liquid limit are also important. This means that an amount of water equivalent only to 25-30 per cent by weight of the sediment is sufficient to cause the sediment to pass into liquid state and thus start to flow. Under periglacial conditions, surface water during the summer season is abundant due to the thawing out of the frozen ground. Since low summer temperatures retard evaporation, and the still frozen layer prevents downward penetration, the meltwater tends to remain in the upper layers of the soil, saturating it. Thus when susceptible materials are present, saturated flowage will take place. It is concluded therefore that, given suitable environmental conditions, both frost

creep and saturated flowage are processes which may have contributed to the formation of these deposits. Benedict, J.B., 1970 Downslope soil movement in a Colorado alpine region: rates, processes and climatic significance. Arctic Alp. Res. 2, 165-226. Beskow, G., 1935 Tjálbildningen och tjállyftningen med sarskild hánsyn till vagar och jarnvagar. Sver. Geol. Under s. Arsbok 26, Ser. c. no. 375. Cailleux, A., and G. Taylor, 1954 Cryopédologie, étude des sols gelés. Exp. Pol. Fr. iv (Paris). Embleton, C., and C.A.M. King, 1968 Glacial and Periglacial Geomorphology (London). King, C.A.M., 1966 Techniques in Geomorphology (London).

P0124

Morphology of the lunar formations SUDERSHAN MUKERJEE National Atlas Organisation, India The investigation in the present decade of the moon by instrument-carrying spacecraft and manned expeditions has made its study one of the most important aspects of current research. However, the study of the lunar surface dates back to the year 1610 when Galileo first discovered craters, mountains, and valleys through a telescope. Subsequently, when more and more telescopic discoveries of lunar formations were made, it was necessary for the International Astronomical Union to prepare a systematic nomenclature of the lunar formations based on terrestrial names and names of eminent persons. The moon is probably one of the few bodies of the solar system where many of the surface features are preserved in the original form because of the virtual or total absence of erosion and weathering agents, though its surface undoubtedly disintegrated under the continuous impact of small meteorites and perhaps the action of cosmic radiations. Two major types of morphological formations can be noticed, viz. the (maria) sea having apparently smooth and level surfaces as seen at low magnification, and rough and roughand-broken ground usually in high or upland forms (terra). The mare or seas (an incorrect term because of its absolute dryness)

occupy about one-third of the visible surf ace of the moon, and on the basis of their general appearance can be classified into circular ones -such as Mare Imbrium, Crisium, and Serenitatus - and irregular ones-such as Mare Tranquillitatis, Fecunditatis, Nubium, and Oceanus Procellarium. On the whole, the floors of the maria are below that of their surroundings, and on macroscopic scale they seem to be fairly smooth between craters though actually they possess many features that become visible under magnification. In addition to ridges and wrinkles on the surface of the maria there are often clefts or cracks up to 5km in width and sometimes a few hundred kilometres long called rules or rils. Low domes are common in some maria. In general, ridges are formed along the coast of a mare though there are cases where ridges of a few hundred metres appear radiating in one direction or crossing each other on the ocean floor. Apart from these broad characteristics, the circular maria apparently possess no ghost craters. Among the lunar seas definite information is available on Imbrium (Sea of Rains), Crisium (Sea of Crises), Serenitatis (Sea of Serenity), Humboltianum (Humbolt Sea), Marginis (Border Sea), and Smythii (Seaof

46 / Geomorphology Smyth), all of circular shape, and Tranquüitatis (Sea of Calm), Fecunditatis (Sea of Plenty), and Nectaris (Sea of Nectar), of irregular shape. Mare Imbrium covering the NE quadrant of the moon is encircled by the Caucasus, Alps, Apennines, and Carpathian mountains and has a diameter of some 1100km. In fact, these mountain chains represent the broken and incomplete ramparts of a gigantic crater, the floor of which is now the bed of the sea. The western bay sea is known as Palus Putredinis (Marsh of Decay) and the eastern one as Sinus Iridum (Bay of Rainbow). The floor of the latter bay is 600m lower than that of the main sea and is of semi-circular shape with Jura Mountain and Laplace (1200m) and Heraclides (2700m) promontories on its northeast, north, and south respectively. Mare Serenitatis, another circular shaped sea, with a diameter of nearly 700km is bordered by the Taurus, Caucasus, and Haemus mountain chains on east, west, and south respectively. The floor of the sea is strewn with ridges, of which one on the western periphery of the sea winds from north to south and attains a height of 200m at places. Mare Crisium, the third important circular formation, located to the west of Imbrium, is on the fringe of the visible lunar hemisphere and has a diameter of 500m. The floor of the sea is strewn with ridges and craters, Picard being the most important with a diameter of 35km and walls of 1500m high. Mare Tranquillitatis, one of the prominent seas among the non-circular ones, has on its southeastern part a few mesas at some distance to give it the resemblance of a desert. Fecunditatis, another important irregular mare, has a dimension of 100km NS and 700km EW and has its floor strewn with ridges and craters. Mare Nectaris, yet another non-circular sea with somewhat square shape, has 300km a side and Pyrenees range on western shore. The other major lunar topographic unit is the rough and broken surface which covers an area of little less than two-thirds of the visible surface. The lunar uplands appear to have little appreciable slope. Even large crater walls or mountains have a gradient of less than 10° to the direction of the base. The foremost formations are the mountain chains, abounding in the southern hemisphere though

no two are identical. These lunar formations though named after the terrestrial mountains are not similar to them. Their craggy ridges, very high but not steep slopes, sometimes make them appear higher than the terrestrial mountains. The majority of the lunar mountains are grouped together to form mountain chains similar to terrestrial formations, but there are isolated mountains like Pico and Piton on the lunar plains. The Alps, Carpathians, Apennines, Caucasus, and Jura, all around the Mare Imbrium have relatively high peaks among which Mont Blanc (4000m) of Alps, and Wolf (3500m), Ampere (3000m), Huygeus (5500m), Hadley (4800m) of Apennines are worth mentioning. The Carpathians, southeast of Mare Imbrium, is a discontinuous low altitude range, while the Caucasus lying in between Mare Imbrium and Serenitatis, have extremely high peaks, the highest being some 6000m. Among the other lunar mountains, Doerfel range located in the uplands near the lunar south pole is the highest. It has peaks between 5000m and 9000m high. It will not be out of place to mention the 'Straight Wall,' a unique lunar cliff resulting from a fault: it is 120km long and 500m high with the western side of the floor higher than the eastern. Among the lunar formations the craters are the dominant feature on the uplands. Their occurrence on the uplands indicates that they preceded maria; further, their existence in the areas of maria, though largely destroyed or covered (yet existing in the form of ghost craters - faint circular outlines), is still discernible in the ocean floors. Discoveries of less common but distinct craters on the maria are of later date. The largest crater, Clavius, close to and slightly west of the moon's south pole, has a diameter of 227km. There are 4 other craters with diameters down to 200km, and 32 with diameters between 100 and 200km. The outlines of such large craters are polygonal with rounded corners rather than circular, which may be due to disintegration after their formation. The height of their walls is rarely more than 5km above the interior floor and generally a few hundred metres above their surroundings on the outside. The maximum slope of the interior side of the wall may be as much as 30° but generally less on the outside. Though

Géomorphologie I 47 the central floors of such wall plains (craters) are generally below the level of their outside surroundings, yet in many instances there is a peak or sometimes a group of peaks near the central floor rising to height of up to 2.5km, and in a few cases having smaller craters on the summit of central peaks. There are a few craters in the uplands which have raised interiors but no central peaks. The badly damaged walls of these craters suggest that they probably are the oldest lunar craters. Apart from these, there are a dozen or so ray craters, a unique formation on the lunar surface. Their characteristic feature is that streaks of light coloured material with insignificant thickness radiate from the crater for a distance of several hundred kilometres, apparently without interruption, across the lunar surface. Copernicus, Kepler, and Aristarchus are the conspicuous ray craters of the maria, while Tycho, near the south pole, is of the upland. Since the ray craters cut into other formations uninterrupted, they are believed to be the youngest lunar craters of appreciable size. This sketchy information about the morphology of lunar formations is not sufficient for any concrete theory on their evolution, yet it can be said that both volcanic processes

and meteoritic impacts are responsible for their structure. Abell, G., 1964 Exploration of the Universe (New York), 177-9. Alter, D., éd., 1968 Lunar Atlas (New York). Beiser, A., 1959 Our Earth (New York). Fesenkov, V.G., 1959 Some Considerations about the Primaeval State of Earth (London), 9-10. Glasstone, S., 1965 Source Book on the Space Sciences (New York). Hamblin, J., 1969 A close watch on the men and their prizes, Life 47 (2), 66-7. Jean, F.C., 1958 Man and His Physical Universe (New York). Kopal, Z., 1964 The Moon: Our nearest Celestial Neighbor (New York). Lascelles, R.G., 1964 Atlas of the Moon (MacMillan). Odishaw, H., 1967 The Earth in Space (New York). Schroth, C., 1970 Moon rock studies disclose new facts on lunar structure, American Reporter, Dec. 2* Wainwright, L., 1969 Apollo's great leap for the moon, Life 47 (3), 28-9.

P0125 Pediment formation in the Mojave Desert, California THEODORE M. OBERLANDER University of California, USA Few relief features have been investigated more indefatigably than the erosional pediments cut across crystalline rocks in the southwestern deserts of the United States. Yet the literature on desert pediments is conspicuously descriptive and speculative, demonstrating a surprising lack of progress in the understanding of this landform. Recent morphometric analyses by Mammerickx (1964) and Cooke (1970) are steps in a new direction, but their results show clearly that even less is understood about pediments than was hitherto realized. The same may be said of Melton's quantitative investigation (1965) of the boulder slopes rising above granitic pediments in Arizona. In the first instance, the concept of the pediment as a slope graded for transportation is made to appear doubtful, and in the second, the 'boulder controlled' or

'repose' slope is shown to be an illusion. Quantitative studies of the evolution of badland slopes and pediments by Schumm (1962, 1966) produced very concrete results but left unclear how findings related to the erosion of weakly consolidated material could be applied to landscapes developed in crystalline rock. All American investigators have approached the pediment landscape in terms of the existing morphogenetic regime. According to the most recent summary of the literature on pediments in the southwestern United States (Hadley 1967), the preponderant opinion is that pediments truncating hard rock originate as a concomitant of the backwearing of escarpments under arid or semiarid conditions, and serve as slopes of transportation for the fine products of weathering

48 / Geomorphology derived from the positive forms rising above them. The conspicuous slope break between pediments and surmounting residual masses in areas of granitic rock has been attributed to the bi-modal size distribution of granitic weathering products. The steep boulder-clad slopes rising above the pediments have been regarded as the angle of repose for blocks loosened by weathering along joints, and the pediment has been assumed to be the minimum slope allowing transportation of grus resulting from weathering in the boulder zone. The association of pediments with granitic rocks has been stressed and attributed to the peculiar size grading of granitic weathering products, suggesting that the presence of a well-defined piedmont angle is critical to recognition of the pediment landform. Rather more significant is the suggestion that the preferential development of pediments on granitic rocks must be the consequence of distinctly rapid parallel rectilinear slope recession in this material during the period of pediment morphogenesis. Schumm's studies in the South Dakota Badlands concluded that parallel slope retreat producing miniature pediments requires only that the slope foot not be a zone of deposition, this condition being met where erosion by slope wash is quantitatively more important than that by gravitational creep. This is true as a consequence of the effect of slope roughness on the velocity of water flows on hillslopes and pediments. My own investigations indicate that the pediments of the Mojave Desert are not a consequence of the present arid morphogenetic regime, but are relict forms inherited from prior non-desert landscapes. The evidence consists of fragments of Tertiary weathering profiles preserved under basaltic lavas having radiometric ages exceeding 8 m.y. The Tertiary land surfaces that can be reconstructed from such remnants were strongly oxidized and weathered to depths between 5 and 50m. They are distinguished by a brick-red soil developed across composite cut and fill surfaces surmounted by steep sided hills mantled with saprolite. A complete cover of vegetation of nondesert composition and type is implied, and has been documented by palynplogical analysis (Axelrod 1958). Slope retreat in this landscape was predicated upon the existence of a mantle of finely comminuted material over-

lying an active subsurface weathering front along which chemical decay of the subsurface rock kept pace with surface erosion. In several localities there is stratigraphie continuity between existing boulder slopes and the basal portions of soil profiles covered by volcanics dated radiometrieally at more than 8 m.y. Present boulder mantles thus appear to consist of former corestones originally formed within pre-Quaternary weathering profiles. Saprolite-covered pediments were an integral part of the Tertiary landscape. Surviving fragments of Tertiary lava flows found on the highest portions of existing pediments reveal that these pediments were fully developed before the onset of desert climatic conditions, and have been modified in the past several million years only by removal of soil and decayed rock inherited from the Tertiary morphogenetic regime. The height of lava mesas presently standing above pediment surfaces approximates the local thickness of the Tertiary weathering profiles. As could be expected in the case of erosion surfaces produced by backwearing, weathering below the Tertiary land surface was significantly deeper towards the distal portions of the individual pediments. During the late Pliocene and Quaternary, pediments and residuals alike have been largely denuded of their former regolith mantle. Cessation of subsurface decay and acceleration of surface erosion at the end of the Tertiary are presumed to be the consequence of gradual climatic desiccation following uplift of the Sierra Nevada, San Gabriel, and San Bernardino ranges. Axelrod has documented the progressive vegetative change produced by the rise of these barriers. The geomorphic consequences were to trigger erosion, stripping the hillslopes down to the irregular bedrock weathering front or to a veneer of corestones let down as a lag deposit, and to produce wholesale exposure of bedrock on the upper margins of the pediments. The latter has been interpreted by several observers as 'exhumation' of a rock-cut suballuvial bench. Examination of relict weathering profiles indicates clearly that the granitic pediments of the Mojave were not cut in rock, but were extended by erosion of regolith that was simultaneously being renewed by chemical decay in the subsurface. Trimming of projecting bedrock irregularities at the sur-

Géomorphologie I 49 face is occurring at present, but such projections have themselves been exposed by the stripping of pre-weathered material, and the total amount of erosion of outcrops by arid processes seems nowhere more than one or two meters. Even this doubtless includes removal of some material pre-weathered in the subsurface during the Tertiary. During their growth the Mojave pediments would not have terminated upward at the sharp piedmont angle seen in the present rocky landscape. The Tertiary landscape was mantled by saprolite under a cover of chaparral or grass, with no contrast in the size or nature of particles on retreating hillslopes and expanding peripheral surfaces. The hillslope profiles were convexo-concave, as are many present boulder slopes. Field experiments by Emmett (1970) have shown how . such slopes are produced by the transition from laminar to mixed to turbulent overland flow, which results in the retreat of soilcovered slopes at a constant angle. The distinctiveness of granite in this respect would seem to lie in the high mass permeability of its decomposition products, which, under an aggressive weathering regime, produce a mantle of irregular depth that is gradational into solid rock. Lacking a well-defined zone of moisture concentration, such a mantle has a relatively low susceptibility to gravitational transfer, and thereby to diminution in slope angle. The findings of Schumm and Emmett thus appear relevant to the morphogenesis of

Mojave Desert pediments - the presence of hillsides erodable by slope wash being crucial to backwearing and the pedimentation process, and the existence of wash slopes in a soil-covered landscape during the period of pediment expansion lending support to Schumm's conclusions regarding slope development. Axelrod, D.I., 1958 Evolution of the Madro-Tertiary geofiora, Bot. Rev. 24, 433-509. Cooke, R.U., 1970 Morphometric analysis of pediments and associated landforms in the western Mojave Desert, California, Am.J.Sci.269 (1), 26-38. Emmett, Wm. W., 1970 The hydraulics of overland flow on hillslopes, usos prof, paper 662-A. Hadley, R.F., 1967 Pediments and pediment forming processes, /. Geol.Educ. 15, 83-9. Mammerickx, J., 1964 Quantitative observations on pediments in the Mojave and Sonoran deserts, Am. J. Sc/. 262, 417-35. Melton, M.A., 1965 Debris-covered hillslopes of the southern Arizona Desert, /. Geol. 73, 715-29. Schumm, S.A., 1962 Erosion on miniature pediments in Badlands National Monument, so, GSA Bull. 73, 719-24. Schumm, S.A., 1966 The development and evolution of hillslopes, /. Geol. Educ. 14 (3), 9.8-104.

P0126 Hillslope evolution on the Nsukka Plateau of eastern Nigeria GODFREY EZEDIASO KiNGSLEY OFOMATA University of Nigeria, Nigeria Nsukka Plateau is one of the major components of the Nsukka-Okigwe cuesta, a prominent geomorphological unit in the East Central State of Nigeria. I have previously described the two groups of landforms on the plateau: residual hills and dry valleys (Ofomata 1967). The present paper is concerned with an analysis of the manner of hillslope evolution of the residuals. The plateau is developed over a sedimentary sequence, the oldest formation of which dates from the Upper Cretaceous. The general level of the plateau (about 330m asl) is underlain by the False-bedded sandstones

(Upper Senonian in age) above which stand residuals of Upper Coal Measures (Upper Senonian to Paleocene), which often rise to over 200m above this general level. The False-bedded sandstones consist mainly of unconsolidated, poorly-sorted, and strongly cross-bedded rocks, white to pale grey in colour, with thin bands of white shale occurring at intervals, especially towards the base. The formation is sandy in texture and highly friable. The more resistant Upper Coal Measures rest conformably on the Falsebedded sandstones and consist of sandstones, sandy shale and carbonaceous shale. Gritty

50 / Geomorphology and pebbly bands are locally present in the sandstones and the shales contain many concretions, bands, and stringers of marcasite and siderite. The plateau lies in a transitional zone between the rainforest region to the south and the savanna region of the north of Nigeria. It has a long rainy season (April to October), and a relatively short dry season which accounts for only 11 per cent of the total annual rainfall of 1655mm, spread over a yearly average of 113 rain days (figures are for Nsukka town). The rains come in the form of torrential downpours which not only aid the incidence of soil erosion, but also make it possible for large amounts of water to be retained on the surface of the ground often for the duration of the rain. Rainwater is evacuated mainly in the form of sheet wash across the characteristic concavo-convex slopes of the hills and the undulating surface of the plateau. Widespread sheet erosion is thus common. Rainwater concentration occurs in the dry valleys as well as in the first order thalwegs and vshaped gullies incised on the slopes. A reasonable quantity of the rainwater also infiltrates into the underlying geologic formations, and is responsible for the weathering to great depth of these formations to produce a thick mantle of weathered materials. Weathering also leads to the precipitation of iron oxides, the absolute and/or relative accumulation of which, within the weathered rock materials, leads to the formation of highly resistant iron-stone concretions (duricrusts), which are responsible for the flattopped nature of most of the residuals (Ofomata 1967, 7). Typical slope profiles on the plateau are associated with the major types of residuals in the area. The conical and the domey residuals, both of which present generally rounded outlines, have predominantly concave slopes, with smooth convex upper portions (profile i, Fig. 1 ). The flat-topped types (profile n) have concave slopes with steep upper portions and gentle lower ends which decrease gradually to adjoining plains. The steeper upper portions here coincide with the resistant caps of the iron-stone concretions already mentioned. The cuesta-like residuals (profile m) combine the concave slopes of the flat-topped residuals with the concavoconvex slopes of the more rounded varieties, while the ridges have rectilinear steep upper

Fig. 1. Typical profiles: Nsukka Plateau, Nigeria. slopes which merge with the general level of the plateau through definite breaks-in-slope. In general, therefore, the residuals on the Nsukka Plateau are characterized by gently concavo-convex slope forms, with the concave wash element in the profile usually dominant. These wash elements are formed over weathered materials, and the author shares the view that they cannot be referred to as pediments, except where the term simply denotes 'any slope displaying a hydraulic curve resulting from surface wash processes ...' (Thomas 1962). The major processes of hillslope erosion in the area are a combination of weathering and rain water run-off. Where weathering leads to the formation of iron-stone concretions, the general tendency is for the hills to present a flat-topped summit which acts as a resistant cap rock and, to a large extent, controls the manner of slope development and evolution. The duricrust cap is dismantled, rather slowly, through a progressive slope retreat, following the collapse of overhanging blocks as softer materials are washed out from below. Where, on the other hand, there is no such resistant cap formation, the hills develop by progressive down-wearing of slopes following the action of sheet wash on pre-weathered materials. In some subtle way the familiar controversy over slope retreat or slope decline reappears here. But field observation helps to resolve the issue more satisfactorily than usual. It is concluded that slope retreat is a dominant process in the development of the flat-topped residuals until the resistant cap rock is removed. After the complete removal of the resistant cap, the summit becomes rounded and, from then on, the process of evolution is through progressive slope decline. By far the most predominant factor controlling slope evolution in the area is lithology. This point was also stressed by Usoroh (1968, 191 ) in his study of the Mamu Basin

Géomorphologie / 51 which covers parts of the Nsukka area. Consequently, certain slope values have come to be associated tentatively with particular geologic formations. In general, slopes developed over iron-stone concretions, especially where they occur as horizontal caps over weaker formations, range from 40 to 55°, with a few nearly vertical, but with a mean value of 45 °. Over the Upper Coal Measures they range from 10 to 45°, with a mean value of 18°. On the False-bedded sandstones, the values lie between 2 and 25°, with a mean value of 6°. Space limitation has not allowed for a fuller discussion of the main points raised in

this paper. In spite of this limitation, however, a few conclusions emerge, the most significant being the overriding influence of lithology and climate on hillslope evolution in the area studied. Ofomata, G.E.K., 1967 Landforms on the Nsukka Plateau of eastern Nigeria, Nigerian Geog. J. 10, 3-9. Thomas, M.F., 1962 On the approach to landform studies in Nigeria, Nigerian Geog. J. 4, 87-100. Usoroh, E.J., 1968 Erosional development of the Mamu River Basin, eastern Nigeria. MSc thesis, U. Ibadan.

P0127 Classification of alluvial plains based on geomorphological characteristics MASAHIKO OYA University of Waseda, Japan Alluvial plains may be classified by the following criteria: agents of deposition, place of deposition, stage of deposition, and deposition material. Agents of deposition fall into two main categories: ( 1 ) the plain was formed by one main river and its tributaries, and (2) the plain was formed by a complex of rivers. Category 1 may be subdivided into (a) the force of the deposition of the main stream predominates over that of the tributaries, and (b) the force of the deposition of tributaries predominates over that of the main stream. Plains with characteristics indicated by ( 1 ) are categorized according to the place of deposition: (w) indicates a wide plain and (N) indicates a narrow plain. In the case of a wide plain, alluvial fans will be formed; on narrow plains they cannot form. The depositional materials of alluvial plains are divided into two categories: sand, silt, clay with gravels (G), and sand, silt, clay without gravels (s). If there is gravel alluvial fans will be formed; if there is no gravel alluvial fans will not be formed. There are also two categories for stages of

deposition: (D) deposition is proceeding; (E) deposition is proceeding in some parts and erosion is proceeding in other parts. By combination of the above mentioned categories, almost all the alluvial plains in Japan and Asia can be classified as the following types: DG-1(1)W

DG-1(1)N

DS-1(1)W

DS-1(1)N

ES-1(1)W

ES-1(1)N

DS-1(2)N EG~1(2)N

DS-2(3) EG-2(3)

EG-1(1)W

DG-Í(2)N

ES-1(2)N

EG-1(1)N DG-2(3)

ES-2(3)

DG-1(2)W

DS~1(2)W

EO-1(2)W ES-1(2)W

DG~2(4)

DS-2(4) EG-2(4) ES-2(4)

Many alluvial plains in Japan belong to the types EG-! ( 1 ) w, EG-! (2)w, EG-! (2)N, EG-2(3), EG-2(4). Almost all plains of Korea belong to the types DS-! ( 1 ) w, DS1 (1 )N. The alluvial plains of tropical areas in East Asia belong to the types DS-2(3), ES-1 ( 1 )w and ES-2(4). This regional distribution of types of alluvial plains can be explained by regional differences in features of crustal movements and climate.

P0128 Tectonic control of morphology on the Canadian interior plains GEORGE F. OZORAY Research Council of A Iberia, Canada The linear elements of morphology on the interior plains of Canada exhibit a definite geometric, grid-like pattern. This is especially

true of the drainage network. River valleys show rigidly straight reaches, orthogonal turns, giant pseudo-meanders, and opposite-

52 / Geomorptiology

Fig. 1. Morphotectonic megalineaments, Canadian interior plains and their neighbourhood. Legend: (1) rivers, lakes, sea coast; (2) morphotectonic trends, including proved faults; (3) ridge of the Rocky Mountains; (4) boundary of geologic-physiographic regions; (5) provincial boundary; (6) inter-

national boundary; c = Cordilleran region; IP = Interior Plains; s = Shield; RM = Rocky Mountains; MC = Missouri Coteau; WL = Wollaston Lake; GR = Geikie River; Q = Queen Charlotte Islands fault. The base map is of a zenithal equal area projection.

directed river stretches along the same lineal trend. Lake shores, ridges, escarpments, geologic boundaries, magnetic and gravity anomaly areas, dry or muskeg-filled depressions, and groundwater discharge features often follow the same pattern, suggesting a common origin. Two systems of linear features, each comprising two mutually nearly perpendicular directions (Fig. 1), are recognized (Ozoray 1970): 1. the linearity of one system is either parallel to (Rocky Mountains trend) or

orthogonal to (Lake Athabasca trend) the Rocky Mountains of west-central Alberta, that is, in the central part of the investigated area, approximately NW-SE and NE-SW respectively; 2. the other system is characterized by N-s and E-W linear features (meridional and latitudinal trends). Some features of the morphologic regularity are a result of glacial control or caused by deposition (beach ridges). Others that cannot be explained in such ways are discussed here. The preglacial drainage network

Géomorphologie / 53 also exhibits a geometric pattern, similar to the present one, although the river courses are different (Meneley, Christiansen, and Kupsch 1957; Stalker 1961; Farvolden 1963; McCrossan and Glaister, eds., 1964, Fig. 14-2). A linear controlling factor must exist, influencing morphogenesis for long distances (1000km magnitude) and through areas of different geologic, orographie, and climatic character. It is tectonism. The tectonic character of some morphologic lines has been proved (Douglas, éd., 1970, maps 125U and 1255A; Sikabonyi 1957). The tectonic origin of the NW-SE and NE-SW trending lineament system of the Missouri Coteau in Saskatchewan is well evidenced by a brilliant lineament analysis (Kupsch and Wild 1958). The main preCretaceous and Cretaceous structural trends of southern Alberta, found by computer analysis of deep test-hole data (Robinson, Charlesworth, and Ellis 1969, Figs. 18-24), match with the ones deduced from surface morphology (Ozoray 1970). Structural effects often start as the revitalization of basement tectonism. The morphologic details are controlled by the local joint-systems of bedrock and depend on local causes; often they may be influenced by nontectonic structures, too. They both are recognizable on airphotos through the driftveneer as lineaments (Blanchet 1957). The regional trends most likely originate in the crystalline basement, as the same morphotectonic trends characterize both the interior plains and parts of the Shield (even partly the Rocky Mountains). The NE-SW Wollaston Lake-Geikie River line in north Saskatchewan is a good example of the Lake Athabasca trend on the Shield. The N54°E trending, Lake Athabasca-Peace River line (at the town of Peace River) even crosses parts of both the Shield and of the interior plains. The investigated morphologic pattern is young (late Tertiary and/or Quaternary) and is controlled by more or less contemporary tectonism (even Holocene: Bay rock 1968, unpublished manuscript). However, this young tectonism must have been affected by reactivation of older trends, particularly in the pre-Cambrian basement. On the Shield some NE-trending faults are probably preCambrian (Green 1958, 7). The meridional and latitudinal trends appear to be secondary

and younger than the other two, but they are not necessarily very young themselves. The discussed system of megalineaments characterizes a vast area (it even extends into the western United States and Mexico). The evaluation of its tectonic character would require a separate study, but some preliminary remarks might be given : 1. The trend-systems may represent two sets of Mohr shear-fractures. 2. They conform to the Rocky Mountains, but in their primary form they may be older than the Cordilleran system. 3. Systems of tectonically controlled megalineaments, having linear features in the same directions as the ones discussed, are widespread. They have been described in Great Britain, Australia, Africa, and South America (Hills 1963, 460, and Figs, xrv-28 and xiv29s ), and in Hungary. Northeastern strikes of tectonic lineaments on the epi-Hercynian platform of the southern USSR have been described, although they have not been evaluated morphologically (Mirchink et al. 1970). The strike of many oceanic features, including transcurrent and transform faults, also agrees with the mentioned trends (van Bemmelen 1965, Figs. 1 and 5; Wilson 1965, Figs. 7-9). 4. This global occurrence points to common global controlling factors. It can be connected with the direction of the convection magmacurrents and with such events of continental drift as rotation of the drifting North American plate. R. Bibby; R. Green; O. Tokarsky; J. Tóth; F.L. Copeland; Council of Alberta. Bemmelen, R.W. van, 1965 Mega-undations as cause of continental drift, Géologie en Mijnbouw 44, 320-33. Blanchet, P.H., 1957 Development of fracture analysis as exploration method, Bull. Am. Assoc. Pet. Geol 41, 1748-59. Douglas, R.J.W., éd., 1970 Geology and economic minerals of Canada, 2 vols. Gebl. Surv. Can. Econ, Geol. Rept. 1. Farvolden, R.N., 1963 Bedrock channels of southern Alberta, in Res. Conn. Alberta Bull 12, 63-75. Green, R., 1958 Precambrian basement features in northern Alberta, ibid. 3. Hills, E.S., 1963 Elements of Structural Geology (London).

54 / Geomorphology Kupsch, W.O., and J. Wild, 1958 Lineaments in Avonlea area, Saskatchewan, Bull. Am. Assoc. Pet. Geol. 42, 127-34. McCrossan, R.G., and R.P. Glaister, eds., 1964 Geological history of western Canada. Alberta Soc. Pet. Geol. (Calgary). Meneley, W.A., E.A. Christiansen, and W.O. Kupsch, 1957 Preglacial Missouri River in Saskatchewan, /. Geol. 65, 441-7. Mirchink, M.F., N.A. Krylov, Yu. A. Romanov, and I.O. Averbukh, 1970 Northeastern strikes of tectonic lineaments on the epi-Hercynian platform of the southern USSR, trans, from Doklady AkademiiNauk S.S.S.R., v. 192, 395-8. Ozoray, G.F., 1970 Structural control of morphology in Alberta, Progr. and Abstr.

,

Ann. Meet. Geol. Assoc. Can. (Winnipeg), 40-1. Robinson, I.E., H.A.K. Charlesworth, and M.J. Ellis, 1969 Structural analysis using spatial filtering in interior plains of southcentral Alberta, Bull. Am. Assoc. Pet. Geol. 53, 2341-67. Sikabonyi, L.A., 1957 Major tectonic trends in the prairie region of Canada, /. Alberta Soc. Pet. Geol. 5, 23-8. Stalker, A. MacS., 1961 Buried valleys in central and southern Alberta. GeoL Surv. Can. Paper 60-32. Wilson, J.T., 1965 A new class of faults and their bearing on continental drift, Nature 207, 343-7.

P0129 The main types of landslides MARTON PECsi Hungarian Academy of Sciences 1. TYPES OF LANDSLIDES

A landslide in the strict sense is a rapid downslope mass movement of bedrock connected with the existence of a slip plane. It is slip plane development that distinguishes landslides from the other types of mass movements. The various types of landslides have not yet been classified in the geologic-geomprphological literature. Distinctions between types should be made not only from the genetic point of view, i.e. on the basis of their particular role in morphogenesis, but also in terms of engineering-geological practice. Among the landslides in the broader sense, phenomena provoking a slow, long-term deformation of the slopes, but not connected with a definite slip plane, are rather frequent. In this case the sliding of the material follows a system of minor creep planes. Similarly a special process is 'block-movement' which, however, is no typical slide, but a slow creep of basalt, etc., blocks down a sloping clay substrate. Movements other than the landslides in the strict sense are falls of rock masses on overturned slopes; the fall and plastic flow of overmoistened materials down the slopethe so-called 'earth flow.' No doubt these processes may occur associated with landslides. That is why authors of some textbooks in engineering geology have classified almost all processes of gravitational mass movements

as belonging to the scope of landslides or to sliding movements taken in the broader sense. It is primarily the three-dimensional geometrical types of the slip plane that the present genetic classification of landslides s.str. has relied upon, with a consideration of the morpho-lithogenic and hydrometeorologic characteristics. The necessity of the formation of a slip plane is the most characteristic feature of a landslide. This distinguishes it from the mere fall of materials on the slopes and provokes a periodical repetition of the phenomenon. 2. CONCISE GEOMORPHOLOGICAL

CHARACTERIZATION

I. Rockslides and slope-slides. In this case the surface of sliding is steep, being situated high on the mountainside, and the volume of the sliding rock mass is catastrophically huge. This mass performs a quick sliding movement starting from the source of the slide down to the foot-slope over a path of several hundred metres comprising passages characterized by downfalls. If the bedrock is covered by a detrital clayey mantle of varying thickness which, as a result of moistening, will slide downslope on the bedrock surface, it is a slope-slide. On the contrary, a rock-slide is the rapid movement of both the weathered mantle and the bedrock, a movement along the slip plane down to the foot of the mountain slope. A rock-slide may come about on the one hand

Géomorphologie I 55 in imconsolidated rocks, along bedding planes preformed during sedimentation; on the other hand, in solid rocks along fault planes, slip planes controlled and preindicated by weathering. The difference between the two consists in that a rock-slide is more localized, occurring rarely in the same place, whereas a slope-slide on weathering-carpeted steep slopes is repeated periodically. In many cases, the material leaves the source of the slope-slide with jerks rolling like an avalanche down the slope. In the case of a slope of 40 to 45°, the overmoistened cover of weathering products will yield to shearing stresses and begin to slide and fall along the resulting shear plane. After heavy rains the weathered rock-cover is saturated with water down to the bedrock and on slope stretches steeper than 45° the elastic solid material will lose its cohesive strength and the intergranular friction will fall below the critical limiting value. II. Slip of stratae. A carpet-like slip of strata takes place where a clayey, sedimentary basement dips slightly downslope, being carpeted by groundwater-bearing sandy clays or products of weathering. In such cases a long, but discontinuous slip plane will form which usually has the same angle as the angle of dip of the underclay. Tapping the waterbearing strata may reduce, and even eliminate, the danger of movement. The movement of the mass is periodical, setting in during the wet season. An environment endangered by sliding is readily revealed on the slope by geomorphological evidence such as waste mounds and tensile fissures, cracks, etc. in. The slice-slides represent one of the common and typical cases of landslides. Prerequisites for their development are an impervious, clayey basement overlain by a thick, (10~50m) and for the most part, permeable sequence which usually ends in the escarpment of a river bank. On the impervious clay the base of the overburden is periodically moistened to such an extent that its inner cohesive strength decreases. As a result, the overburden will be fractured along slices parallel to the escarpment and partly sheared off of it. Still resting on the escarpment along fibrous joints, the slices will then be further displaced. During this process, pressure and moisture increase in the basal strata. At a critical value of these characteristics the basal layers will lose their resistance to compression and be suddenly sheared off. A few narrow

earth slices will then be torn and slide away on the surface of the clay as on a preindicated and moistened slip plane (Fig. 1, not published here). iv. The slump may be treated as a classic representative of rotational landslides. This process takes place on a steep slope of clays and loams. The semi-cylindrical slip plane is formed within the clay body along a surface of rupture. The slip plane may terminate exactly at the base of the slope, above it or below it (undercut). The shear plane in the clay is not a geologically controlled surface. v. Slump-earth flows have arched centres and slip planes, being formed in the steepest upper part of the slope. The rock-waste slides rhythmically in a narrow channel. Afterwards, the waters of newly-formed springs will overmoisten the slipped mass which periodically becomes plastic and 'flows' down, to finally accumulate in the form of a tongue at the base of the slope. Viewed from above, this type shows a glacier-like pattern. Morpholithological prerequisites for its formation are a marked slope covered by rather thick (5-20m), detrital products of weathering or by lopssic-loamy sediments. Morphohydrological prerequisites are a steep slope where groundwaters or artesian waters of the permeable strata cannot cross the overburden to reach the surface directly in the form of springs. A channel-pathed landslide is a complex phenomenon comprising both 'slump' and 'earth flow.' The movement may last for years with temporary pauses, the motions being connected for the most part with the humid seasons. It would be desirable to supplement observations and measurements with the theories and methods of geomechanics and engineering geology and to continue work in the domain of genetic classifications by undertaking complex fundamental research. Kézdi, A., 1970 Soil mechanics n (Budapest). Pécsi, M., 1971 The main types of landslides. Foldr. Kozl. 2-3, 125-43. Varnes, D.I., 1958 Landslide Types and Processes 'Landslide and Engineering Practice/ Hwy. Res. Bd. Special Report 29 (Washington, DC). Zaruba, O., and V. Mencl, 1969 Landslides and Their Control (Prague), 205.

56 / Geomorphology P0130 Solifluction rates in the Ruby Range, Yukon Territory: a preliminary report LARRY w. PRICE Portland State University, USA A project for measuring long-term rates of mass wasting in a periglacial environment was initiated during the summers of 1967 and 1968. This paper presents the initial results. The Ruby Range is located in the Yukon Territory at 61 °23'N lat. and 138° 13'w long., about 215km NW of Whitehorse. The range is composed of granitic intrusions. The higher peaks average 2130m. There is a distinct flattening of the upland surface, which may be an old erosion surface (Bostock 1948, 72). The area has been glaciated, however, and is quite rugged. The Gulf of Alaska lies only 230km to the west but the towering St Elias Mountains (highest peak, Mt Logan 6050m) intervene and eliminate direct marine influence, so the Ruby Range has a continental climatic regime (Taylor-Barge 1969). Tree line occurs at about 1200m. The area has a subarctic alpine tundra environment. Intensive investigations were carried out on four slopes formed by the confluence of two ridges. These four slopes, facing SE, sw, E, and N, all have the same rock type (coarsegrained biotite granodiorite), similar gradients (14° to 18°) and elevations (16751980m). The slope environments differ considerably, however, because of exposure to solar radiation and meltwater from late snow melt. Solifluction lobes and tundra vegetation are best developed on the SE facing slope, and sequentially less so on the E, sw, and N facing slopes (Price 1971). Two principal techniques were used to measure mass-wasting: (1 ) insertion of polyethylene tubes vertically to permafrost and monitoring subsequent deformation with a strain gauge probe, (2) painting in situ surface rocks and establishing measurement grids. Measurement of slope movement with polyethylene tubes and strain gauges was largely pioneered by Williams (1957; 1962a and b; 1966): and his methods have been followed with a few modifications. Polyethylene tubes 13mm ID and 16mm OD were inserted in auger holes to permafrost on different slopes at depths of 1-1.5m. The tubes were kept straight in the ground by a steel rod fitted snugly inside; after repacking the soil, the rod was removed. Subsequent move-

ment of the tube was measured by a strain gauge probe consisting of two strain gauges fastened on either side of a thin strip of spring steel and encased in silastic rubber. As the probe bent following the curvature of the tube, the flow of electricity through the circuit varied and was measured on a Wheatstone bridge. The probe was calibrated and field readings were taken at 3.75cm intervals down the tubes several times during both summers. Thirty tube sites were established during the two field seasons but movement at only a few of these has been analysed (Price 1970). The southeast facing slope is 900m wide, has a gradient of 14°, and is the alb of a glacial trough bounded upslope by a rocky ridge 150-200m high. Snow patches accumulate in the lee of the ridge providing meltwater through the summer. Solifluction lobes are well developed averaging l-2m in height but the larger ones reach 5-7m. The tundra vegetation is well developed, and sharply delineated communities occur in horizontal bands across the step-like form of the lobes (Price 1971 ). Four measurement tubes were analysed, all being on lobe treads. Surface movement at these sites ranged from 1321mm/yr. and averaged 16mm/yr. Movement decreased gradually to an average depth of 22cm. Two of the sites were excavated at the end of the second field season to check the accuracy of the probe readings and were found to coincide closely. The east facing slope is similar in form to the SE facing slope but is only 300m wide with a gradient of 15° and a 100m high ridge upslope. Snow patches provide meltwater until mid-July and Solifluction lobes are well developed, ranging from 1-3 m in height. Several tube measurement sites were established on the E facing slope but not nearly enough to allow the collection of data for two field seasons. Painted rock measurement sites were established, however, and are discussed later in this paper. The southwest is the driest of the study slopes and Solifluction lobes are virtually absent. One reason for its xeric quality is the prevailing northwesterly winter wind (TaylorBarge 1969, 214) which sweeps the slope clean of snow. It is 220m wide with an aver-

Géomorphologie I 57 TABLE 1. Average annual surface movement rates across solifluction lobes (mm) Slope faces

Tread

SE E N

16 14 12

Riser

Near lee (within 4.5m below riser)

Far lee (4.5-30.5m below riser)

6 7 6

3 3 4

5 2 5

age gradient of 17°, and the surface consists of 1.5m or more of unconsolidated weathered detritus. The plant cover is of the tundra fell field type. One movement tube was put at midslope on a dry and thinly vegetated site. Annual surface movement was 5mm and extended to a depth of 10cm. Due to the lack of moisture, movement here is probably more a result of frost creep than solifluction (Washburn 1967, 49-50). The north facing slope is the steepest (18°) and narrowest (180m) of the study slopes. The sun strikes the slope at a very low angle, and mountains to the north and west prevent the sun from shining on the slope during the night hours. As a result, snow patches remain until early July and vegetation is discontinuous and sparsely developed. Solifluction lobes occur but are not well developed, reaching a maximum height of only 1.5m and occurring singly rather than coalesced as on the SE and E facing slopes. One tube was installed on the lower part of the N facing slope where there were no solifluction lobes. The site was very moist, and sparse patches of moss and sedge occurred amidst bare ground. Surface movement from 9 August 1967 to 9 August 1968 was 2mm and extended to a depth of 15cm. Average annual rates of surface movement on the different slopes as revealed by the mass wasting tubes are as follows : SE facing slope, 16mm; E facing slope, —; sw facing slope, 5mm; N facing slope, 2mm. These results take on more significance when combined with the painted-rock measurements. Although the mass wasting tubes give a good idea about rate of movement on and within the lobes, they do not indicate how or at what rate the risers move. To measure this important characteristic over 200 rocks were painted in the lobe risers on all slopes except the sw where suitable sites were lacking. Rocks in the riser were selected for measurement and base rocks downslope for control

(Price and Alexander 1971). Ideally riser rocks were in situ, so whatever movement occurred would accurately represent surface movement but in some cases where no such rocks were available others were placed on the surface. Analysing the rock movement data, it was clear that rates of movement varied a great deal from the tread to the riser, to the area below the riser. Therefore the data were analysed on the basis of these four areas: tread; riser; near lee, and far lee (Table 1 ). The figures support the evidence presented by the basic form of solifluction lobes, that movement is more rapid on the tread than below the lobe. They establish an order of magnitude in rate of movement from one segment of the lobe to another. For example, on the SE facing slope the tread is moving five times faster than the near lee area. The riser is moving more slowly than the tread and movement increases again farther downslope in the far lee area (Table 1 ). The ratio between the tread and near lee area on the east facing slope is 4 to 1 and on the north it is 3 to 1. This correlates nicely with the relative development of solifluction lobes on each of these slopes since they are best developed on the SE, followed by the E, and then the N facing slope. Arctic Institute of North America; Icefield Ranges Research Project; Jim Sij; Charles Volk; Nancy Price; J. Peter Johnson; Richard Ragle; Walter Wood; Mel vin Marcus; Charles S. Alexander. Bostock, H.S., 1948 Physiography of the Canadian Cordillera with special reference to the area north of the fifty-fifth parallel, GeoL Soc. Can. Memoir 247. Price, Larry W., 1970 Morphology and ecology of solifluction lobe developmentRuby Range, Yukon Territory. PhD dissertation, U. Illinois.

58 / Geomorphology Price, Larry W., 1971 Vegetation, microtopography, and depth of active layer on different exposures in subarctic alpine tundra, Ecology 52. Price, Larry W., and C.S. Alexander, 1971 Methods of measuring mass wasting: review and critique, Proc. Assoc. Am. Geog. 3, 135-9. Taylor-Barge, Bea, 1969 The summer climate of the St Elias Mountain region. Arctic Institute of North America, Res. Paper no 53. Washburn, A.L., 1967 Instrumental observations of mass wasting in the Mesters Vig District, Northeast Greenland, Meddelelser om Grfinland, Bd 166, no 4. Williams, P.J., 1957 The direct recording of

solifluction movements, Am. J. Sci. 255, 705-14. Williams, P.J., 1962a An apparatus for investigation of the distribution of movement with depth in shallow soil layers, Division of Building Research, National Research Council, Ottawa, Canada, Building note no 39. Williams, P.J., 1962b Quantitative investigations of soil movement in frozen ground phenomen, Biuletyn Peryglacjalny, no 11, 353-62. Williams, P.J., 1966 Downslope soil movement at a sub-arctic location with regard to variations with depth, Can. Geofech. J. 3 (4), 191-203.

P0131 Outline for a genetic classification of major types of sinkholes and related karst depressions JAMES F. QUINLAN McMaster University, Canada Much of the literature concerning the origin of sinkholes (dolines) and related karst depressions is seemingly contradictory. An author may assert or prove that sinkholes in a particular area are developed by a given process. He may or may not try to extend his proof as a general model, but subsequent authors, in a different locale, have tended to refute the first model and offer a second or a third, also as a general model. The purpose of this note is to show that-as in cave development - there is no general model for sinkhole development that is universally applicable. This paper lists the various genetic types of sinkholes and shows their relation to one another; a future paper will discuss the rock mechanics, soil mechanics, and hydrology of sinkhole development (Quinlan and Jennings, in prep.). The major processes involved in the development of various types of sinkholes are: collapse, subsidence, solution, lixiviation, nivation, suffosion, and consolidation. It is the extent to which any of these processes may be dominant that determines what type of sinkhole may form. Ultimately, of course, all sinkholes are a consequence of solution. Collapse differs from subsidence in that the latter process occurs slowly rather than catastrophically. In a broad general sense, the American

term sinkhole and the European term doline are synonymous; doline, however, is the international term for any type of moderately sized karstic depression, be it a.pit or a pan. The following classification is based on major morphogenetic process and the major surface or near-surface material (rock or soil) that is lowered. Where appropriate, subclassification is on the basis of where the processes operate, type of recharge to the doline, fill, and topographic expression. The types of dolines are named and a description is given of how the processes cause their development. Dolines may also be classified according to size, shape, hydrologie function, and position in a landscape, but these criteria are not considered herein. i. COLLAPSE DOLINE. Produced by collapse of rock, soil, or other clastic material. A. ROCK-COLLAPSE DOLINE. Produced by the collapse, and sometimes subsidence* of rock and overlying soil. 1. Shallow collapse doline. Collapse is into a cavity at shallow depth and is as a result of: (a) denudation of the ground surface and loss of stability of the roof of a cave, (b) weakening of a rock arch by solution, or (c) stoping of a cave roof after a and b. Usually the collapse is through karstifiable rock that is similar to that of the subjacent cavity. Such

Géomorphologie I 59 dolines commonly allow access to an underlying cave. 2. Deep-seated collapse doline. Collapse and subsidence is a result of stoping into a cavity at great depth. These processes operate because a stable dome is not developed; stoping continues and, in some of the breccia pipes so formed, it may be aided by solution by artesian water. Usually the collapse is through non-karstifiable rock or rock that is different from that of the subjacent cavity. Some rare deep-seated collapse dolines have developed in response to shrinkage caused by oxidation (by groundwater) of sulphide replacement ore bodies in limestone and by accompanying solution of cavities in the bedrock. B. SOIL-COLLAPSE DOLINE. Collapse OÍ SOU

(or other clastic material) may occur because of: (a) lowering of a water table, with consequent surcharge and collapse of soil arches, (b) raising of a water table, with consequent ¿ loss of cohesion of soil, (c) flooding of the ground surface, with consequent cohesion loss, suffosion, and surcharge on the soil over any subjacent void, or (d) seismic shock from earthquakes, blasting, or mining subsidence. Soil-collapse dolines are the most common type of doline in urban areas. They commonly are anthropogenic because their development is triggered by or accelerated by groundwater pumpage and other kinds of alteration of the natural environment by man. ii. SOLUTION DOLINE. Produced by lowering of the ground surface because of subaerial, subsoil, or interstratal solution of bedrock or sediment. A. SUBAERIAL SOLUTION DOLINE. Produced

by subaerial solution of bedrock. Some may have a very thin mantle of soil. Subaerial solution dolines are transitional to subsoil dolines, q.v. 1. Rock doline, solution pit, or karren well, depending upon size or shape. Corroded by rain water, snowmelt, or icemelt, that falls directly or flows as rills. Any of these may develop beneath a temporary cover of snow or water. 2. Ponor doline, sink, or swallow hole. (All 3 terms are equivalent and commonly used.) Corroded and abraded by water from a surface stream. If the doline sometimes reverses flow and functions as spring, it is an estavelle. 3. Sub glacial doline. Corroded and abraded subglacially.

B. SUBSOIL DOLINE (or SUB-ALLUVIUM DOLINE, etc.). Produced by subsoil (or suballuvium) solution of bedrock and concomitant lowering of the ground surface; a high concentration of soil-CQ2 favours relatively rapid corrosion of carbonate bedrock. c. INTERSTRATAL DOLINE. Produced by interstratal solution of relatively soluble bedrock and concomitant subsidence of overlying relatively insoluble rock or sediment; beds dip inward from all sides. Subclassified according to whether there is a fill, and if one is present, according to its nature. Most interstratal dolines do not have topographic expression as a depression. 1. Subsidence doline. No significant fill; forms a topographic depression. 2. Structural sink. With fill consisting of relatively unweathered bedrock; rarely has topographic expression as a depression. 3. Geological organ. With fill consisting of unconsolidated sediment; rarely has topographic expression as a depression. Depending upon the degree of weathering of the fill, geological organs are transitional both to structural sinks and subsoil karren. D. LIXIVIATION DOLINE. Produced by lowering of the ground surface because of local solution of the relatively soluble particles within sediment (e.g. calcite within loess) and consequent consolidation of the sediment. This lixiviation process is also known as 'chemical suffosion.' in. SOLUTION-NIVATION DOLINE. Subaerial solution doline that is enlarged and deepened also by spalling induced by repeated cycles of freeze and thaw; characteristic of karsts in cold climates. Nivation may be the initial cause of some dolines. iv. SUFFUSION DOLINE. Produced by washing of soil (or other sediment) into an underlying cave or solutionally enlarged cracks; the topographic form may be entirely within the soil cover, the subjacent bedrock surface being essentially flat. v. SLUMPAGE DOLINE. Produced by solution of rock fragments and bedrock, plus downwashing (suffosion) of soil or other sediment Into cracks, plus repeated step-like collapse of soil; loss of cohesion by wetted soil causes local collapse of small arches over voids produced by solution and suffosion. VI. COMPACTION-SUBSIDENCE DEPRESSION.

Produced by lowering of the water table in soil (or other sediment), with consequent re-

60 / Geomorphology duction of void space and differential settlement of the ground surface as water drains from soil. The occurrence of some collapse dolines, suffosion dolines, slumpage dolines, and compaction-subsidence depressions can be accelerated or triggered by earthquakes, other seismic activity, or faulting. Some might be triggered by mining subsidence. Any doline, no matter how it initially formed, is usually further developed by subaerial or subsoil solution of bedrock. Accordingly, the dolines classified above are endmembers in several genetic continua. It is recognition and understanding of the endmembers, however, that leads to an understanding of the others. I fully realize that the origin of many sinkholes is not always obvious. But it may become more so after excavation of some and after a study has been made of the surface and subsurface geology. There are many other types of topographic

depressions but unless they are at least partially a result of solution, i.e. at least partially karstic, they are not sinkholes. The above classification does not include several types of polygenetic depressions that probably are only in part karstic — mardelles, Carolina bays, and some of the high plains depressions found in the midwestern USA - or depressions produced by mining subsidence or the discharge of industrial wastes. A comprehensive discussion of the various types of dolines, and their international literature, is given in my doctoral dissertation (Quinlan 1972). Quinlan, J.F., 1972 Karst, pseudokarst, and dolines: classification and a review. PhD thesis, U. Texas, at Austin. Quinlan, J.F., and I.E. Jennings Major causes of sinkhole development (ms in k: preparation).

P0132 Anomalous behaviour of model rivers

S.A. SCHUMM, R.S. PARKER, R.G. SHEPHERD, D.E. EDGAR, and M.P. MOSLEY

Colorado State University, USA

During the erosional development of a landscape an interruption of the normal progress of erosion will leave its imprint. For example, valley-fill deposits and river terraces must reflect significant changes of base level, climate, or hydrologie regimen. Thus, an erosional or depositional modification of the landscape is readily explained by climatic, tectonic, or eustatic influences. One needs to recall, however, that in hydraulics certain threshold values exist above which significant changes in the character of fluid flow occur (e.g. critical values of Reynolds and Froude numbers). For example, with a progressive increase in velocity and depth of flow in an alluvial channel, marked alterations of bed forms occur at critical values of stream power or Froude number (Simons and Richardson 1966). In addition, threshold values of tractive force and shear occur at which sediment movement begins. Therefore, the geomorphologist should expect that at least some secondary features of the landscape may reflect the effect of thresholds, variations in erosional efficiency, and even a complex geo-

morphic response to a simple change in an independent variable. We use three examples from our model studies on drainage pattern development and river morphology to illustrate these possibilities. The research is supported by grants from the National Science Foundation and Army Research Office-Durham. 1. A series of experiments were performed in a SOm-long and Sin-wide flume to study the effect of increased gradient on channel morphology. At a constant discharge of 0.004cms (0.15cfs) three channel patterns developed, as slope was increased. The increase of the slope of the surface on which the channel flowed was, of course, accompanied by an increase in sediment load, velocity of flow, and other hydraulic variables. It was determined that, as valley slope was increased from 0.001 to 0.02 at the same discharge, two significant alterations of channel pattern occurred. At very low slopes the channel remained straight, but at a slope greater than 0.002 a meandering thalweg developed, and at slopes greater than 0.015 a

Géomorphologie I 61 braided pattern developed. In this series of experiments threshold values of slope were detected at which major changes of channel patterns occurred. One may readily visualize similar, abrupt, and significant alterations of channel morphology in response to slope changes resulting from hydrologie changes of the Quaternary period. Preliminary results from a continuation of the above-described experiments suggest that not only are threshold values of slope significant, but at a given slope there are threshold values of discharge above which major changes of channel pattern occur. In these continuing experiments the occurrence of abrupt changes of channel morphology are now expected. 2. A very different type of channel response was observed during experiments involving channel incision. A trapezoidal channel was cut into cohesive materials, and a constant discharge was introduced into the model channel. It was assumed that scour would deepen the channel, and the objective of the experiment was to study variations in the scour patterns along the length of the channel. However, the channel did not deepen over its entire width; rather a narrow and deep inner channel formed. Such channels in bedrock have been observed and logically explained by erosion at low flow. This inner channel, however, was the result of variations of velocity and tractive force on the floor of the channel during bankfull discharge. 3. Another type of discontinuity developed in a model study of drainage network evolu-

tion. In a 9m by 15m box, which was filled to a depth of 1.5m with erodible sediment, a drainage pattern developed as simulated precipitation was delivered at intensities ranging from 1.2cm to 8cm per hour. The experiment was designed so that the base level of the system could be lowered and the response of the system to lowered base level documented. As might be expected, lowering of the base level caused incision of the main channel. Initially the channel scoured to produce a relatively low gradient. However, as tributaries were rejuvenated and bank caving occurred upstream, the sediment load of the channel increased to a point where deposition occurred, and an alluvial valley-fill was deposited by a braided stream. Later, as the system stabilized at a constant base level, sediment loads were reduced, and a narrower and deeper channel formed. Thus, a single lowering of base level produced incision, deposition, and renewed incision, as the entire drainage system adjusted to altered base level. During this series of events precipitation intensity and run-off were constant. In each of three very different studies unexpected responses of the model system occurred. In each case, if evidence for similar changes had been noted in the field, probably a very different interpretation of the sequence of events that produced the feature would have resulted. Simons, D.B., and E.V. Richardson, 1966 Resistance to flow in alluvial channels, us Geol. Survey Prof. Paper 422-J.

P0133 Erosion surfaces in the lower Chambal Valley, India H.s. SHARMA University of Rajasthan, India The Lower Chambal Valley lies between 24°45'-26°45'N and 75°20'-79°20'E (Fig. 1, diagrams not published here). It is a geotectonic zone; the younger Vindhyans were faulted against the oldei* Gwalior formations at some time in the Cretaceous period. The general elevation of the valley near Kota is about 250m to 270m above mean sea level while it decreases to 150m near the confluence with the Yamuna. The maximum elevation of 650m is attained in the NW border hills (Fig. 2). Three morphometric techniques were em-

ployed for the recognition of erosion surfaces and subsequently field studies were made. The altimetric frequency histogram (Fig. 3) shows (1) that more than 1200 spotheights in the Lower Chambal Valley and adjoining areas fall between 180m to 300m above mean sea level and that there is greater frequency near the lower elevations; (2) that about 800 summits also occur between 360m-450m and a good part of them tend to be near the lower limits; and that there are (3 ) a few summits above 480m. Superimposed profiles (Fig. 4) give the

62 / Geomorpholagy same results as the altimetric frequency histogram. The two sets of superimposed profiles give a bird's eye view of various erosion surfaces in the valley. Moreover, it is evident that summit plains occur at three heights, viz. 180m to 305 m, 360m to 450m, and 480m to 610m. The longitudinal profiles (Fig. 5) of the Chambal and its tributaries support the results of these two methods. Proceeding downstream along the Chambal there is one bench or knick point at about 375m near the boundary of Malwa plateau and the Triple scarps. Thereafter, at the terminus of the Singoli scarp near Rawatbhata the second bench occurs. Further downstream, a last knick point at a height of 225m is found where the Chambal emerges from the gorge section near the town of Kota. All these three inflections in the course of the Chambal bring out the fact that in the development of the normal cycle of erosion during the geological past the Chambal was interrupted three times, giving rise to successive rejuvenation of various erosion surf aces and knick points at three heights. The following erosion surfaces may be recognised in the area: 1. Pre-Cretaceous erosion surface between 480m and 600m; 2. Mid-Miocene erosion surface between 360m and 450m; 3. Pleistocene erosion surf ace between 180m and 305m. The greater part of the valley is covered by the Vindhyans and recent unconsolidated alluvium (Fig. 6). The Gwalior formations occur only to the northwest of the Great Boundary Fault in a narrow and elongated belt. Generally, the summits over 490m occur in the Kaimur and Bhander sandstones of the upper Vindhyan system and the Gwalior quartzites. Spot heights around 360m occur entirely in the Vindhyan sandstones and shales. There is a greater frequency of spot heights below 305m; these are generally confined to the flood plain bluffs and alluvial knolls from the city of Kota to the confluence of the Chambal with the Yamuna. It has been observed during the field study that the summits of the Vindhyan formations are characteristically elongated, flat, and continuous while the tops of the Gwaliors are rounded and isolated.

PRE-CRETACEOUS EROSION SURFACE (480M-

600M) The relics of the pre-Cretaceous erosion surface are restricted to the bevelled surfaces of the Akoda hill, the Satur plateau, the Palki mesa of the Kota plateau, and the Ranthambhor syncline (Fig. 7). The bevelled surfaces of Satur and Palki are composed of Bhander sandstones. On the contrary, the Akoda hill and Ranthambhor syncline are comprised of the Kaimur sandstone and older Gwalior quartzites respectively. The uniformity of spot heights above 480m over the Vindhyans and the Gwaliors, irrespective of their lithology and structure, suggests partial peneplanation, representing an interrupted cycle of erosion. The structural and geomorphic evidence leads to the conclusion that the pre-Cretaceous erosion surface of the central Vindhyan Basin was subsequently uplifted and faulted along the NW fringe against the pre-Cretaceous surface of the Aravalli mountain prior to the outpouring of the Deccan lava. Choubey (1967) found some evidences of a pre-Cretaceous erosion surface at 379m515m in the Sagar, Damoh, Jabalpur, and Hoshangabad districts of Madhya Pradesh, while Dubey (1968) also observed erosion remnants of a contemporary surface at around 546m-63 8m on the Rews plateau. In the Lower Chambal Valley this surface is at about 480m-600m. Its occurrence in three different parts of the Vindhyan Basin suggests a partial peneplanation of the region. MID-MIOCENE SURFACE (360M-450M)

The termination of the Cretaceous cycle may be marked by the sequence of events in the Himalayan region, tectonic movement along the northwestern part of the Vindhyan Basin, and the subsequent eruption of the basaltic lavas in western and central parts of the peninsular India. The mid-Miocene erosion surface stands about 30m to 45m below the pre-Cretaceous surface. The remnants of this surface have a similar distribution to those of the pre-Cretaceous surface, though they are more extensive and widespread. The relicts of this cycle may commonly be observed on the northern and southwestern parts of the valley (Fig. 7). The perfect bevelled surfaces of the Bundi-Sawai Madhopur hills, the cuestas and mesas of the Kota plateau, the top of

Géomorphologie I 63 the Chambal scarp, and the Mukandawara hills represent the remnants of such a surface. PLEISTOCENE SURFACE (18ÛM-305M)

The termination of the Miocene cycle by cymatogenic upwarping may be dated back from the sequence of events in the Himalayan region and in the southern part of India. The studies of Krishnan (1953) and Radhakrishnan (1967) reveal that peninsular India was disturbed by epeirogenic movements during the mid-Miocene period. The Pleistocene surface in the valley is the most extensive of the erosion surfaces and gives a general uniformity of height and gently undulating character to the riverine plain of the Chambal. The remnants of this surface occur as extensive valley-side benches merging gradually with the lower elevations of the mid-Miocene erosion surface of 360m and above.

It is concluded from the above study that the various erosion surfaces of the region are the result of the intermittent upheavals of the Himalayas. Choubey, V.D., 1967 A study of erosion surfaces in the Sagar, Damoh, Jabalpur, and Narsingpur districts of Madhya Pradesh, Proc. Geomorph. Studies India, 16470. Dubey, R.S., 1968 Erosion surfaces in the Rewa plateau, Madhya Pradesh, 21 IGC, I (New Delhi). Krishnan, M.S., 1953 The structural and tectonic history of India, Memoirs Geol. Survey India, \. 81. Radhakrishnan, B.P., 1967 The western Ghats of the Indian peninsula, Proc. Geomorph. Studies India, 4-14.

P0134 Rock glaciers on Aquarius Plateau, Utah JOHN FORD SHRODER University of Nebraska at Omaha, USA The Aquarius and Table Cliffs Plateaus in southern Utah are located in the eastern part of the High Plateaus section on the west edge of the Colorado Plateau province. The area lies between lat. 37°37'30" and 30°15'N, and long. 111 ° 15' and 112°w, and measures about 48km east to west, and 68km north to south. The maximum relief is over 1800m, with a maximum elevation of 3453m in the north on Boulder Mountain and about 1525m in the canyons to the southeast. Annual precipitation ranges from about 76cm on the plateau, to about 13cm in the canyons. The climate is humid microthermal (Db and DC) in the highlands and mid-latitude steppe (BSk) in the lower elevations. The plateaus generally have a precipitous bounding scarp developed in Tertiary basalt. During the Pleistocene, part of flat-topped Boulder Mountain was covered by a small ice-cap glacier which had over seven outlet glaciers flowing down over the sides (Flint and Denny 1958). Several other smaller cirque glaciers also occurred on the periphery (Smith, Huff, Hinrichs, and Luedke 1963). Elsewhere the plateau surfaces were unglaciated but subjected to intensive periglacial action. Extensive dry valleys, block fields,

and patterned ground were produced. Mass wasting throughout the Quaternary, because of underlying unstable Mesozoic and Tertiary sedimentary rocks, produced vast areas of debris-flows, landslip blocks, talus, protalus lobes, block lobes, and rock glaciers below the scarps. This activity produced a prominent landslip bench with generally gentle slope. Steep slopes, which are largely the result of headward stream erosion, occur below the bench. Some of the valleys of the former outlet glaciers on Boulder Mountain contain a landslip bench and lower landslip-bench slope, which are both moraine-covered. The term rock glacier, as used in this report, applies to accumulations of unsorted rock debris which have been transported a considerable distance; which show pronounced ridges, furrows, cones, pits, steep fronts, and other features described by Wahrhaftig and Cox (1959), and which are caused by slow interstitial-ice or ice-cored flow. About 109 rock glaciers occur in the area; 30 were investigated in the field, and the remaining 79 were analysed from maps and aerial photographs. LITHOLOGY Rock glaciers investigated in the

64 / Geomorphology field are over 99 per cent basalt. Reconnaissance of most of the remainder indicates the ubiquity of this lithology, which is a result of its strong jointing and predominantly bloeky fracture. Wahrhaftig and Cox (1959, p. 415) noted that this is a common phenomenon of rock-glacier clasts because large interconnected voids are formed in which ice can accumulate. SOURCE In the southern unglaciated area the landslip scarp produced 33 rock glaciers and the landslip-bench slope generated 20, In some places the landslip bench is narrow (a few hundred metres) and rock glaciers, which develop from talus, block-slides, and rock-falls at the main landslip scarp, flow across the landslip bench and down the lower slope. Those rock glaciers which originate on the landslip-bench slope do not have obvious source rock but it is probable that buried landslip blocks contribute significantly. On Boulder Mountain no rock glaciers occur on the landslip-bench slope. This is probably due to the presence there of fine elastics from till and from long-term weathering which infill voids and exclude ice. Rock glaciers on the landslip-bench slope in the southern area may be due to the absence of till and to the greater youth (hence less weathering) of the landslip bench there. Source material for 43 rock glaciers on Boulder Mountain came from glacially overridden cliffs and cirque headwalls. Extensive faulting on the west side of that area caused scarps which generated 13 rock glaciers. TYPE According to the classification of Wahrhaftig and Cox (1959, p. 389), only two types of rock glaciers occur in this area, the tongue-shaped type which is longer than it is wide, and the lobate, which is wider than it is long. About 64 rock glaciers are lobate; the remaining 45 are tongue-shaped. The predominance of short lobate types is due to several causes at the sources: (1) gentle slopes below glacial scarps and cirque headwalls (35 rock glaciers) ; (2) gentle or inward-dipping slopes opposite fault scarps (13 rock glaciers) ; (3) landslip blocks which impede flow; and (4) relatively moderate climate caused by low elevation and latitude. Twenty-one of the tongue-shaped rock glaciers occur on the steep landslip-bench slope and 15 occur below landslip scarps. ELEVATION On Boulder Mountain the aver-

age elevation of rock-glacier heads is 3145m (range 3060m-3240m). Average elevation of toes there is 3080m (range 2940m-3165m). Average elevation of heads in the remainder of the area is 3030m (range 2808m-3150m); toes average 2900m (range 2712m-3045m). Rock glaciers thus tend to be higher on Boulder Mountain, which is a reflection of the greater altitude of the contributing basalt there. SLOPE EXPOSURE Direction of flow and slope exposure are essentially the same for the rock glaciers. The direction of flow of each was plotted on the eight octants of the compass. Fourteen flowed N-NE; 29 NE-E; 17 E-SE; 6 SE-S; 3 s-sw; 9 sw-w; 21 W-NW; and 10 NW-N.

Slope exposures of the contributing basalt caprock were sampled at 0.5km intervals around the plateaus. Forty occur N-NE; 44 NE-E; 54 E-SE; 62 SE-S; 40 s-sw; 27 sw-w; 30 W-NW; 21 NW-N. The large number of occurrences in the NE-E and W-NW octants is a reflection of the north-south linearity of the plateau with a large number of east- and west-facing slopes. North and south slope exposures are consequently more uncommon, which explains the overall deficiency of rock glaciers with these directions. Nevertheless, the northern octants have many more rock glaciers than the southern, in spite of the proportionately larger number of slopes there, which is, of course, because the interstitial ice and contributing snowbanks are less likely to melt in a northern exposure. GRADIENT Average gradients of field-investigated rock glaciers originating at glacial scarps are 15° (range 100-20°), those at fault scarps are 16° (range 14°-19° ), those at landslip scarps are 17° (range 11 °-22° ), and those at landslip bench slopes are 20° (range 13°-27° ). Average gradients are remarkably similar and no statistical significance can be attached to these values at present. Further work is in progress. HISTORY OF MOVEMENT Study of tree rings, lichens, and aerial photos indicates detectable modern movement on only about four rock glaciers in the south. Lobes (Wahrhaftig and Cox 1959, p. 424), compound fronts one above another, with increasing soil, vegetation cover, and lichen diameters in the lowermost; and multi-lobate fronts where reactiva-

Géomorphologie I 65 tion caused diversion around a stagnant front: all attest to at least three stades of rock glacier movement, undated at the present. University of Nebraska at Omaha, Senate Research Grant; Colin Fallat; Frank Stehno. Flint, R.F., and C.S. Denny, 1958 Quaternary geology of Boulder Mountain,

Aquarius Plateau, Utah, us Geol. Survey Bull. 1061-D. Smith, J.F., Jr., L. C. Huff, E.N. Hinrichs, and R.G. Luedke, 1963 Geology of the Capitol Reef area, Wayne and Garfield counties, Utah, us Geol. Survey Prof. Paper 363. Wahrhaftig, C., and A. Cox, 1959 Rock glaciers in the Alaska Range, Bull. Geol. Soc. Am. 70, 383-436.

P0135 Some general regularities in development of relief of the Ukrainian SSR Territory IGOR SOKOLOVSKY Academy of Sciences of the Ukrainian SSR, USSR Recent extensive research in the Ukrainian SSR enables one to approach in a new way the study of general regularities in relief development, particularly the problems of interaction of endogenous and exogenous processes, the formation of levelling planes and river terraces, regularities in development of recent relief-forming processes, etc. Recent neotectonic investigations have confirmed the presence of relations between the relief and structure, the latter being considered in its dynamics, and have provided a quantitative characteristic for some regions as well. Studies of structure and neotectonic movements in modern relief ('morphostructural analysis'), were given graphical expression in a map of morphostructures. Morphostructures of three different orders are distinguished in it. First Order morphostructures are major units within the geostructural regions of the pre-Cambrian platform, Scythian platform, and the adjacent Alpine geosynclines. The socle elevated plain of the Ukrainian crystalline shield, plains in the south-western margin of the Russian platform, Dnieper-Donets, and Black Sea area depressions, the peneplaned elevated plain of the Donets folded structure are the morphostructures of the Russian (Pre-Cambrian) platform. Beddedstage, bedded, and structural-denudation plains in the southern part of the PrutDniester interfluve and Steppe Crimea correspond to the Scythian platform (on the Baikal and Gertsinian basement) ; foldedclumpy ridges of the Carpathians, the accumulation terrace plain of the Forecarpathian foredeep, structural-denudation elevated

plain of the Upper Tissa depression and accumulation lowland of the Chop-Mukachevo depression correspond to the Carpathian orogen; mountain structures and plains correspond to the folded system of the Crimea and Indol-Kuban trough. Second Order morphostructures are units within the first order morphostructures. They are big blocks undergoing neotectonic movements differentiated in time and space. Thus, within the socle elevated plain of the Ukrainian crystalline shield one can observe a number of benches which correspond to big blocks of the shield; similarly, within the plain in the southwestern margin of the Russian platform, the Dnieper-Donets and Black Sea depressions and the Scythian platform plains, particular blocks are differentiated by their sharp mobility during the NeogeneQuaternary period. Within mountain structures there are ridges and depressions corresponding to definite structural subdivisions of the Carpathian orogen, etc. Third Order morphostructures are smaller surface areas corresponding to local structures, e.g. brachy-anticline structures of the Dnieper-Donets depression, the Forecarpathian foredeep external zone and other regions. Study of the tectonic activity of these structures promotes the solution of problems of oil and gas content. Within the Ukrainian crystalline shield, faults and faulted zones dividing separate blocks are the main types of local structures studied by morphostructural methods. Crustal movements are manifested on the earth's crust surface as vertical movements of blocks within the limits of shields only

66 / Geomorphology where the sedimentary cover is of insignificant thickness or is entirely absent. In troughs on the surface there are manifested not only block movements of the crystalline base but also dynamics of the sedimentary cover, which is of particularly great importance when plastic saline masses are present in sedimentary strata. Studies of the effect of exogenous processes upon formation of the modern relief make it possible to distinguish surface areas with characteristic complexes of morphosculpture forms : (a) complexes of chiefly accumulation morphosculptures of glacial and periglacial regions of the Oka and Dnieper Glaciations as well as of alluvial and deltaic morphosculptures; (b) complexes of accumulation-denudation morphosculptures of extraglacial regions; (c) complexes of chiefly denudation morphosculptures. The principles of finer subdivision of the accumulation and denudation morphosculptural areas are different: accumulation morphosculptures are subdivided according to accumulation conditions, denudation morphosculptures according to the leading denudation processes, the structure and composition of rocks on which they are formed. Simultaneous analysis of morphostructure and morphosculpture of the Ukrainian SSR territory shows that the development of the morphostructures exerted a decisive effect on formation of the relief sculpture forms. They determined location of the boundary for continental glaciations, the periglacial zone, the location of a shore line for the Black Sea and Sea of Azov, the distribution of elevated and subsided areas, etc. A study of levelling planes is of importance to detect the relation between morphostructure and morphosculpture as well as to

study origin and age of the watershed reliefs. Levelling planes in the Ukrainian SSR territory have formed since the Paleozoic, but the earliest regionally distributed levelling planes playing an ascertained role in modern relief were formed only in the Late Mesozoic era. There are Late Mesozoic, Paleogene, Miocene, Pliocene, and Quaternary levelling planes. According to genetic types, these are subdivided into: (a) marine accumulation planes formed with sufficiently complete compensation for local tectonic movement, by accumulation and denudation; (b) alluvial-deltaic accumulation levelling planes formed under conditions of stable sea level with a minor predominance of subsidence; (c) continental levelling planes - peneplains, accumulation-denudation and chiefly accumulation planes. On the whole, accumulation levelling planes are most characteristic of the plain parts of the Ukrainian SSR. Denudation planes are found only within the boundaries of the Ukrainian Shield and Donets Basin where levelling of the tectonic-created relief took place. The second important peculiarity of the levelling planes of the Ukrainian SSR is their stage arrangement, the superposition of one plane on another. This feature complicates their mapping but at the same time it offers new opportunities for establishing the history of the relief development and for using these findings in the search for minerals. Essential differences are determined in quantity, heights of terraces, and alluvium structure within separate geostructures. This enables the results of such research to be used for detection of quantitative characteristic of neotectonic movements.

P0136 Rock control in coastal erosion at Arasaki, Miura Peninsula, Japan TAKASUKE SUZUKI and KEN'ICHI TAKAHASHI Chuo University, Japan TSUGUO SUNAMURA University of Tokyo, Japan On the Arasaki coast, southwest of Tokyo, the uplifted wave-cut benches of a height of 1 to 3m above msl are well developed. Mean tidal range is about 1.6m. This coast is com-

posed of steeply inclined Miocene strata, which are of a rhythmic alternation of mainly mudstone and scoriaceous tuff layers 5 to 400cm in thickness, striking about N 60°E,

Géomorphologie I 67 diagonal to the general trend of the coastline, and dipping about 70°s. A number of weathering joints are observed in the superficial part of mudstone layers, but not in tuff. Frequency of weathering joints in mudstone increases with increasing altitude of ground surface on the benches, but tends to be constant on the landward coastal cliffs. Micro-topographic feature on the benches and their coastal cliffs is characterized by a washboard-like relief. On the benches, the ridges and the furrows of the relief, which attain 5 to 300cm in amplitude, are always composed of tuff and mudstone layers, respectively. In contrast, on the cliff faces higher than about 3m above msl, such correspondence of micro-topography to geology is reversed, the ridges being composed of mudstones, attaining about 30cm in amplitude. In the zone between msl and about 1.5m below msl, the tuff is also slightly furrowed and the amplitude of both layers is smaller, not exceeding about 10cm. The zone below 1.5m below msl was not investigated. Mean physical properties of respective specimens of mudstone and tuff from this coast are 2.6 and 2.6 in apparent specific gravity; 1.6 and 1.7 in dry bulk weight (g/cm 3 ) and 39.3 and 34.0 in porosity (per cent). Mean mechanical properties are respectively 235 and 170 in compressive strength (kg/cm2, at the state that specimen's water content (w) is 0 per cent) ; 43 and 15 in tensile strength (kg/cm2, at w=0 per cent) ; 74 and 30 in shear strength (kg/cm2, at w=0 per cent) ; 2.0 and 1.3 in elastic wave velosity (km/sec, at w=0 per cent) ; and 36 and 75 in abrasion loss by Los Angeles Machine (per cent/500r, at w=0 per cent). These mechanical strengths decrease with increasing water content of specimens. For example, compressive strength under a fully saturated condition is about one-tenth of that under a fully oven-dried condition. In order to examine the resistance of rock specimens to wind and wave abrasion, the following experiments were carried out. Air (A), sand (s), and water (w), and their combinations (A+S, A+W, and A+s+w) were perpendicularly jetted on the flat surface of rock specimens at a speed of 16m/sec during 15 minutes from the jet gun which was set 5cm distance from the surface of rock specimens. Volume (cm3) of abrasion

loss of mudstone and tuff, for example, were 3.4 and 6.0 in case of the combination of A+S, and 1.1 and 1.8 in case of A+S+W, respectively. When the specimens of mudstone and tuff were put into water for more than 10 days, their responses to water absorption were as follows, respectively: 25. 9 and 17.7 in maximum water content (per cent) ; 0.16 and 0.52 in absorption rate (water content per cent/ minute); 0.15 to 0.40 and 0.0 to 0.05 in swelling strain (per cent) due to water absorption at fully saturated state; and also 400 to 1200 and 0 to 50 in swelling pressure (g/cm 2 ). When fully dried specimens of mudstone were put into water, and their water content attained about 6 per cent, the specimens began to crack into many minor flakes. Such phenomena, however, were not observed in case of tuff. The cracking of mudstone was probably caused by tension. The above-described results of some mechanical tests prove that, as compared with tuff, mudstone has higher resistence to wear, impaction, compression, tension, and shear failure, while mudstone itself very easily disintegrates owing to swelling and shrinking due to wetting and drying. According to these laboratory experimental results and also to the exogenetic agencies acting on this coast, the aforementioned difference of relative credibility of mudstone and tuff layers in the various altitudes of ground surface is explained as follows: On the cliff faces higher than the limit of spray of waves at the highest high tide, that is, about 3m above msl, wind abrasion is the most important erosional process, hence tuff layers are much eroded as compared with mudstone. In the zone between msl and about 3m above msl, the alternation of wetting and drying affects the rock surface. Owing to the repeated expansion and shrinkage due to such wetting and drying, mudstone is more easily cracked than tuff, and weathering joints are developed much more in the uppermost part of mudstone layers than that of tuff. Many minor flakes of mudstone, separated by such weathering joints, are washed away by storm waves and waves at high tide. As a result, mudstone layers are more eroded than tuff, and the most remarkable washboard-like relief results on the wave-cut benches. In the zone between msl

68 / Geomorphology and about 1.5m below msl, such weathering joints are not developed in mudstone, because of little effect of the repeated wetting and drying; hence tuff has a higher erodibility with wave action. It is stressed in conclusion that relative erodibility of rocks must be examined in relation to the acting geomorphic agencies, because rocks with various physico-mechanical properties behave themselves in various ways under different geomorphic

conditions, as already suggested by Prof. Yatsu. Suzuki, T. et al., 1970 Rock mechanics on the formation of washboard-like relief on wave-cut benches at Arasaki, Miura peninsula, Japan, Geographical Review of Japan, 43, 211-22 (in Japanese with English abstract). Yatsu, E., 1966 Rock Control in Geomorphology (Tokyo) (inEnglish).

P0137 Landslide damage in mountain areas in Tanzania and its control P.H. TEMPLE University of Dar es Salaam, Tanzania Landslide damage in some of the mountain areas of Tanzania is of considerable significance. It particularly affects steep-angled soil, debris, and regolith-mantled slopes in areas of high rainfall which have been recently deforested to extend the area available for cultivation. Areas where damage has been serious include the Uluguru mountains and the Rungwe area; areas where such damage has been recorded include the Usambara mountains, Kilimanjaro, and the Southern Highlands around Iringa. Probably because of the rapid rates of weathering and vegetation recolonization of slide-affected zones in such tropical mountains, this process of soil erosion and mass transfer has been largely overlooked. On 23 February 1970 a densely settled area ( 188 sq km) of 75 sq km in the western Uluguru mountains centred upon Mgeta (7°2's; 37°3-f'E) was devasted when over 10cm of rain fell in under 3 hours. Almost 1000 landslides, mostly small debris slides and mudflows, with some bottle-flows, were recorded in this area, while flooding, river erosion, and deposition of flood debris, partly supplied by landslides, contributed to the damage both within the storm area and over a wide area beyond it (Temple and Rapp in press). Damage was severe to (a) cultivated land and growing crops and to (b) roads, bridges, and culverts, causing consequent disruptions of communications. Six people were killed, 9 houses destroyed, and 20 goats killed. Damage was also sustained by neighbouring engineering works.

Some 1570 fields (shambas averaging 0.3 hectares in area) or approximately 500 hectares of cultivated cropland were partially destroyed; 13.6 per cent of the households of the area were directly affected. The storm occurred just before the harvest of the main staple crop, maize. It was estimated that in financial terms this harvest loss amounted to over 40,0005. The landslides also destroyed and removed a significant proportion of the best and deepest soils in the area, leaving behind barren scars which will take many years to recover. Damage to roads prevented vehicular access to Mgeta for 6 weeks and to Kienzema for over 9 weeks. Much of this damage was caused by flood waters which destroyed bridges and embankments, but a large amount was caused by the release of landslides from oversteep, artificial roadcuts. Repairs to roads (many of them very temporary) in the western Ulugurus resulting from this single storm cost 507,5005. The disruption of vehicular communications over the period after the storm necessitated the transport of all goods into and out of the area as head loads. This area normally exports considerable quantities of fresh vegetables to the urban markets of Morogoro and Dar es Salaam; during March and April 1970 40,000 kilos of vegetables, conservatively valued at 13,0005, rotted for lack of transport. The reduced quantities carried out by head load cost consumers an extra 42,0005 above normal prices and producers untold expenditure of effort. Flood waters generated by this storm damaged a water intake and the engineering

Géomorphologie I 69 works of an irrigation scheme on the mountain fringe, for which repair costs are estimated at 10,000s. An approximate addition of some of the direct costs of this single event can thus be placed at over 620,000* ($90,000). This figure is over half the total sum allocated in colonial times for a 10-year comprehensive rehabilitation scheme for the whole Ulugurus (Temple 1971). The catastrophe of 1971 is not an isolated event. Damage of comparable magnitude was experienced on the night of 23-4 March 1969 south of Kienzema and in 1968 in the Matombo area in the eastern Ulugurus. The recurrence interval of such events in one locality may be under 10 years. On the night of 9-10 April 1955 a fairly densely settled area (58 sq km) of 70-80 sq km on the southern slopes of Rungwe centred upon Itagata (9°14's; 33°39'E) was devastated when 53cm of rain fell during the night. No count was made of the total number of landslides nor was any inventory made of the extent of damage. Haldemann (1956) noted that several people had been killed, houses destroyed, and a large number of fields wasted. A further 12cm of rain fell on 5 May, causing a landslide of such magnitude that it temporarily blocked the main river of the area. The size of area damaged, the nature of the geomorphological processes recorded, and the nature of the area affected show strong affinities with the event at Mgeta. There are no written reports covering comparable events elsewhere. Observations by this writer indicate that numerous landslide scars are present in the northwestern Usambara mountains and in the Rombo and Old Moshi areas of the slopes of Kilimanjaro. In the Usambaras the areas affected by sliding appear to be zones where annual crops or pasture dominate the land use pattern. The same is true of the slide-affected zones on Kilimanjaro. In this context it should be noted that in the north Pare mountains where tree crops form a major element in the land use (Thomas 1970) no important landslides have been observed. In the Ulugurus less than 1 per cent of the slides recorded in 1970 affected wooded or forested slopes while 94 per cent had occurred on cultivated fields or fallow or grassland.

There is no feasible and economic means as yet of controlling rainfall. Slope angles may be changed by bench terracing; this was tried and proved unacceptable in the Uluguru and Usambara mountains. Furthermore it is technically unsound in such areas. The only feasible means of restricting landslide damage is by the modification of current land use practices to restrict rather than prevent damage, for the areas presently experiencing severest hazard are agriculturally valuable and often densely settled. Haldemann (1956) concluded from his studies in Rungwe that remedial measures normally recommended from a geological viewpoint (mainly drainage) would be economically unjustifiable as the costs would be out of all proportion to the value of the land so protected. Furthermore he concluded that reafforestation, which would effectively control soil erosion, was only possible to a limited extent without disrupting the agricultural economy of such areas. These conclusions appear to have a general validity. Conservation measures, to be acceptable, must bring (a) a demonstrable advantage to individual farmers in the short run and (b) long-term advantages to the whole community. Past measures considered only the latter point and were rightly rejected (Cliffe 1970; Temple 1971). Measures should be soundly researched before implementation; otherwise they are counterproductive in effect, actually accelerating the wastage of land resources (Temple 1971). The most feasible measures would be (a) to maintain existing reserved forest, (b) to encourage tree planting in lines below ridge crests, along road cuts and stream sides, and (c) to encourage a change towards perennial tree crops and away from annual cropping. Cliffe, L.R., 1970 Nationalism and the reaction to enforced agricultural change in Tanganyika during the colonial period, Taamulil, 3-15. Haldemann, E.G., 1956 Recent landslide phenomena in the Rungwe volcanic areas, Tanganyika, Tanganyika Noies Rec. 45, 3-14. Temple, P.H., 1971 Conservation policies in the Uluguru mountains. BRALUP Res. Pap. Temple, P.H., and A. Rapp, in press Land-

70 / Geomorphology slides in the Mgeta area, western Uluguru mountains; morphological effects of sudden heavy rainfall, Geog. Ann. 54A.

Thomas, I.D., 1970 Some notes on population and land use in the north Pare mountains, BRALUP Res. Notes 9,1.

P0138 Litho-control of morphometry in the Upper Gondwana terrain of central India V.K. VERMA University of Delhi, India Studies of the Gondwana formations, a continental suite of sedimentary rocks, have attracted the attention of many earth scientists for almost a century. However, there is insufficient geomorphological data on them, particularly from the morphometric analysis standpoint. The present paper embodies the results of investigations on morphometry as lithofunctions in the Upper Gondwana terrain of central India covered by the Survey of India one inch topographic sheet numbers 55 j/6 and 55j/7 (in parts) delineated by latitudes 22°25'N and 22°45'N and longitudes 78° 15'E and 78°30'E encompassing an area of approximately 806km2. Horton's method as modified by Strahler has been followed in ordering the drainage net work (Horton 1945; Strahler 1965) to establish an hierarchy of tributaries, which is a measure of stream position. It envisages the restriction of the designation of the order of stream segments. Small natural units as delineated by the various third order drainage basins were marked on the map for determining the various morphometric parameters and subjecting them to further statistical operations. The number of streams of various orders was counted, linear measurements were made with a rotameter, areal determinations by planimeter and epidiascopic enlargements for relief evaluation. Geology after Meddlicott (1873) was superposed on the drainage basin map. Basin distribution with corresponding lithology was as shown in Table 1, which excludes those drainage basins not confined within one stratigraphie unit. As is perceptible from the table, the number of basins present in a stratigraphie unit offers in itself a ready means of distinction. Apart from shades of variations in the mean values of stream numbers of first and second orders in different horizons, some simple statistical parameters unveil their peculiarities more convincingly. The Denwas-

Bagras and the Pachmarhis both are characterized by leptokurtic peaks but their frequency distribution is symmetrical and asymmetrical respectively. The recent formation however depicts platykurtic normally skewed distribution. Statistical inferences in case of Talchirs-Barakars which enclose only one third order basin have not been possible. The maximum channel length, various stream orders, is found in the recent formations though the deviation therefrom is also highest in contrast to others. The frequency distribution is normal, the two lower orders being platykurtic and the third order leptokurtic in nature. The Denwas-Bagras and Pachmarhis reveal positive right-handed skewness and a relatively sharp peak deviating away from the normal distribution (except for the second order streams of DenwasBagras which are within the limits of Normalcy). However, no definite trend is discernible so far as the length ratios are concerned. The maximum length of the basin is measured as the basin length. Its highest mean value lies in the recent formation followed by the Pachmarhis. Whereas both the skewness and kurtosis values indicate a normal distribution of basin length for the Denwas-Bagras, the recent formation differs from it in its platykurtic nature of almost identical magnitude and the Pachmarhis by its asymmetrical, positive right-handed skewness. The average area occupied by the third order drainage basin in the case of DenwasBagras and Talchirs-Barakars is identical and is about two-thirds that of Pachmarhis and a little over one-third that of the recent formation. The frequency distribution is normally symmetrical with a flatter peak for the latter. The basin areas of Pachmarhis are nonnormal in distribution though those of Denwas-Bagras are asymmetrically skewed, their peakedness lies well within normalcy. Basin shape in a way is an expression of

Géomorphologie / 7 l TABLE 1 Number of basins

Formations

Lithology

Récent Denwas-Bagras

Loose alluvium Mottled clays & shales with some sandstones Medium coarse grained sandstones, pebbly at places Boulders, conglomerates, clays, sandstones, etc. Crystallines

Pachmarhis Talchirs-Barakars Metamorphics

TOTAL

its catchment area per unit length and is therefore determined accordingly. It reveals that the frequency distribution in case of the Denwas-Bagras and Pachmarhis is nonnormal. It is so for the peak in the recent formation also but the symmetry is normal. It is rather unusual to observe an apathie behavioural pattern of bifurcation ratio frequency distributions. The first (F1/F2) ratio shows a non-normal but the second (F2/F3 ) indicate a normally skewed though flatter pattern in the recent formation. In the Pachmarhis the situation is just the other way. Further the Denwas-Bagras are distinguished from either by their right-handed positive symmetry and leptokurtic nature in the first ratio. So far as the second ratio is concerned though, the leptokurtic peak is within the limits for normalcy. The entire region falls into the category of very fine drainage texture though variation is considerable. Occasionally ultrafine textures are supported by the Pachmarhis, but the mean value for the Denwas-Bagras is higher. The distribution for stream frequency, drainage density, and texture in the case of former is asymmetrical and is therefore distinguishable from the normal ones of the DenwasBagras and the recent formations. Further the stream frequency and the drainage texture

18 52 86 1 0 157

for the former being platykurtic is differentiable from the latter's leptokurtic peak. Relief and slope characteristics are even more conspicuous. The average basin relief and the slope in the Pachmarhis is about 10-11 times greater than that in the recent formation and approximately 5 times the latter in the Denwas-Bagras. Similarly, the ruggedness number also varies. It works out to about Í 4 times for the Pachmarhis and about 6 times the recent formation in the Denwas-Bagras. The Upper Gondwana terrain of central India amply demonstrates the control of constituent rock units on the various morphometric attributes, climatic, and biogeographical effects being broadly constant. University Grants Commission, New Delhi; A.G. Jhingran. Horton, R.E., 1945 Erosional development of streams and their drainage basins: hydrophysical approach to quantitative morphology, Bull. Geol. Soc. Am. 56, 275370. Meddlicott, H.R., 1873 Notes on Satpura Coal Basin, Mem. GSI, I. Strahler, A.N., 1965 The Earth Sciences (New York).

P0139 Relief evolution of old mountains of central Europe during the Tertiary, with the Sudetes as an example WOJCIECH WALCZAK Wroclaw University, Poland Since 1924 when Walter Penck published his notable theory about piedmont planations, 3-4 horizons of Tertiary planation in European mountain ranges of medium height have been interpreted as the result of a cyclic

uplifting of particular rock massifs during the Alpine orogeny. For the most part their age is estimated on the basis of an analysis of their orography and their relation to correlated sediments in the mountain foreland, less

72 / Geomorphology often on analytic studies of regolith covers in situ. In recent years it has become possible in the Sudetes to establish age estimates of the Tertiary relief from palynological analyses and from plant remnants recovered from kaolin clays extending in younger horizons of Tertiary planations. Judging from what so far has been determined, the evolution has proceeded as follows: The Laramide movements of the YoungSaxon orogeny caused a regression of the Upper Cretaceous Sea and a subsidence of basins lined with deposits of this sea, combined with the emergence of their periphery. Afterwards, in the Paleocene, the Eocene, and the Lower Oligocène, tectonics were at rest for a long period of time and an intensive planation of the relief took place under conditions of a then-ruling humid subtropical climate. The result was that on magmatic and metamorphic rocks thick latérite and kaolin covers developed which have partly been preserved in situ on the plateau of the Sudetic foothills or, as correlated sediments, in the northern Sudetic foreland. Chemical weathering occurring at that time, was selective; thus, after denudation had removed covers of weathered material, dome-shaped hard-rock ridges or mogotes typical of tropical karst became exposed in a number of outcrops of carbonate rocks. This long-continued destructive action taking place during the time of tectonic rest for some 40 million years until the wane of the Middle Oligocène resulted, in the area of the Sudetes and the foreSudetic block, in a widespread planation surface covered by a thick mantle of weathered material and called the Paleogene planation surface. Fragments of this first surface have survived to this day in the shape of flat plateaus rising to different heights in the majority of Sudetic mountain massifs, being higher at the southern margin of the Sudetic block and lower in the zone adjoining its northern margin. Apart from the marginal fault, counterparts of the Palaeogene planation surface occur in the Sudetic foothills in the form of elevated plateaus at an altitude of 250 to 400m, containing remnants of the former regolith covers in the shape of deposits of kaolin, red earths, and latérites, 20 to 40m thick.

Next, as the result of the Savan phase of the Young-Saxon orogeny at the wane of the Middle Oligocène, the Sudetes were uplifted along faults and flexures which frequently were simply revivals of former Variscan dislocations. The most characteristic among them was the marginal fault of the Sudetes which separated the higher uplifted Sudetes from the Sudetic foothills, which up to then had constituted a joint body affected by the Palaeogene planation. The age of these movements is indicated by the correlated deposits mentioned, which are derived from erosive incisions in the emerging Sudetes made by rivers flowing in their foreland and underlie Upper Oligocène sediments and the Miocene brown-coal formation. Linked with this uplifting of the Sudetes during the Middle Oligocène was also the transformation of the river and valley system in the middle and eastern Sudetes; in the Paleogene these chains had been drained towards the Carpathian geosyncline. Now, the uplifting movement blocked the valleys which were turned southeastward, in favour of a concentric drainage pattern. An indication of the former flow direction of the Paleogene rivers diverted at that time can be seen in today's left-bank tributaries of the Nvsa Klodzka river. The amplitude of the uplift of the Sudetes during the Oligocène is estimated at 700800m at the southern margin of this massif, and 200-400m in the northern zone. These were the relative altitudes of the Sudetes of Oligocène times with regard to the surface of the foothills which were uplifted only some 100m. This Oligocène emergence of the Sudetes and the fore-Sudetic block was accompanied by volcanism. Streams of basaltic lava were emitted. This lava filled in some of the former valleys, and on interfluve fragments of the Paleogene planation surface in the foreland and the foothills of the Sudetes there developed a relief-protecting mantle of basalt covers and domes, underlying which here and there relics of the old regolith were preserved. Due to these facts the Upper Oligocène basalt eruptions represent an important stratigraphie stage in the evolution of the Tertiary relief of the Sudetes and their foothills. They represent the final stage in the development of the polygenic Paleogene planation surface which, due to its being

Géomorphologie / 73 locally covered by basalts, has been called the pre- or sub-basaltic surface. These basalts also separate the older planation surface from the younger planations which have developed partly within the basalt covers and are therefore called the post-basaltic planation surfaces. The Upper Oligocène and the Lower Miocene were in the Sudetes a period of pediplanation which took place while tectonics were at rest. The planation surfaces were developing due to recession of valley scarps and slopes and the formation of sub-slope pediments; this considerably widened the valleys which gradually changed from erosive ravines into wide flat-bottomed valleys. Today this evolution is illustrated by a widespread planation surface extending at 600800m or even 900m asl, also taking in the lower surfaces of lateral ridges and slope planations. Here and there on this horizon, relics of regolith covers survived in situ. The correlated deposits dating from this phase occur as fine-fractional quartz sands and as kaolin clays. They occur as intercalations amidst beds of brown coal of Upper Oligocène and Miocene age in foothill areas of Lower Silesia. This planation was brought to an end by a new phase of emergence of the Sudetes during the Middle Miocene. This led to a revival of erosion, causing the dissection of the wide Lower Miocene valleys and planation surfaces at 600-800m asl by new valley incisions reaching to 400-500m asl. In the Sudetic foothills the Lower Miocene surface of 400-3 Om altitude was at that time lowered to the level of 300-60m asl. Sandy-gravelly correlated deposits of this new phase of erosion were laid down in the foreland, in depressions filled in by the Miocene brown-coal formation. The Upper Miocene and Lower Pliocene was the third period of tectonic rest and of planation in the Tertiary evolution of the Sudetes. The climate was temperate, warm, and fairly dry, with occasional humid periods in which episodic heavy downpours washed down the products of weathering - clayey kaolin mantles, quartz sands, and loose rock debris - and deposited them in the foreland of the Sudetes in the vast basin of the freshwater Poznan lake. These deposits are evidence that there was created, from material

denuded from the Lower Miocene planation surfaces, the lowermost planation horizons at 400-500m altitude in the Sudetes and at 30060m asl in the foothills. These surfaces developed in consequence of denuding processes on valley scarps and slopes which caused the formation of pediment-type planation surfaces at the slope bases. This denudation affected all soft and less resistant rock types also, while from more resistant formations numerous hard-rock hills were left standing, characteristically dome-shaped and 3 0-5Om high. The relative age of the last-named planation surfaces has been determined by palynological methods. These surfaces, extending at a mean altitude of about 500m, locally contain shallow bowl-shaped undrained depressions filled with kaolin clays and a strongly weathered debris of surrounding rock formations. These are relics of Tertiary boisons, as indicated by the fact that in the clayey beds 40-48 per cent sporomorphs of light-loving plants of species Hystrix were found as well as a variety of other Mio- and Pliocene plants like Podocarpus docrydioides, Taxodiaceae, Cupresaceae, Nyssa, Caria, and Pterocaria. Similar species were also identified in the kaolin clays deposited in the watershed pass of the Swierka river in the Middle Sudetes, at 570m asl. Further proof of the Mio-Pliocene age of the valley relief is a bed of clayey kaolin debris more than 12m thick, which lies at 360m asl, some 40m above today's floor of the Scinawka Kiodzka valley in the Middle Sudetes. In this bed, leaf remnants of a MioPliocene flora were identified. This evidence shows that as early as the wane of the Miocene and in the Lower Pliocene the relief of the Sudetes, although somewhat lower than it is today, had been completely formed in its mean features resembling those seen today. Correlated deposits of the last Tertiary phase of deep erosion in the Sudetes (following their Middle and Upper Pliocene uplifting), have accumulated in the foothills to a thickness of fifty or so metres, forming cover sheets overlying the Miocene and Lower Pliocene clays or filling the valley channels eroded in these clays. These deposits are impoverished quartz sands and gravels containing inclusions of kaolin clays; these same

74 / Geomorphology deposits also appear in lower sections of the main Sudetic valleys where they constitute former valley floors below the contemporaneously accumulated layer of ground terraces or pediments of high Pleistocene terraces. Penck, W., 1924 Die morphologische Analyse (Stuttgart). Jahn, A., 1953 Morphological problematics of the Western Sudetes, Prhegl. Geog. xxv, 3 (Warszawa) (in Polish). Klimaszewski, M., 1958 Geomorphological

evolution of Poland's territory in the preQuaternary, Przegl. Geog. xxx, 1 (Warszawa ) ( in Polish ). Piasecki, H., 1964 Morphological evolution of the Kaczawa foothills. (PhD thesis, Wroclaw) (in Polish). Pulina, M., 19 64 Karst phenomena in the Sudetes. PhD thesis, Warszawa (inPolish). Walczak, W., 1968 Lower Silesia, I: The Sudetes. PWN Warszawa. - 1970 Lower Silesia, II: The fore-Sudetic Area. PWN Warszawa (in Polish).

P014Q

Fluvial cave sediments: a description and interpretation THOMAS E. WOLFE State University of New York at Buffalo, The literature concerning clastic cave fills generally doesn't discuss the origin and structures of cave sediments deposited by flowing water. Most passages in limestone caves are generally acknowledged to have formed below the water table by slow percolation of groundwater (White and White 1968). Although considerable literature on cave deposits exists, fluvial sediments are poorly described by most authors. In nearly all caves there is a sequence of clastic sediments that have sedimentary structures, a systematic vertical variation in grain size, and bedding characteristic of fluvial transport. Wolf e (1971) studied Appalachian caves and estimated that their sediments are generally composed of over 75 per cent allochthonous deposits (all of which is stream transported). About 20 per cent can be classed as autochthonous deposits (mainly ceiling breakdown and rockfall). The remaining less than 5 per cent is classed as chemical deposits (mainly carbonates). The material most common to cave fills, fluvial sediment, has received the least attention. Most of the literature concerning fluvial sediments in caves is vague or characterized by misinterpretations. The purpose of this paper is to pose and answer some basic questions concerning the origin and nature of fluvial cave sediments. 1. At what stage in the development of caves do most fills enter the caves? 2. How are cave deposits similar or dissimilar to surface fluvial deposits? 3. What are the relations between ancient

USA

and modern allochthonous fluvial deposits in caves? 4. Is there a cycle of sediment deposition in caves? Three cave systems in the Greenbrier Limestone karst of West Virginia are discussed in this paper: 1. Poor Farm Cave; 2. Bob Gee Cave; 3. Culver son Creek Cave. Cave fills are probably deposited in the later stages of passage evolution by higher velocity streams than those which are responsible for the original development and solutional enlargement of the passage. If fills are deposited in the first stage little evidence for them remains. Although pipe-full flow conditions exist during the phreatic generation of the cave passage, the conditions of deposition at this stage are not understood. The deposition of coarse bedload probably occurs under vadose (free-surface) flow quite similar to surface conditions. It might be argued that even under para-phreatic pipe-full flow conditions, bedload acts the same in a cave as in a surface stream, because bedload movement is independent of such channel conditions. Suspended load, on the other hand, may be affected by the depth of water in.a cave stream. Channel deposits predominate; because of a lack of space in a passage there is little area in which to develop a flood plain. During times of low water, point bars and braided streams develop in sediments laid down under flood conditions and are occasionally preserved.

Géomorphologie / 75 The confinement of the system means that lateral shifting during floods is impossible; on some occasions this can result in reworking of sediments in upper levels of a cave. As also observed by Palmer (1969), roof falls and ceiling breakdown may cause ponding and the deposition of fine-grained sediment, thus completely blocking the active route of a cave stream. This may help to explain why many passages show distinct intervals of vertical stacking rather than the continuous downcutting characteristic of surface valleys. The stacking of cave passages has long been observed and both filling by sediment and the shielding of bedrock by sediment may help to explain the discontinuous nature of such vertically-stacked, horizontal cave passages. Although provenance changes, surface and subsurface diversions, and other complex variables make it difficult to assess the exact differences between present and past conditions, some quantitative statements can be made: 1. There appears to be a smaller quantity of material being deposited in channels of contemporary cave streams. If a simple comparison of maximum depth of sediment to passage height ratios between abandoned and active passages is made, it is found that active passages have lower ratios (example: Poor Farm upper passage 30/30=1.00, lower passage 5/15=0.33). 2. The mean grain size of bedloads in contemporary and ancient channels appears to be comparable. 3. Most contemporary cave streams are accumulating more sediment than they discharge (based on field measurement from sediment budget studies of five active cave streams). 4. In both conditions, there is strong evidence for slow phreatic flow generation followed by rapid vadose deposition of fills. As indicated by measurement of solutional scallops, velocities of contemporary cave streams are greater than those of ancient cave streams. The stacking of passages with discontinuous downcutting of channels, each with a distinct vadose invasion of bedload that generally coarsens upward and eventual cessation of flow and deposition of suspended load, appears to repeat itself vertically many times in caves of the Appalachians. As many as four such cycles have been observed in one cave.

Sediments may shield the bedrock floor of a cave from further solution, thus preserving the phreatic features (profile and solutional scallops) of a passage. This shielding effect may help to explain the nature of the discontinuously downcut passages. Complete filling of passages causes stream diversion and relocation, and such filling may be cyclically repeated as long as regional base level continues to be lowered and as long as there is limestone available for solution. Brain, C.K., 1958 The Transvaal ape-man bearing cave deposits. Mem. 11, Transvaal Museum, Pretoria. 32-48. Bretz, J.H., 1942 Vadose and phreatic features of limestone caverns, /. Geol. 50, 675-811. - 1953 Genetic relations of caves to peneplains and big springs in the Ozarks, Am. J. Sci. 251, 1-24. - 1962 Dynamic equilibrium and the Ozark landforms, Am. J. Sci. 260, 427-38. - 1965 Geomorphic history of the Ozarks of Missouri, Mo. Geol. Survey and Water Resources 41, 2nd Ser. Collier, R.C., and R. F. Flint, 1964 Fluvial sedimentation in Mammoth Cave, Kentucky, uses Prof. Paper 475-D. Davies, W.E., 1958 Caverns of West Virginia, W.Va. Geol. and Econ. Survey, 59A. Davies, W.E., and E.C.T. Chao, 1959 Report on sediments of Mammoth Cave, Kentucky, uses Admin. Rpt. Goodchild, M.F., 1969 A study of erosional scallops, PhD thesis, McMaster University. Harmes, J.C., and R.K. Fahnestock, 1965 Stratification, Bed forms and flow phenomena, in G.V. Middleton, éd., Primary Sedimentary Structures and their Hydrodynamic Interpretation. Soc. Econ., Paleon., and Min. Special Pub. 12, 84-115. Helwig, J.A., 1964 Stratigraphy of detrital fills of Carroll Cave, Camden Co., Missouri, Missouri Speleology 6, 1-15. - 1965 Geology of Carroll Cave, Camden Co., Missouri, Nat. Speleo. Soc. Bull. 27, 11-26. McKee, E.D., E.J. Crosby, and H.L. Berryhill, Jr., 1967 Flood deposits, Bijou Creek, Colorado, June 1965, /. Sed. Petrol 37, 829-51. Palmer, A.N., 1969 A hydrologie study of the Indiana karst. PhD thesis, Indiana U. Reams, M.W., 1968 Cave sediments and the

76 / Geomorphology geomorphic history of the Ozarks. PhD thesis, Washington U., St. Louis. Schmid, B., 1958 Hohlenforschung und sedimentanalyse. Schriften Institut, Urund Fruhgeschicte Schweiz, 13. Vanoni, V.A., 1966 Sediment transportation mechanics: initiation of motion, /. Hydraulics Div. Proc. Am. Soc. Civ. Eng. 92, 291-314. White, W.B., 1964 Sedimentation in caves: a review, Nat. Speleo. Soc. News, 21,

152-3. White, W.B., and E.L. White, 1968 Sediment transport in limestone caves, Nat. Speleo. Soc. Bull. 30, 115-29. Wolfe, T.E., 1970 Sediments of Bob Gee Cave, West Virginia, Canadian Caver, 314. - Surface and subsurface sedimentation in basins along the Allegheny front of southeastern West Virginia. PhD thesis, McMaster U.

P0141 A classification of Martian terrain features on two Mariner 6 photos JAMES F. WOODRUFF University of Georgia, USA

This paper tests the applicability of a terrain classification system such as Fenneman's to the Martian landscape shown on two nearencounter frames from the Mariner 6 mission. 6N21 is a wide angle (A) camera photograph taken during the closest approach to the planet. It includes an area of about 902 X 701km centred on 15°E and 16°s. The northern 1 / 4 of the photo lies within Meridiani-Sabaeus Sinus, early recognized by telescopic observations as part of an eastwest equatorial dark band. To the south is the light area of Deaucalionis Regio. These two areas comprise the primary categories of terrain in the study area and are comparable to provinces in Fenneman's systems. The boundary separating Deaucalionis Regio and Meridiani-Sabaeus Sinus, based upon difference in reflectance, is fairly easy to trace over 2/3 of the width of 6N21. These provinces, however, if they warrant recognition as distinct geomorphic regions should have some terrain quality other than albedo which might be identified on the photos. Spectra analysis indicates that Meridiani Sinus is a highland area with peak altitudes of 2.43km above a datum plane established by the 6 millibar atmospheric level (Herr et al. 1970). The regional slope to the south averages about .4° and has a drop of more than 2.2km to Deaucalionis Regio. Although such differences of elevation are not apparent on 6N21 there are some supporting features. A number of faint sinuous east-west lines might be interpreted as south-facing escarp-

ments, particularly a rather conspicuous promontory which appears to project southward for 8-10km. Using a maximum slope angle of 20° estimated by Murray, this promontory appears about 800m high (Murray et al. 1971 ). The line of which this escarpment is a part continues to the west for more than 300km. Over much of its length it looks canyon-like and in several instances is joined by tributary escarpments or canyons in a dendritic pattern. These features, whether a series of escarpments resulting from lava flows, faulted forms, or canyons formed by collapsed lava tubes, suggest a regional slope accomplished by several plateau levels and bordering escarpments. Cutts concludes that there is no connection between crater abundance and light or dark areas and certainly data from a single photo is statistically invalid; however, the number of craters larger than 5km in diameter seems to decrease consistently to the south while the number of vestigial craters increases (Cutts et al. 1971). These relationships would not be inconsistent with a process which was deflating the highlands and depositing materials in the lowlands to the south. Crater walls in the northern highlands appear to be smoother and there are few inner concentric rings. Such morphological differences may imply lithological differences and, with the other factors, would seem to validate the recognition of two major physiographic regions on this photo. 6N21 fails ot reveal any secondary physiographic units comparable to Fenneman's 'sections' or perhaps no section boundaries

Géomorphologie I 77 occur hi the area covered by 6N21. However, variations in elevations in both provinces, indicated by spectra studies, could form the basis for 'section' regionalization. Small area geomorphic assemblages analogous to Fenneman's 'districts' and individual terrain features are visible on narrow-angle B camera photos. One of these, 6N22, shows a small portion of the area also covered by 6N21. Cutting diagonally across this frame is a section of the wall of a crater more than 250km in diameter. A flat-floored crater about 25km in diameter is superposed upon this rim in the centre of the southern margin of the photo. A number of bowl-shaped craters and other relief features are scattered across the area. The most obvious terrain unit is the floor of the larger craters. Although this almost level surface, under magnification, shows a vague suggestion of soft linear hills the largest relief features on the craters floors are the cup-shaped craters. A typical one is about 3km in diameter, has a rim which rises 50m above the surrounding surface, and a central depression of about 250m. The third recognizable relief feature on the crater floor is a low rectangular section along the foot of the large crater wall. In some aspects it resembles a single lava flow but faulting might be an origin more consistent with other linear features of the crater wall. The crater wall itself is the roughest and most complex morphological sub-unit. It rises 450 to 900m to an irregular crest above the crater floor. Notable along the walls are NE-SW lineaments which may be scarps formed by secondary faults transverse to the crater rim. From its crest the rim drops 250 to 500m to the general surface of the intercrater areas. On 6N22 this intercrater surface is a rolling upland of irregular chains of hills and a grouping of higher rough terrain along the extreme northern margin. Ridges on the upland are 1 to 2km wide, convex, broad crested and anastomatic in pattern. Assuming average slopes of less than 10°, these ridges have a maximum relief of about 80m. Within this district there is the impression of softness of features which may result from erosion of the éjecta blankets of both the large crater and

the superposed blanket of the 25km crater. Although erosion by base surge at the impact forming the smaller crater may have modified the older éjecta blanket, some other process must be responsible for the softness of the surface relief forms and the absence of secondary crater swarms so typical of similar lunar situations. The last distinct topographic district on 6N22 is the rough area along the northern margin. Reference to 6N21, the wide-angle photo, indicates that this is part of another extensive escarpment system several hundred metres above the general intercrater area to the south. These hills may be the edge of a very large plateau system forming Sabaeus Sinus, and may mark a major division boundary rather than a discrete minor topographic unit. In summary the surface features of Mars shown on these two photos are amenable to a classification system similar to Fenneman's. Even though eroded, the fact that morphological detail results primarily from impact will result in minor topographic units being radial and related to each of the larger craters. Very similar terrain units may, thus, be discontiguous but replicated many times. Not only will organization at the small-unit level be point oriented, as have stratigraphie studies on the moon, but complete photo coverage may necessitate organizing physiographic regions, even at the province level, to major impact events. For instance, how far do the features associated with Nix Olympica dominate the Martian landscape? Such questions wait for solution upon the imagery from Mariner 9. NASA NCR 11-003-033 contract. Gutts, James A., L.A. Soderblom, R.P. Sharp, B.A. Smith, and B.E. Murray, 1971 The surface of Mars, 3, light and dark markings-, /. Geophys. Res. 76, 343-56. Herr, K.C., D. Horn, J.M. McAfee, and G.C. Pimentel, 1970 Martian topography from the Mariner 6 and 7 spectra, Astron. J. 75, 883-94. Murray, B.C., L.A. Soderblom, R.P. Sharp, and J.A. Cutts, 1971 The surface of Mars, 1, cratered terrain, J. Geophys. Res. 75, 313-30.

78 / Geomorphology P0142 The reworking of shores in the permafrost zone

FELIX ARE Academy of Sciences, USSR

The processes of reworking of shores formed by permafrost have some peculiarities, connected with the rocks' frozen state and ice content. Particularly, the shores formed by quaternary deposits with large content of constitutional ice and ice-wedges, widespread in the lowlands of eastern Siberia and eastwest of the USSR, have original development. Such deposits are called the ice complex. The main motive force of development of shores formed of the unfrozen rocks is the mechanical energy of waves. In reworking of shores formed by permafrost the thermal factor, which can cause the retreat of shores even in the absence of significant wave action, plays the important role. In the complex of intercommunicated processes of reworking of shores formed by permafrost, three main processes are distinguished: (a) thermoabrasion - the destruction of shore zone under the action of mechanical and thermal energy of water; (b) thermodenudation - the destruction of cliffs under the action of the air thermal energy and solar radiation; (c) thermokarst - the thawing of the basin's bottom under the action of thermal energy of water leading to subsidence of the bottom surface, conformably to reworking of shores. The consequence of thermoabrasion is the shore's retreat: Thermodenudation leads to the flattening of cliffs only: thermokarst, deepening the basin, promotes abrasion and, in certain conditions, interacting with thermodenudation, leads to shore retreat without thermoabrasion. In the permafrost absence the retreat of shores under the action of abrasion continues until the formation of the end profile of submerged shore slope (Zenkovich 1962, 153). The character of reworking of shores formed by permafrost depends radically on the ice content in rocks. Shores formed by ice only retreat very rapidly. Thus, in the permafrost zone there is a critical value of total ice content for each shore, above which the shore becomes unstable. After full thawing of rocks having critical ice content their subsidence level may coincide with water level in the basin. If the subsidence level is above the water

level, the shore is stabilized and can retreat under the action of abrasion only. In this case the end profile of the submerged shore slope forms when the thermokarst process is over. If the subsidence level is lower than water level the permafrost thawing under the basin bottom inevitably leads to its exposure and thawing in the cliff. In such a case the shore is unstable and will retreat even without the action of abrasion. The comparative roles of thermoabrasion, thermodenudation and thermokarst in the processes of shore development depends on the climatic conditions. So the regional peculiarities of shore reworking are defined by the climate. The climate of Arctic coastal lowlands is characterized by high winds and cold summers. The temperature of water in the basins is low. That is why the thermoabrasion plays the main role in shore reworking in these regions. The rate of shore retreat has reached great values. The thawing of basin bottom is developing slowly. Closed taliks are widespread even under the bottom of large lakes. The climate of the eastern Siberia taiga zone is characterized by weak winds, hot summers and large insolation. In these regions the role of thermoabrasion in shore reworking is unimportant. Basically, the shores retreat under the action of thermokarst and thermodenudation, at a slower rate than in the northern coast, other things being equal. According to the form of the profile, the cliffs may be divided into four main types: sloping, sloping with plumb lower part, plumb and stepped. The cliff form is the approximate indicator of shore retreat rate. In this respect the stepped cliffs give the biggest possibilities, but they are met seldom (Are 1968,97). If the cliffs are plumb and have wave-cut notches, or there are remains of fallen blocks at their foot and no large accumulations of products of thermodenudational destruction, then it means that the shore is retreating rapidly, and the rate of retreat exceeds the intensity of thawing of permafrost exposed in the cliffs. If the cliffs are sloping or sloping with

Géomorphologie I 79 plumb lower part, with ice-wedges exposed on their surface and there are no large accumulations of thermodenudational destruction products at their foot, the rate of shore retreat is approximately equal to the rate of thermodenudation. If the cliffs of any form have large trains of thermodenudational destruction products at their foot, the rate of shore retreat is less than the rate of thermodenudation. In the process of thermodenudation of cliffs formed in the ice complex, thawing of exposed enclosing rock material may be slower than the thawing of the ice wedges, but on the whole the rate of thermodenudation of such cliffs is defined by the intensity of ice-wedges thawing. This value may be easily calculated by climatic characteristics and be used for approximate definition of the rate of retreat of shores formed in the ice complex on the basis of indications stated above. On the Yakut coast of the Laptevs' Sea the rate of thermodenudation is 4-5m a year (Are, Molochushkin 1965, 138), and in central Yakutia is 9-10m a year. At the present time there are no data for the satisfactory prediction of reworking of intensively retreating shores in the destruction of which thermoabrasion plays the main role. The observations made by N.R Grigorjev (1962, 71 ) and the author show that during recent years the mean rate of retreat along the shores of the Laptevs Sea does not exceed 4-6m a year. In Arctic coastal lowlands the rate of lakeshore retreat has reached 10m a year under the active influence of thermoabrasion (lakes more than 1km across), and under the weak influence of thermoabrasion l-3m a year

(Tomirdiaro, Rjabchun, Golodovkina 1969, 249). In the taiga zone of eastern Siberia the destruction of lakeshores is going on under the small influence of thermoabrasion. In central Yakutia plumb cliffs with wave-cut notches, or even without trains of destruction products at their foot, are met seldom, even at the largest lakes with dimensions of some kilometers across. That is why the rate of shore retreat is less than the rate of thermodenudation and on the average does not exceed 0.5m a year. The prediction of reworking of such shores is possible with the help of calculations concerning rock thawing and subsidence by methods which are known. Are, F.E., and E.N. Molochushkin, 1965 The rate of destruction of arctic cliffs in Yakutia under the influence of thermodenudation. Processy teplo- i massoobmena v merzlykh gornykh porodakh, 130-8. - 1968 The development of thermoabrasive shores relief, Izv. Akad. Nauk SSSR, ser. geograf., no. 1, 92-100. Grigorjev, N.F., 1962 The role of cryogenic factors in the formation of sea shores in Yakutia. Mnogoletnemerzlye parody i soputsîvujushchie im javlenija na territorii Yakutskoi ASSR, 68-78. Zenkovich, V.P., 1962 The principles of teaching on sea shore development, 153. Tomirdiaro, S.V., V.K. Rjabchun, and A.D. Golodovkina, 1969 The reworking of ice saturated shores of basins and reservoirs in arctic and subarctic lowlands of northeast of the USSR, Trudy soveshchanija po izucheniju beregov vodokhranilishch i voprosov drenaga v uslovijakh Sibiri 1, 244-60.

P0143 Commentaires relatifs à une coupe transversale de la moraine terminale de SaintNarcisse à Saint-Gabriel-de-Brandon, Québec

ROBERT DENIS

Université du Québec à Montréal, Canada

La région de Saint-Gabriel-de-Brandon se situe à quelques 110 kilomètres au nord-est de Montréal. Elle se localise près de la limite topographique de la plaine et du plateau laurentiens et près du contact géologique des basses-terres du Saint-Laurent et du Bouclier canadien (voir la carte de localisation). La moraine terminale de Saint-Narcisse, dont

nous analyserons une coupe transversale qui se situe au sud-ouest de la ville de SaintGabriel, suit grossièrement ces frontières physiques. Cette moraine fut étudiée par maints auteurs géologues ou géographes dont Mawdsley 1927; Osborne 1950; Laverdière et Courtemanche 1959; Gadd et Karrow

80 / Geomorphology glaciaire, soit immédiatement avant ou après ce premier dépôt morainique. La déglaciation dans les basses-terres du Saint-Laurent aurait débuté il y a plus de 12,000 ans BP (Lasalle 1966, 120). Cette retraite coïnciderait avec une période de x réchauffement climatique (Two Creeks, 'sensu lato' selon Lasalle 1966, 99). La Mer de Champlain, après le dégagement par les glaces du goulot l'étranglement situé en amont de la ville de Québec, s'infiltra dans la plaine du Saint-Laurent et ses eaux se haussèrent au niveau de la région à l'étude tout en talonnant l'inlandsis en retrait. Dans les bas-fonds se décantèrent les particules fineá en suspension, tels les limons et les argiles Fig. 1. (V). Un refroidissement climatique, signalé par un changement phytogéographique (Lasalle 1959; Elson 1962; Parry 1963; Lasalle 19661966, 95), entraîna une récurrence glaciaire 70; Denis 1970-71; Lasalle et Hardy 1971. qui semble s'être traduite ici par une oscillaLa réalisation de la coupe transversale de tion de faible amplitude de l'ordre de 3 kilola moraine terminale de Saint-Narcisse à mètres, ce qui, à l'échelle continentale, peut Saint-Gabriel a été rendue possible, grâce être considéré comme une simple halte. aux campagnes de forage menées dans cette Durant ce stade, fut mise en place une région, par une équipe du Service d'Hydromoraine terminale (3) gravelo-caillouteuse à géologie du Ministère des Richesses Natumatrice sableuse qui recouvra des sédirelles du Québec sous la conduite de Claude ments fluvio-glaciaires ou marins (V). Grenier, ingénieur (Grenier 1965, 1970; D'ailleurs, en surface, certaines coupes nous Grenier et Denis 1971). C'est particulièrefont voir ce type de recouvrement (Denis ment à partir de l'examen approfondi des 1970, en préparation). L'âge de cette moraine nombreux échantillons prélevés à cette occaserait d'environ 10,500 ans BP, d'âge Valders? sion, de celui des diagrammes électriques en(Lasalle 1966, 95). registrés dans chaque puits et des observaAu bourrelet morainique terminal se tions consignées dans le journal du foreur rattache un immense delta (3a) (6.2km2) que nous avons pu dresser cette coupe. En qui se forma dans les eaux de la Mer de outre, l'expérience du terrain acquise au cours Champlain, grâce à l'apport de sédiments des années 1970-71, alors que nous avons véhiculés par tous les cours d'eau s'échappant, cartographie les formations superficielles de de façon radiale, du front glaciaire. Ce delta cette même région pour le compte du Service se hisse à une altitude de 215 mètres environ. de l'Exploration géologique du même minisL'ensemble des lits sommitaux et des lits tère, nous a permis de saisir les liens qui frontaux atteint une puissance totale voisine existent, à une échelle régionale, entre ces de 50 mètres (voir forages nos 22024-6). diverses formations et d'esquisser une hypoCes lits recouvrent des lits basaux de limon thèse relative aux conditions paléo-géoet d'argile qui dans le cas du forage no 22024 graphiques de leur mise en place respective. comportent 5 séquences de puissance variable Sur le socle précambrien (1 ), nous notons séparées par autant de séquences de till de la présence d'une moraine de fond (2) défond. L'épaisseur totale de ce dernier enposée par la calotte glaciaire du Wisconsin. semble atteint près de 72 mètres. Nous inElle comporte des blocs, des cailloux et des terprétons ces séquences de till comme autant graviers (classification granulométrique de des dépôts morainiques largués dans la mer Wentworth 1922, modifiée) enrobés dans une par des plates-formes flottantes ou par des matrice sablo-limoneuse. Nous n'avons pu icebergs que vêla l'inlandsis lors des oscilladistinguer de ce till de fond les dépôts fluviotions de la ligne de contact glacier-mer. glaciaires, mis en place par les eaux de fonte Un autre réchauffement climatique survint

Géomorphologie I 81 COUPE

TRANSVERSALE

DE LA MORAINE TERMINALE DE SAINT-GABRIEL-DE-BRANDON

SAINT-NARCISSE

QUÉBEC

Fig. 2. qui eut pour effet de faire reculer la calotte glaciaire sur les plateaux qui bordent le bassin du lac Maskinongé au nord et à l'ouest, tout en laissant mourir dans celui-ci un énorme culot de glace. Ce dernier se fragmenta en plusieurs blocs de tailles différentes particulièrement dans le secteur qui s'étend au nord de la moraine terminale (forage no 22025 ). La fusion sur place de tous ces blocs créa une topographie typique de bosses et de

creux (voir profil topographique au sud du forage no 22025). La fonte de l'inlandsis entraîna un relèvement du niveau des eaux marines qui envahirent très tôt le bassin formé par la moraine terminale de Saint-Narcisse. Des limons et des argiles (5) se décantèrent en eau profonde, surmontés de sables et graviers deltaïques et de rivage (6). (Nous n'avons point distingué les sédiments de rivage du début de

82 / Geomorphology la transgression marine.) Le relèvement isostatique concurrença le relèvement eustatique, ce qui entraîna le départ des eaux marines et la formation de différentes terrasses qui s'emboîtèrent les unes dans les autres. Il semble que l'épisode de la Mer de Ghamplain aurait duré 'de l'an 11,400 jusqu'à un peu avant 9,500 ans avant aujourd'hui' (Lasalle 1966, 91). Enfin, le même travail d'accumulation et d'érosion fluviales s'est poursuivi depuis cette période jusqu'à ce jour. R. Roy, C. Grenier et leur équipe du Service d'Hydrogéologie du Ministère des Richesses Naturelles du Québec. Denis, R., 1970 Géologie des dépôts meubles de la région de Saint-Gabriel-deBrandon (feuillet ouest), Min. des Rich. Nat. Qué. R.P. Elson, J.A., 1962 New England Int. Geol. Conf., Guide-Book, McGill U. (Montréal). Gadd, N.R., et P.F. Karrow, 1959 Surficial Geology of Trois-Rivières Area, Canada Geol. Sur. (Québec), Map 54-1959. Grenier, C., 1965 Levé hydrogéologique à St-Gabriel-de-Brandón, Min. Rich. Nat., R. no. 550-A.

- 1970 Levé hydrogéologique à St-Gabrielde-Brandon, ibid., no. 550-B. - et R. Denis, 1971 (en préparation). International Water Supply Ltd., 1967 Essai de pompage à Saint-Gabriel-deBrandon (non publié). Lasalle, P., 1966 Late quaternary vegetation and glacial history in the St Lawrence Lowlands, Canada, Leidse Geol. Med. 38, 91-128. - 1970 Notes on the St Narcisse morainic system north of Quebec City, Can. J. Earth Sci. 7 (2), 516-21. - et L. Hardy, 1971 (en préparation). Laverdière, C., et A. Courtemanche, 1959 La géomorphologie glaciaire de la région du Mont-Tremblant, Rev. Can. Géog. 13, 103-34. Mawdsley, J.B., 1927 St Urbain area, Charlevoix district, Canada Geol. Surv. (Québec), Mem. 152. Osborne, F.F., 1950 Marine crevasse fillings in thé Lotbinière région, Québec, Am. J. Sci. 248, 874-9Q. Parry, J.T., 1963 The Laurentians: a study in geomorphological development, PhD thesis, McGill U. Wentworth, C.K., 1922 A scale of grade and class terms for clastic sediments, /. Geol. 30, 377-92.

P0144 The role of wind in periglacial environments, with special reference to northwest Banks Island, western Canadian Arctic HUGH M. FRENCH University of Ottawa, Canada The role and importance which should be given to wind in periglacial environments is not yet fully understood. However, because of the absence of vegetation, the frequency of strong winds, the lack of precipitation, and the presence of large amounts of comminuted debris produced by cryonival processes, it has generally been thought that the wind plays a dominant role in both erosion and transportation (e.g. Embleton and King 1968, 564-85). In those temperate latitudes which experienced severe periglacial conditions during the Pleistocene, such as central and eastern Europe, the presence of fossilized stone pavements, wind-modified pebbles and blocks, loess deposits, and frosted sand grains also suggest a similar interpretation. On the other hand, more recent studies in the Cana-

dian arctic have stressed that the wind is only important in the sense that it gives rise to a number of minor features or secondary effects associated with deflation, deposition and aeolian activity (e.g. Pissart 1966; Bird 1967, 237-41). Based upon observations made in northern Banks Island in the western Canadian Arctic, this paper attempts to illustrate the complex role which the wind plays in some periglacial environments today, and in particular, to stress the importance of this factor in the production of asymmetrical terrain. Field studies undertaken in 1968 and 1969 involved the mapping of slopes, soils, and surficial materials on the Beaufort Plain of Banks Island. The Beaufort Plain is a distinct physiographic region in the extreme north-

Géomorphologie I 83 west corner of the island and is distinguished from adjacent areas by the remarkable uniformity of the surface, the absence of thaw lakes and depressions, and the lack of a recognisable veneer of glacial sediments. The nearest permanent climatological station is at Mould Bay, on adjacent Prince Patrick Island and this may be considered to be the most representative station for northern Banks Island. At Mould Bay, throughout the year, the dominant winds are from the northwest and north, with average speeds in excess of lOmph. On the Beaufort Plain, the wind is even more important since this part of Banks Island is directly exposed to strong westerly winds moving across the Beaufort Sea, and is not sheltered to the same extent as Mould Bay. Because of the absence of topographic barriers on the plain, the modifying effects of these winds are carried well inland and the Beaufort Plain experiences a damp maritime periglacial climate. Precipitation and snowfall is small in all parts of the western arctic and in many respects the region may be regarded as an arid one. Nowhere is there more than 4 inches of precipitation per annum and over 50 per cent occurs during the summer months falling as either rain or snow. The major landform characteristic of the plain is the widespread fluvial dissection which is currently taking place. A series of major streams run approximately parallel to each other in an east-west or southeast-northwest direction to form the main drainage lines, and dissect the plain to depths of over 150' in places. Observations in the field and from air photographs has revealed the following further characteristics: (a) An asymmetry of the valley side slopes. In those valleys or sections of valleys which aré aligned east-west or southeastnorthwest, the southwest or south facing slope is usually the steeper while the northeast or east facing slope is gentler. Flat bottomed, braided stream channels occur at the immediate foot of the steeper slopes which are, themselves, closely dissected by minor gullies, the probable result of erosion associated with ice wedges. The form of the gentler slope is also distinctive since it is characterised by a series of small bluffs of between 10° and 15° in angle, separated from each other by long gentle rectilinear slopes of very low angle. (b) An asymmetrical pattern of soils and

materials. While Polar Desert or Storkerson soils cover much of the surface of the plain, tundra soils (meadow tundra, Kelleft series; upland tundra, Bernard series) are the next most widespread. They occur, however, predominantly upon slopes of north and east exposure, i.e. upon the gentle slopes of the asymmetrical valleys. Regospls, by contrast, occur as narrow strips coinciding with the steeper south and southwest exposures. This gives a strong cross-valley contrast between vegetated tundra slopes and unvegetated gravel slopes. (c) An asymmetry of surface tributary patterns. Although the major drainage lines are either east-west or southeast-northwest, there is a preponderance of small tributaries and gullies which enter the master streams from the south and west, thus dissecting the gentle and vegetated tundra slopes. The tributaries take on a variety of forms ranging from shallow, bowl-like depressions below snowpatches to distinct stream channels. The asymmetrical nature of the slopes, soils, and stream patterns appear to be closely interrelated. The asymmetry is the result of differing microclimates existing upon differently exposed slopes, which favours the development of solifluction processes upon the north and northeast facing slopes. As a result, the stream in the valley bottom moves laterally towards the slope producing the least colluvium. That slope is then undercut and steepened in angle by fluvio-thermal erosion processes (French 1971, 726-27). The basic microclimatic differences on the two slopes can be related to the dominant westerly winds in this part of the arctic. During the winter, these winds deposit snow preferentially upon lee slopes and in small gullies, while the exposed south and southwest facing slopes, and the surface of the plain, are continually swept clear. During the summer, the snow distribution and the continuation of the westerly wind produces a complex geomorphic situation. The exposed snow-free slopes are kept cool by evaporation and consequent latent heat loss, resulting in a thin active layer and relatively little solifluction. By contrast, the lee slopes do not suffer from evaporation cooling and are therefore warmer, with deeper active layers. In addition, the snowbanks provide moisture for the rapid operation of solifluction processes and also, by the gradual concentration

84 / Geomorphology of sheetwash below the snowpatch, they favour the development of surface drainage. As a result, the north and northeast facing slopes become characterised by nivation hollows and snowbanks, solifluction and tundra soils, and sheetwash and surface runoff during the summer months. The major conclusion to be drawn from this study is that the prevailing wind, through its influence on the deposition of snow in winter and the soil temperatures on exposed slopes during the summer, exerts a control over nivation, solifluction, and ultimately, fluvial processes. This control is reflected in the asymmetry of the landscape. In that the asymmetry is restricted to the extreme northwest corner of the island, the wind factor may be regarded as being essentially a local factor varying in importance from locality to locality, and depending upon the strength and variability of the wind. In more general terms, it is clear that in this part of the arctic, the wind is more important in determining local microclimatic conditions than is insola-

tion. The theoretical temperature variations with orientation, as described in many texts (e.g. Geiger 1959, 215-29), do not occur. Finally, it would appear best to regard the wind as being capable of exerting a fundamental control over several different periglacial processes rather than being, itself, either a primary or secondary process. Bird, J.B., 1967 The Physiography of Arctic Canada (Baltimore). Embleton, C., and C.A.M. King, 1968 Glacial and periglacial geomorphology (Arnold). French, H.M., 1971 Slope asymmetry of the Beaufort Plain, Northwest Banks Island, Can. L Earth Sci. 8, 717-31. Geiger, R., 1959 The Climate near the Ground (Cambridge, Mass.). Pissart, A., 1966 Le role géomorphologique du vent dans la région de Mould Bay, He Prince Patrick, Zeitsch.f. Géomorphologie 10, 226-36.

PO 145 Beach erosion at Durban MUNSAMY B. NAiDOO ex-Sastri College, South Africa The fury of the advancing spring tides, at times accentuated by powerful north easterly winds, is well seen at Durban (Republic of S. Africa) where a perennial problem of beach erosion has engaged the attention of experts. The purpose of this paper is to present my views on the Durban shoreline, derived from many years of close and continuous observations. The erection of groynes to restore a shore with wide beaches has been a recent attempt at re-establishing an eroded foreshore. Drift sands carried by a longshore current are no longer available and the groynes are in effect structures to trap the meagre supply of wavetransported particles from the south coast as well as sand made available by pumping operations. The groynes have succeeded in some measure to rehabilitate the beach but one wonders whether their construction at right angles to the shore does really solve the problem even though the structures are lacking in aesthetic appeal. It is best perhaps to consider the position of the 'Back Beach' a name by which the shore is known in Harbour Reports and to

early residents of Durban. It is recorded that in the year 1860 the bar at the Point had an incredibly low depth of twelve inches. Six years earlier, in 1854, a maximum depth of approximately eleven feet was registered. In its early stages Durban Bay was a lagoon receiving the discharge from the Umgeni, Umbilo, Umshlatuzan, Isipingo and Umlaas rivers which in former times built the alluvial flats on which the City stands. Mud flats appeared during tidal recessions and a network of channels was scoured during ebb tide. The harbour entrance was blocked by a sand-bar breached by a channel of uncertain navigability. To the north of the harbour entrance and at the south end of the beach lay a sand bank, the result of eddies caused by the deflection seawards by the Bluff of the longshore current. The growth of Durban and the importance of the harbour as a port led to engineering projects to prevent the northerly flow of sands from inundating the channel. Construction of piers intensified the scouring action during ebb tide and by continuing a south pier seawards, the longshore current with its load of drifting sands was

Géomorphologie I 85 pushed further out to sea. Whereas before these operations, the longshore current near the harbour entrance was confined to shallow waters, the deflection seawards must have caused the deposition of sands at greater depths thereby depriving the current of its normal load and intensifying its erosive action. As a surface current, the estimated average velocity is between half to one knot along the shore but on being directed to greater depths the effect of the current on the sea bed is negligible. Therefore, the vast accumulation of transported sands must remain practically undisturbed out at sea. But eddy currents generated on the leeward side of the south breakwater could create sufficient agitation of the sea bed to cause sand particles to be transported southwards and northwards by the longshore current. The effect of this is seen in the North Beach area beyond the Country Club where coarse beach sands such as appear on the Bluff lie in graded sequence. It is apparent thaf the north-bound shore current, deflected seaward by the extension of the south pier, flows towards the north beach in a wide arc while its littoral components drag the suspended material in a northerly direction. I believe that a curved North Pier running parallel with a curved South Breakwater of increased northerly orientation might have led to more satisfactory results in respect of the deflected shore current and the eddies on the leeward side of the North Pier. Perhaps, the offshore submarine topography determined the present orientation. In the deep water where wave oscillations diminish with depth, deposition of transported sand from the south decreases the load of the incoming wave and increases its erosive power. The zone of breakers now appear closer to the shore, wave height is increased and the swash advances further landwards. Beach restoration was initiated by the pumping at the South Beach of sand dredged from the harbour entrance. While this has raised the hopes of citizens as a solution to the disappearing beach sands, one striking weakness seems apparent in this procedure. The scouring action of the ebb tide has been intensified by the extension of piers and considerable amounts of argillaceous sand, fine grained in texture, flow outwards from the bay only to be trapped in the deep channel caused by dredging. It is sand of this type

that is being pumped and little wonder that so much of the south beach and portions of the north beach consist of it. In addition to discolouring the water, it is readily transported on windy days to the great annoyance of bathers. Such sands will cause dangerous shoals and channels affecting the safety of bathers. The effects of sand pumping are noticeable all along from the south beach to the vicinity of the Country Club. The fine grained sands do not constitute normal beach material and the powerful undertow during Spring Tides has undoubtedly redeposited much of them, affecting marine life on the sea bed. The Bluff is nature's groyne and the Anabella sand bank built up by the eddies on the north side of the Point was a feature of the south beach before harbour developments took place. Groynes are valuable structures to trap shore drifts along coasts devoid of breakers. They have fulfilled the expectations of engineers in many parts of the New England Coast. One wonders whether the Durban groynes are sufficiently rewarding and whether their spacing is not a contributory cause to increased erosion of the sea bed by the incoming waves. The swash and undertow cause distribution of beach deposits on the basis of weight and size of sand grains. But where the grains are uniformly small, the advancing swash is bound to pile the sand into irregular mounds with gradients greatly disturbed by the action of winds. Again, on the north side of the groynes powerful subsurface agitation of the waters adds to the risks of bathing in their proximity. For a systematic study of transported beach deposits recourse could be had to the practice of tracing shingle movement with tracer pebbles treated with radioactive isotopes. Barium 140-lanthanum 140 with a half life of twelve days has been used with great success along British coasts and there is no reason why this method, admittedly costly, should not be tried here for a more accurate mapping of sand movements along the sea bed, using geiger counters mounted on heavy metal sledges. Such a study is imperative for any successful scheme aimed at beach reclamation. Wellington, Physical Geography, 173-8. Wooldridge and Morgan,^ The Physical Basis of Geography, 346.

86 / Qeomorphólogy P0146 Formes particulières de l'érosion différentielle dans les tillites PIERRE ROGNON Institut de Géographie, Paris, France Les affleurements de tillite couvrent des surfaces réduites comparées aux superficies représentées par l'ensemble des roches élastiques. Ils ont cependant une extension appréciable sur les vieux socles qui ont connu les glaciations éocambriennes, ordoviciennes (Sahara) ou permo-carbonifères (fragments du Continent de Gondwana). Les reliefs de ces régions de tillite ont pour particularité d'être difficilement explicables sans une connaissance précise des mécanismes d'érosion et de dépôt intervenus au moment de la sédimentation glaciaire. En effet, les variations de faciès des grès ou des conglomérats et leur imbrication avec des lentilles de siltstones ou d'argilites présentent souvent la même complexité que celle qu'on rencontre dans les dépôts glaciaires quaternaires. On a déjà signalé, dans des tillites, certains reliefs énigmatiques provenant de l'exhumation d'ensembles gréseux homogènes et bien individualisés qui correspondent à des formes originelles de dépôt, comme par exemple les eskers d'âge permp-carbonifère dans la diamectite des Falklands (L.A. Frakes et J.C. Crowall 1967). La variété des structures sédimentaires dans les tillites est, en effet, beaucoup plus grande que l'individualisation de chenaux fossiles ('shoestrings') dans les grès d'origine fluviàtile. On connaît d'autre part, dans ces formations, quelques exemples spectaculaires de paléovallées glaciaires exhumées qui jouent un rôle dans la topographie actuelle ou dans le tracé du réseau hydrographique, par exemple en Australie méridionale ou dans le sud-ouest africain. Plus généralement, la fréquence des surfaces de discontinuité à l'intérieur des tillites accentue encore la complexité de la répartition des faciès: elle est en rapport avec les phases successives d'avancées des anciens inlandsis mises en évidence aussi bien dans la glaciation permo-carbonif ère que dans celle de l'Ordovicien supérieur. Dans les déserts, le rôle joué par ces facteurs paléogéographiques est très spectaculaire. Ainsi, au Sahara central, la variété déconcertante des topographies liées aux affleurements de la tillite tardiordovicienne s'oppose à la régularité et à la monotonie des grands plateaux de grès du Paléozoïque in-

férieur dans les auréoles sédimentaires entourant le bouclier précambrien du Hoggar et, au nord, celui des Eglab. Dans ces auréoles sédimentaires dont la synthèse géologique a été publiée récemment (Beuf et al. 1971 ), la tillite affleure sur près de 3500km de longueur et sur une largeur variant de quelques centaines de mètres à plus de 150km (nordest du Hoggar). Sous le climat hyperaride actuel (moyenne annuelle des pluies: 10 à 50mm), l'action du ruissellement épisodique et du vent sur des roches dépourvues de sol et de végétation favorise le dégagement des moindres différences lithologiques à l'intérieur de la tillite. Or les différences de comportement de ces roches détritiques vis-à-vis de l'érosion sont en rapport étroit avec la disposition originelle du matériel glaciaire ou périglaciaire. En effet, celle-ci n'a guère été modifiée par les déformations tectoniques qui se sont traduites presque essentiellement par des réseaux de fractures et de diaclases sur cette portion du craton africain. La diagenèse elle-même a souvent respecté les limites des unités sédimentaires originelles, ce qui favorise leur individualisation par l'érosion actuelle ( Rognon et al. 19 68 ). Ce rôle particulier de la diagenèse est mis en évidence lorsque les anciens planchers glaciaires sont dégagés en grandes surfaces pseudo-structurales. Ces surfaces d'érosion glaciaire, authentifiées par de fréquentes stries, cannelures, crescentic gauges etc. remarquablement conservées, ont été ferruginisées ou parfois silicifiées de façon préférentielle parce que la circulation des eaux a rencontré là une brusque discontinuité de porosité entre le matériel compacté par l'inlandsis sous la surface et le matériel plus meuble situé au-dessus. Ainsi, l'érosion actuelle a exhumé des champs de 'glacial flutings' (à ne pas confondre avec des yardangs de grande taille! ) à la bordure sud du Hoggar et des surfaces moutonnées au nord des Eglab. Au nord du Hoggar, on assiste à l'exhumation partielle d'une série de paléovallées glaciaires profondes de 100 à plus de 300 mètres, qui constituent des sections anormalement élargies sur certaines vallées actuelles. D'une façon plus générale, la succession de plusieurs phases d'érosion glaciaire

Géomorphologie I 87 lors des fluctuations de l'inlandsis ordovicien se traduit par une topographie en gradins, assez caractéristique à la bordure nord des Eglab par exemple on dans le Tassili d'In Ebeggi (Biju-Duval et Rognon 1971). D'autres surfaces pseudo-structurales correspondent à des paléotopographies d'accumulation fidèlement exhumées par l'érosion actuelle lorsqu'elles étaient recouvertes par des siltstones ou des argilites. Les détails de leur morphologie ancienne sont parfaitement conservés, surtout vers le haut de la tillite, c'est à dire après les dernières avancées de l'inlandsis ordovicien. On peut citer ainsi: dans le Tassili des Ajjers, l'exhumation d'anciens chenaux d'un vaste sanour lié à la dernière des glaciations (Liandovery inférieure?) qui se traduit par la mise en relief de plusieurs dizaines de grands 'shoestrings' méandriformes ou anastomosés, longs de plusieurs kilomètres. Ce paysage de chenaux exhumés s'étend sur plus de 100km d'ouest en est. Les grès intermédiaires, moins résistants, sont constitués par des remplissages périglaciaires à fentes de glace et à grandes rides nivéo-éoliennes. Un peu plus au nord, au toit de la tillite ordovicienne, on observe une quarantaine de grandes rides gréseuses longues de plusieurs centaines de mètres, larges de 50 à 150 mètres et hautes de quelques mètres. Ces curieuses rides parallèles seraient des vestiges des rivages successifs d'une région soulevée par glacio-isostasie à la fin de la glaciation ordo vicienne et qui ont été fossilisées par les argiles marines siluriennes. - Dans le Mouydir, sur une surface d'accumulation périglaciaire, l'érosion actuelle a exhumé une série de formes circulaires de 200 à 400 mètres de diamètre, qui sont d'anciens pingos ordoviciens parfaitement conservés. On comprend, après l'énumération de ces quelques exemples spectaculaires, pourquoi la morphologie de ces grès ordoviciens était restée incompréhensible jusqu'à la mise en

évidence du caractère glaciaire de cette formation (Beuf et al. 1971). Il existe des quantités d'autres exemples de formes étranges qui sont des chenaux sousglaciaires exhumés, des amas de moraine, des dépressions arrondies provenant de paléo-kettles, etc. Ainsi au nord des Eglab, nous avons pu interpréter un bourrelet long de près de 15km, en forme de lobe, comme un ancien arc de moraine de poussée. Il est évident que devant de telles formes, le géomorphologue qui ne s'intéresse pas à la paléogéographie de ces roches détritiques est totalement désorienté! Les relations entre la topographie actuelle et les conditions paléogéographiques d'érosion et de dépôt sont particulièrement frappantes dans le cas des tillites, surtout sous climat désertique actuel. Mais le but de cet exposé est d'attirer l'attention sur l'importance de ces facteurs paléogéographiques (paléoformes exhumées le long des surfaces de discontinuité, 'corps gréseux' correspondant à des entités sédimentaires originelles), pour l'explication morphologique dans tous les types de roches détritiques. Beuf, S., B. Biju-Duval, O. de Charpal, P. Rognon, O. Gariel, et A. Bennacef, 1971 Les grès du Paléozoïque au Sahara (Paris). Biju-Duval, B., et P. Rognon, 1971 Phénomènes glaciaires et périglaciaires dans POrdovicien supérieur au sud du Hoggar, Abstracts* du VIIIe Congrès internat. Sédimenîol. (Heidelberg), 9. Frakes, L.A., et J.C. Crowell, 1967 Faciès and paleogeography of the Late Paleozoic Lafonian Diamectites, Falklands, Geol. Soc. Am. Bull. 78, 37-58. Rognon, P., O. de Charpal, B. Biju-Duval, et O. Gariel, 1968 Les glaciations 'siluriennes' dans l'Ahnet et le Mouydir (Sahara central), Publ Serv. Géol Alg. n 38, 5381.

P0147 L'aggravation de l'érosion dans l'Ouarsenis (Algérie)

DJILALI SARI

Université d'Alger, Algérie

On n'insistera jamais assez sur l'érosion accélérée au Maghreb, dans la conjoncture actuelle, en général, et dans l'Algérie occi-

dentale montagnarde, en particulier. Dans ces montagnes déboisées et relativement peuplées, les phénomènes érosifs sont devenus

88 / Geomorphology classiques, et même l'une des composantes dominantes des paysages. L'inverse en serait une exception de plus en plus rare. Cependant, on ne prend conscience de leurs effets à moyen et à long terme que par un examen continu et régulier du terrain au cours d'une période, aussi courte soit-elle. C'est ainsi que dans l'Ouarsenis, région que nous suivons de près, nous avons noté au cours de ces derniers mois, en plus de nombreux signes attestant une reprise générale de l'érosion (Benchetrit 1954; Sari 1971), des phénomènes fort impressionnants. Nous nous limitons dans cet exposé à la présentation de trois exemples de glissements: les glissements liés aux merjate, les solifluxions généralisées et les complexes de glissements. I. LES GLISSEMENTS LIÉS AUX MERJATE

Ces phénomènes se signalent avant tout par la présence d'une ou plusieurs merjate (= pluriel de merja: petite étendue.d'eau) de dimensions variables suivant les conditions topographiques locales. Alimentées dès les premières pluies d'automne, ces merjate persistent tardivement. Cependant, il ne faut pas exclure les apports par infiltration semisuperficielle. La difficile infiltration des eaux et leur stagnation sont liées en grande partie à la lithologie, à la prédominance des argiles montmorillonignes (Laboratoire 1964). Tout autour des merjate et jusqu'à la base du versant, apparaissent les signes caractéristiques des glissements classiques. Mais ces glissements se distinguent aussi, dans les sections aval et amont, par de grands affaissements qui deviennent à leur tour le siège de nombreux phénomènes secondaires: accumulation d'eau, bourrelets, généralisation de fissures. Par ailleurs, ces phénomènes trouvent dans les versants modelés par la solifluxion ancienne, un terrain de prédilection. On assiste non seulement à une réanimation mais aussi à une intensification générale des actions antérieures. Aussi les déformations en surface sont-elles considérables, comme le montrent notamment les beaux exemples du Haut-Bassin de l'Oued Fodda, caractérisé par une topographie très vallonnée (800m en moyenne ). Ainsi, parmi les facteurs directement responsables de ces glissements, il convient de souligner le rôle de la topographie héritée, la lithologie et surtout l'intensité des précipitations, d'une part, et les surcultures annuelles de pentes et de versants, d'autre part (Sari in préparation).

II. LA SOLIFLUXION GÉNÉRALISÉE

Ces phénomènes affectent de nombreux secteurs des deux versants du Massid. Ils se signalent non seulement par les phénomènes classiques (fréquence de cavités plus ou moins profondes, reliées par des ravines souterraines), mais aussi et surtout par des sortes de grands affaissements dont les axes sont grossièrement parallèles aux fonds des vallons. De plus quand on les examine de près, les formations semblent passer, durant les grosses pluies, de l'état solide, à l'état presque fluide, avec une décroissance sensible du volume total. Ceci est particulièrement mis en évidence après la saison pluvieuse. Par ailleurs, c'est également à cette période précise de l'année que l'observation devient fort intéressante, pour parvenir à une première approche d'interprétation du phénomène d'ensemble. Ainsi à la mi-juin, nous avons été attiré par les faits suivants, le long de l'Oued el Merdja, immédiatement en aval de Sendjes, ex-Bougainville (Carte 1953), Un contrastre très net est à souligner entre les versants qui sont dans l'ensemble très peu élevés (de 10-20 à 30-35m) et de pentes variables (de 10 à 30° suivants les secteurs) et le fond. Celui-ci est encore humide comme l'attestent la présence de végétation encore verte, non loin des chaumes, et un sol très imbibé d'eau. De plus le fond du vallon est assez plat et recouvert dans la section avale, par une couche blanchâtre de cristaux, provenant de l'écoulement latéral, au niveau de la base des versants, couche qui ne semble point provenir du fond même. Par ailleurs, deux autres faits non moins importants sont à préciser au sujet des deux versants. Ceux-ci sont régulièrement ravinés et surtout accidentés par des cavités de dimensions variables, résultats des phénomènes de solifluxion qui sont très répandus dans la région et qui sont à l'origine de la formation de ravines. En effet, l'élargissement de proche en proche des cavités dues à la solifluxion et leur approfondissement assez rapide causé par des averses rapprochées expliquent la fréquence des ravines et leur multiplication soudaine. Ainsi les affaissements importants soulignés ci-dessus sont, sans aucun doute, à mettre en relation avec l'altération des roches, la dissolution de certains éléments et l'entraînement de ces derniers sous l'effet des eaux d'infiltration. Ces dernières réapparaissent à

Géomorphologie / 89 la base des versants comme l'attestent d'une part, la conservation tardive de l'humidité, et d'autre part, la formation de la couche blanchâtre des éléments dissouts puis réprécipités. De plus la prédominance de la solifluxion souligne bien l'ampleur de la dissolution et de l'altération des formations en place, d'autant plus que ces dernières sont gypseuses et renferment des pointements triasiques, bien visibles. Dans ces phénomènes, les processus physico-chimiques jouent donc un rôle considérable après les chutes de pluies, dans la conjoncture actuelle. Ils finissent par imprimer à la topographie d'ensemble de nouvelles empreintes (élargissements des ravines par solifluxion et leur multiplication, petits affaissements). Cependant si l'on parvient assez facilement à reconstituer le processus général dans ses grandes lignes il n'en va pas de même quand on assiste à des glissements non seulement généralisés mais aussi fort complexes, comme dans le secteur situé au sud de Sendjes, et délimité par la route no 19 et sa bifurcation à l'ouest. III. LES COMPLEXES DE GLISSEMENTS

Dans de pareils exemples, on décèle la coexistence et l'enchevêtrement de plusieurs formes de glissements dont certaines sont apparentées aux familles décrites ci-dessous. L'échantillon suivant (carte précitée, x = 385.5, y = 303.4) nous en donne un aperçu. L'enchevêtrement des compartiments dans la section centrale, où l'on observe en particulier des déformations importantes dans les assises marneuses, constitue ici le fait le plus spectaculaire. Certaines de ces assises sont littéralement soulevées, d'autres tranchées, et d'autres, enfin, épanouies, en forme de chouxfleurs. Par ailleurs d'une saison à l'autre, on se rend compte de l'évolution. Après avoir bien noté des déformations évidentes à l'aide de repères, de nouveaux remaniements ont été observés par la suite: des biocailles, taillées dans un matériel relativement dur, jonchent la surface et un grand olivier perché sur la contre-pente nous donnent une idée de ces déformations. Quant au reste du glissement, il est classique, à l'exception de la section amont qui s'individualise par un grand affaissement que souligne en particulier une sorte de 'hort.' En revanche, la section

avale s'écoule jusqu'à la route pendant les pluies et bloque la circulation momentanément. S'étendant sur une pente moyenne de 25 à 30°, ce glissement atteint 200m et ne cesse de se prolonger en amont. Ainsi, cet exemple met en lumière l'interférence de plusieurs actions. Outre les processus déjà signalés, il y a lieu de faire intervenir dans des cas, et notamment dans la section centrale, des forces internes, des pressions et compressions, comme l'attestent nombre de déformations en surface. De l'examen de ces phénomènes, il y a lieu de souligner : la reprise générale de l'érosion et son extension à des secteurs jusque-là immunisés, préservés par les forêts et une occupation adaptée au milieu, l'arboriculture; la complexité des phénomènes et leurs effets de plus en plus spectaculaires; l'aggravation de ces phénomènes menace l'équilibre de la région à moyen et long termes. Le ruissellement s'accroit aux dépens de l'infiltration. La nécessité et l'urgence d'une révision des techniques de Défense et de Restauration des Sols (DRS) mises au point ces dernières années. Ces techniques ne sont pas toujours valables. De plus, la construction de nouvelles routes et l'ouverture de pistes exigent de plus en plus une reconnaissance approfondie du milieu, notamment de la géomorphologie. Les organisateurs du 22e Congrès International de Géographie. Baurens, J.; 1966 Eléments sur l'érosion dans le bassin versant de l'Oued Fodda, Ann. Alg. Geog. 1, 13-40. Benchetrit, M., 1954 L'érosion accélérée dans les Chaînes telliennes d'Oranie, Re. Geom. dy., 145-67. Carte d'Etat-Major au 1/50000, feuille no 106 (1953),Orléansville. Laboratoire d'Hydraulique de France, 1964 Résultats des analyses effectuées en vue de déterminer l'origine des sédiments de la retenue de l'Oued Fodda (rapport dact). Sari, Dj., La désorganisation de l'agriculture traditionnelle dans l'Ouarsenis, à paraître dans 'Etudes Rurales' (Paris). - 1971 Les villes précoloniales de l'Algérie Occidentale, lile partie (Alger).

90 / Geomorphology P0148 Some problems in the correlation of landslide movement and climate

M.J. CROZIER

Trent University, Canada

Any attempt to correlate climatic data with landslide movement, particularly for the formulation of predictive models, can result in only very limited success. The reason for this is both the extremely multivariate nature of the relationship (Fig. 1 ) and the highly variable, active nature of the system containing these variables. A model based solely on independent climatic variables and a dependent movement variable will be extremely restricted in its predictive capacity. Spectacular landslide movement so often occurs during rainstorms that it is a natural reaction to expect a reliable model for slope failure to depend heavily on climatic factors. What is more, these factors are much more readily measured than, say, those of a geological or biological nature. Consequently, many attempts at correlation have been made (Miyabe 1935; Campbell 1966; Jackson 1966; Crozier 1968; Prior, Stephens, and Douglas 1970). The concept has been put forward (Crozier 1968) that climatic factors play three distinctive roles in respect to landslide movement: formation of unstable slopes (preparatory factors), precipitation of localized slope failure (triggering factors), and perpetuation of already moving material (controlling factors). Incidentally, non-climatic factors may also fulfil these three roles. For example, a slope weakened by the undercutting action of a river (preparatory factor) may fail during an earthquake (triggering factor) and continue to move in response to vibration from passing traffic (controlling factor) (Roed 1966; Kerr and Drew 1968; Benson 1940). Nevertheless, observation indicates that climatic factors frequently fulfil these roles, particularly that of triggering. Of the 66 major regolith landslides studied in detail by the author (in eastern Otago, New Zealand) only eight can be listed as having been triggered by some factor other than a rainstorm (Crozier 1970). A reliable climatic model for the prediction of landslides would be most welcome, not the least because of the defence it would offer from natural hazards of this nature. Nonetheless, the contention is that this task,

certainly with present techniques, is fraught with almost insurmountable problems. One approach to finding a forecasting model is simply to record the rainfall characteristics immediately preceding triggering. This has resulted in a number of convincing contradictions. Wright and Miller (1952) state that slips are reported to be more common during the first heavy rain following a drought; on the other hand the four devastating periods of landslip occurrence in New Zealand, since 1956, have all been associated with wetter than usual seasons. Then again, Jackson (1966), working on the Eastern Hutt hills, records a prolonged dry spell prior to a period of slippage. To come up with a universally applicable generalization from local studies of this nature is extremely difficult. This realization is implicit in an astute observation made by Prior, Stephens, and Douglas (1970). Of antecedent weather conditions they said: 'While the amount and character of the rainfall in the period leading up to the instability is undoubtedly important in the preparation of slope for failure, each of the cases examined appears to have little in common.' The discrepancies resulting from the different attempts to relate antecedent climatic conditions to triggering are obviously due to local variation in the relative importance of non-climatic variables. Even if the effect of non-climatic variables could be removed it is unlikely that an accurate forecasting model could be obtained. As an illustration of the problems involved consider an attempt in the field to exclude non-climatic variables by selecting and monitoring a single slope of uniform soil, form, drainage, vegetation, and other environmental conditions. Climatic factors such as rainfall, temperature, humidity, and evaporation are measured and recorded prior to slope failure. The magnitude and relative effect of climatic conditions at the time of triggering provide a working hypothesis which is theoretically a predictive device, at least for the chosen slope and its uniform conditions. However, the model is already virtually useless as a forecasting tool and little better as a working hypothesis. Even if

Fig. 1. Slope stability model.

92 / Geomorphology it were possible to duplicate by experimentation the exact climatic factors that produced the original failure (not an easy task, as the Japanese tragically discovered) it would not be possible to reproduce the results exactly. Replication under these conditions is impossible because now there is the interference of non-climatic factors through a time, rather than areal dimension. Depending on the initial slope conditions, the post-failure slope material will be either more stable (for example, all the regolith may be removed in the original movement thereby exposing material of higher shear strength) or more unstable than the pre-failure slope material, since certain climatic conditions can leave a landslide in an unstable sequential stage of its development (Crozier 1969). A fair test of or even an accurate prediction from this kind of model could thus be expected only in reference to another hillslope of identical non-climatic conditions. Finding such a hillslope is a virtually impossible task, as the aforementioned studies indicate, and having to look for a situation to fit the model, rather than vice versa, is a highly dubious scientific pursuit at any time. Despite confusing results to date it might still be expected that the desired predictive model for triggering of movement would come from examining a number of landslide occurrences through space in conjunction with their related climatic events. This approach has a number of problems apart from that of determining local variation in the importance of non-climatic environmental variables. The single most important problem is that of time equivalence in movement and climatic events. Meteorological records are usually of sufficient quality to yield accurate dates for climatic events, whereas determining the date on which a landslide was triggered is^much more difficult :e^en? inline-;unlikely event that there were witnesses. Experience has shown that farmers, on whose property landslides may be present, tend to relate the occurrence to some pastoral event rather than a calendar date. Relative dates for triggering can be obtained in reference to features that can serve as chronological datum points. For example, a landslide that disrupts a fence (pipeline, building, or other structure whose date is known) is usually assumed to post-date the

period of fence construction and, conversely, one that has an undisturbed fenceline traversing it is assumed to pre-date this period. It is clear, however, that many landslides undergo periods of reactivation, so the law for relative dating just suggested need not apply to triggering but only to the most recent period of activity. Besides the better known techniques of radio-isotope dating (Eden 1967), pollen dating (Johnson 1964), and dating with airphoto coverages, the presence of loess, volcanic ash, or other dated deposit on the surface of rupture can also afford a relative date of occurrence. The problem with these last two techniques, as with palynological methods, is the possibility of secondary deposition onto the surface of rupture. Many of these problems can be avoided if a climate/movement model is constructed by correlating data from a continually moving landslide with closely monitored climatic factors. Its application as a forecasting device is now of course restricted to predicting variation in movement rather than the more socially significant instant of triggering. As can be deduced from Fig. 1, the most important factor to include in any model is the water content of the landslide material. This factor is linked with no less than seven recognizable and important climato-hydrological variables and has the most direct physical link with landslide movements. The likelihood of success of a model derived from measurement of moving material lies in the ability first to determine what constitutes water content and, secondly, to find suitable techniques for its measurement. The low degree of correlation frequently obtained from this last approach presents a challenging problem pointing to the necessity of unravelling the tangle of interrelated factors^ accounting for the time lags between climatic events and peaks of movement, and finding ways of precisely measuring the desired factors. Benson, W.N., 1940 Landslides and allied features in the Dunedin district in relation to geological structure, topography and engineering, Trans. Proc. Roy. Soc. New Zealand 70, 249-63. Campbell, A.P., 1966 Measurement of movement of an earthflow, Soil and Water

Géomorphologie I 93 2, 23-4. Crozier, M.J., 1968 Earthflows and related environmental factors of eastern Otago, /. Hydrol. (N.Z.) 7, 4-12. - 1969 Earthflow occurrence during high intensity rainfall in eastern Otago, New Zealand, Eng. Geol. 3, 325-34. - 1970 Mass-movement in eastern Otago, PhD thesis (unpublished ), U. Otago, New Zealand. Eden, W.J., 1967 Buried soil profile under apron of an earthflow, Bull. Geol. Soc. Am. 78, 1183-4. Jackson, R.J., 1966 Slips in relation to rainfall and soil characteristics, /. Hydrol. (N.Z.) 5, 45-53. Johnson, R.H., 1964 A study of the Charlesworth landslides near Glossop,

North Derbyshire, Trans. Inst. Brit. Geog. 37, 111-26. Kerr, P.P., and I.M. Drew, 1968 Quick clay studies in U.S.A., Eng. Geol. 2, 215-38. Miyabe, N., 1935 Study of landslides, Tokyo Imp. Univ., Bull. Earthquake Res. Inst. 13, 85-113. Prior, D.N., N. Stephens, and G.R. Douglas, 1970 Some examples of modern flows in north-east Ireland, Z. Geomorph. 14, 27588. Roed, M.A., 1966 Report on the Geology of the North Garneau District. River Bank Stability Study (Dept. Geology, U. Alberta). Wright, A.C.S., and R.B. Miller, 1952 Soils of South-West Fiordland, New Zealand Soil Bureau Bull. 7.

P0149

Comments on a neglected landform: the meteorite crater G.V. DOJCSAK University of Saskatchewan, Canada The science of landforms is geomorphology. Its aim is to understand the shape of the earth and to explain the processes at work on its surface. To date geomorphology has clarified the landforms produced by the constant interaction between endogenous and exogenous forces, considering all but one of the tremendous variety to be found on the surface of the earth. The exception is the meteorite crater. The origin of this feature is due to the mass and kinetic energy of a cosmic body - that is, it is extraterrestrial. The presence of such features was recognized towards the end of the last century, but because of insufficient proof their impact origin was only suspected by some and totally ignored by most earth scientists. They were treated as 'pseudovolcanic' structures with probable cryptovolcanic origin by some authors, and possibly meteoric origin by others. Their presence was regarded more as a curiosity than as of any major significance

until high-flying craft made possible the recognition of the larger meteorite craters. The stage for the scientific investigation of these features was not reached until the middle of the present century. Canada harbours perhaps the most complete assemblage of meteorite craters on the earth and has played a leading role in this research. As a result of the investigations, acceptance of impact cratering is gaining recognition. Yet, to date, most geographers neglect this landform. It is felt that with their recognized presence (both as minor and major landforms) and their origin scientifically proved, their recognition (as a unique and perhaps the most interesting landform) is overdue. In order to remedy this neglect they should be included in the proper categories of the landform classification and dealt with in texts dealing with the surface of the earth.

P0150

Karsts de type tropical sous climat tempéré PAUL FENELON Institut de Géographie de Tours, France Sur le versant septentrional des Pyrénées, en Quercy et en Périgord, des reliefs karstiques, sous forme de hums et de chicots,

rappellent ceux des régions tropicales : Asie du Sud-Est et Antilles. Ils paraissent dus à des variations latérales de la sédimentation,

94 / Geomorphology à des phénomènes de diagénèse et à des formations coralligènes. A la suite d'un soulèvement épeirogénique, les processus

d'érosion ont déblayé les zones meubles ou tendres, laissant en relief les zones massives, dures et peu solubles.

P0151 Geologic and physiographic control of individual karst landscapes in Cuba VLADIMÍR PANOS Polackeho University, Czechoslovakia Numerous geomorphologists now contend that climate is the most important factor in the process of landscape development and is able to produce certain specific types of landscape forms in particular climatic morphogenetic zones. In accordance with this concept, H. Lehmann, the eminent German karst authority, proclaimed many years ago that the characteristic type of karst landscape in both the constantly and seasonally humid tropics is cone karst or some of its varieties such as tower karst. His opinions were published in several studies of the tropical regions and they evoked so much attention that a great number of undistinguished, imitative papers dealing solely with cone karst throughout the world soon appeared in the climatic geomorphology literature. Surprisingly, the various authors pretended not to notice that in the regions studied there are also karst landscapes that differ substantially from that of cone karst. Finally, such narrowminded concentration on cone karst resulted in the conjecture that this attractive landscape is not only the specific but also the sole karst type in the tropics. Because of this questionable conclusion the term 'tropical karst' even became synonymous with 'cone karst.' The comprehensive karst research and geomorphic mapping in Cuba that was done as a team effort in 1964-7 clearly showed that cone karst is neither the sole karst landscape nor the most common and extensive one, even though all of Cuba has a uniform, seasonally humid, tropical climate. It became obvious that the naive climatic idea mentioned above differs from sober reality. It also contradicts the basic principles of climatic geomorphology that maintain that geologic factors, specifically lithologie variability and tectonic variability, may cause a considerable differentiation of exogenous landscape forms within individual climatic morphogenetic zones. Naturally the first publication advocating geologic and physiographic control of the development of the

Cuban karst cones known as mogotes encountered an unkind and often unwarranted reaction from some authors. Nevertheless, as a result of a comprehensive analysis of all Cuban karst landscapes, it was concluded that geologic and physiographic control is generally the most important influence on karst development as well as the major determinant of its considerable morphologic and hydrologie variability. The most important geologic factors are: ( 1 ) stratigraphie and hypsometric position of soluble rocks in relation to the insoluble ones; (2) thickness and distribution of soluble rocks; (3) geologic structure; (4) lithologie properties; (5) porosity and degree of lithification; (6) chemical composition; (7) thickness of beds; (8) type and density of joints and fissures of any origin; (9) type, velocity, and frequency of crustal movements that took place before, during, or after individual phases of exogenous development. The most important physiographic factors are: (1 ) relation of soluble beds to certain tectonic units that are the initial macroforms of a constructional landscape; (2) the position of the soluble beds in time relative to erosion and/or corrosion base-level; (3) isostatic and glacio-eustatic sea level fluctuations as well as resulting processes and changes; (4) duration of individual phases of exogenous development; (5) protective function of erosion-resistant duricrusts and calcareous evaporitic coatings; (6) influence of permeable sediments and weathered mantles; (7) type of vegetation cover and its density. The geologic and physiographic factors are complexly interrelated, and they control mechanical weathering, chemical decomposition (solution), soil formation, areal erosion and corrosion, linear and lateral erosion and corrosion of streams, and sedimentation - all of which are processes influenced by climatic conditions, such as the

Géomorphologie I 95 temperature and the annual radiation as well as the annual and diurnal distribution of precipitation. The Cuban insular platform consists chiefly of Laramide structures that include infolded relics of probably Acadian (Carboniferous) tectonic units, and it is widely fringed or mantled by post-Laramide sediments and structures. These structures are, from a geomorphic viewpoint, the initial endogenous or constructional basis of an exogenous landscape. Since much of all of these landscapes consists of soluble rocks, chiefly limestone, the exogenous landscape is characterized by karst landforms. Constructional macroforms that have the same or similar geologic characteristics clearly display the same or similar sets of corrosional, corrosional-erosional, and corrosional-suffosional mesoforms and microforms. Such forms constitute particular types of landscape that are differentiated because of physiographic factors into several subtypes. According to the relation between certain constructional macroforms and the prevailing influence of certain physiographic factors, six main types of karst and several subtypes all of which occur in a seasonally humid tropical climate-have been recognized in Cuba. Different developmental stages were recognized in some of them. These types and subtypes are: 1. Karst of coastal plains and lowlands: (a) inundated by the sea; (b) covered with marshes and brackish or fresh-water inland swamps; (c) emerged from the sea during Holocene time; (d) covered with thick and continuous, fluvial, deltaic, and mixed deposits; (e) covered with thin and discontinuous, chiefly fluvial deposits. 2. Karst of uplifted carbonate plateaus. 3. Karst of simply folded and faulted mountains, consisting of: (a) sedimentary rocks, chiefly carbonates; (b) sedimentary rocks, chiefly non-carbonates; (c) interbedded volcanic and both carbonate and noncarbonate sedimentary rocks; (d) intrusive rocks, commonly serpentinized. 4. Karst of complexly folded and faulted mountains, consisting of : (a) thick-bedded heterogeneous sedimentary rocks; (b) thinbedded heterogeneous sedimentary rocks; (c) carbonate schists and marble; (d) insoluble schists with isolated blocks, partings,

and lenses of soluble rock; (e) interbedded volcanic and carbonate and non-carbonate sedimentary rocks and intrusive rocks. 5. Karst of diapiric structures. 6. Littoral karst. The karst of diapiric structures was classified as a distinct karst type because of their specific initial constructional forms as well as their complex lithologie composition. The littoral karst was classified as a special karst type because it occurs not only in all kinds of constructional relief modelled by the sea's activity, but also in coral barriers, fringing reefs, and cemented calcarenitic bars. Furthermore, it is a result of processes different from those that form inland karsts. It may seem surprising to some that cone karst does not occur in the classification as one of the types of Cuban karst landscape. This is because cone karst must be regarded as a set of forms characteristic of rather advanced phases or stages of exogenous development of the landscape. It has been clearly shown that only some parts of certain constructional macroforms are able to achieve the morphology of cone karst. Consequently, the geologic and physiographic factors are again the determinant of the development of this very attractive karst landscape. With regard to these statements one must consider that even where there is climatic influence on karst development, the individual basic climatic morphogenetic zones include regions with entirely different endogenous and exogenous conditions that determine a priori not only differentiation of weathering and denudational processes but also differentiation of the consequent karst morphology and hydrology. These endogenous arid exogenous differences are of such obvious importance and often so very striking that they must not be ignored when studying the resultant sets of karst landforms. In order to achieve valid results in climatic morphogenetic classification of landscapes it is first necessary to study carefully all the non-climatic morphogenetic factors that might modify the active development of karst within individual climatic morphogenetic zones - and it is necessary to make such a study with more care than is required for non-karst terrains. It must be emphasized that this concluding statement is not a crusade against climatic karst geomorphology in general. On the

96 / Geomorphology contrary, it is only an attempt to rectify some of its false, yet ostensibly orthodox concepts and to stress the importance of variations in lithologie and structural influence in a given

climate as a key to correct classification and understanding of karst landscapes, specifically those in Cuba.

P0152 Les terrasses d'abrasion de la côte du Chili semi-aride R.P. PASKOFF Université de Tunis, Tunisie La côte du Chili semi-aride (30°-33°s) est remarquable par l'existence de terrasses d'abrasion marine étagées du type rasa dont l'ampleur et la netteté avaient attiré l'attention de Darwin lors de son voyage en Amérique du Sud. La baie de Coquimbo (30°s). Autour de la baie de Coquimbo, cinq terrasses, taillées dans la série gréseuse tendre (formation de Coquimbo) laissée par la transgression du Pliocène moyen à supérieur, s'étagent avec une étonnante régularité au-dessus de la plage actuelle. Elles sont séparées entre elles par des falaises mortes très nettes dont le pied se situe successivement à 5-7m, 15-20m, 3540m, 75-80m et 120-13Om (Fig. 1; les diagrammes non publiés ici). Aux alentours du 31 °s. La présence d'une seule grande plate-forme d'altitude variable donne à ce secteur son originalité. Vers l'intérieur des terres, cette vaste terrasse est limité par un abrupt, haut de plus de 500m, qui représente la retombée des chaînons côtiers. A l'opposé elle se termine par un escarpement de plusieurs dizaines de mètres de commandement qui domine l'Océan, soit directement, soit par l'intermédiaire d'une ou deux banquettes étroites. L'ampleur de ce niveau mérite d'être soulignée puisque sa largeur moyenne est de l'ordre de 3km. L'arasement, pour ainsi dire parfait, a été réalisé sur des roches granitiques d'âge primaire ou secondaire d'une part, sur des formations gréso-schisteuses anciennes, plus ou moins métamorphisées d'autre part. Des mouvements tectoniques ont affecté la rasa après son emersion: des gauchissements expliquent que sa hauteur au-dessus de l'océan varie d'un endroit à un autre; des escarpements de faille de rejet appréciable (de quelques mètres à quelques dizaines de mètres) créent à sa surface des dénivellations brusques (Fig. 2). A la hauteur du 32°s. Dans ce secteur les terrasses marines sont au nombre de trois. La

plate-forme supérieure s'abaisse d'environ 13 0-140m, en contrebas des chaînons qui la dominent de quelque 500m, jusqu'à 90-100m sur son bord externe; sa pente est de l'ordre de 2 pour cent. La terrasse moyenne est séparée de la précédente par un escarpement raide, d'une quarantaine de mètres de hauteur; elle s'incline de 35-40m à 15-20m, avec une pente voisine de 1.5 pour cent. Une falaise morte dont le pied se situe à 5-7m, la limite du bas niveau qui s'étend jusqu'à la plage actuelle. La largeur de l'ensemble est d'environ 3km. Ces plates-formes ont été taillées dans une série sédimentaire à faciès rythmique (alternance de schistes et de grès) d'âge triasique (Fig. 3 ). Différents types de preuves permettent d'affirmer que ces terrasses ont bien une origine marine. Preuves paléontologiques. On les rencontre autour de la baie de Coquimbo (30°s). Là, les cinq terrasses sont couvertes par des formations de plage épaisses de quelques décimètres, mêlant sables^ galets et coquilles, le plus souvent cimentées en dalle. Il est parfois possible de distinguer dans ces dépôts qui ravinent les grès tendres du Pliocène marin sous-jacent (formation de Coquimbo), ce qui revient à la transgression - un conglomérat comprenant des fragments arrachés au substratum - et ce qui a été laissé par la régression, généralement des sables peu caillouteux et stratifiés. Le foisonnement des fossiles atteste l'origine marine des terrasses. Les associations faunistiques des formations de plage corrélatives de l'abrasion de ces plates-formes sont très différentes de celles du Pliocène moyen à supérieur puisque 70 pour cent environ des espèces et sous-espèces de la formation de Coquimbo ne s'y retrouvent plus. Au contraire, elles sont par leur composition très proches de celles qui se rencontrent aujourd'hui le long du littoral et témoignent d'affinités à tendance fraîche. On peut ainsi dater les terrasses du Quaternaire

Géomorphologie I 97 et situer à la charnière du Pliocène et du Pleistocene, sinon rétablissement du courant de Humboldt, du moins l'acquisition de l'anomalie thermique qui le caractérise encore actuellement. Par suite de la persistance de cette anomalie pendant le Quaternaire, la faune littorale est restée à l'abri de toute influence extérieure; cet isolement explique ses traits accusés d'endémisme comme la diminution du nombre des genres et des espèces d'une part, la multiplication des individus d'autre part. Preuves sédimentologiques. Aux alentours du 31°s, les dépôts qui couvrent la vaste plate-forme littorale sont azoïques mais leurs caractères morphométriques et granulométriques, tout comme leur disposition, montrent qu'ils sont incontestablement marins. La meilleure coupe à cet égard est celle qui peut s'observer au km 269 de la route panaméricaine du Nord. On y observe, immédiatement au-dessus de phyllades infracambriens irrégulièrement tronqués, des galets et des graviers stratifiés horizontalement, très bien classés et remarquablement arrondis (ils sont désignés dans le parler local sous le nom de porotos qui signifie pois) auxquels se mêlent quelques blocs de phyllades. Il apparaît donc clairement que la roche en place a été rabotée par l'action de l'Océan qui a laissé un dépôt corrélatif typiquement littoral. Preuves géomorphologiques. Aux environs du 32°s, ce sont surtout des arguments d'ordre géomorphologique qui permettent d'être convaincu de l'origine indiscutablement marine des terrasses côtières. Ainsi, la plate-forme moyenne est hérissée de rochers aigus de plusieurs mètres de haut que le regard le moins averti n'a pas de peine à reconnaître comme étant d'anciens écueils. Quant à l'escarpement dont la base se situe à 5-7m et qui limite le bord interne du bas niveau, c'est bien une ancienne falaise morte puisqu'elle porte encore les marques bien visibles d'une action pas très ancienne de l'Océan: anfractuosités básales, grottes béantes, cordons de galets plaqués à son pied. Plusieurs facteurs permettent d'expliquer l'existence de terrasses marines remarquablement développées le long de la côte du Chili semi-aride. La houle du Pacifique Sud, forte et constante, constitue un agent d'érosion mécanique efficace. Son action régulière et continue

est renforcée par celle, épisodique, des ondes de tempête et des tsunamis. La topographie initiale de la marge continentale exerce aussi une influence. La côte du Chili semi-aride est d'origine tectonique. A la fin du Miocène ou au début du Pliocène se sont produits des mouvements verticaux de grande ampleur qui ont dénivelé une topographie d'érosion élaborée pendant l'Oligo-Miocène. D'un point de vue structural la marge continentale du Chili semi-aride est ainsi constituée par une série de blocs plutôt aplanis et plus ou moins effondrés. Or si l'océan trouve presque à son niveau un gradin déjà bien raboté par l'érosion continentale (Fig. 4), il va le retoucher et parfaire l'aplanissement en taillant un escarpement qui reculera au fur et à mesure du progrès de l'attaque des vagues. Cette falaise continuera ainsi à céder du terrain jusqu'à ce qu'elle vienne se confondre avec la faille qui limite le bloc situé en arrière. Elle prendra alors l'aspect d'une grande falaise et, parce que le volume à déblayer devient trop important, son évolution sera bloquée, surtout si l'on tient compte de l'instabilité chronique, provoquée par le glacio-eustatisme, du niveau marin pendant le Pleistocene. Chaque cycle océanique du Quaternaire pourra ensuite tailler sa propre plate-forme en contrebas de la précédente. Ainsi s'expliquent en grande partie les belles terrasses marines étagées qui accompagnent le littoral du Chili semi-aride sur presque toute sa longueur. La nature et la disposition du matériel rocheux de la marge continentale conditionnent aussi le travail morphologique de l'océan. Certaines roches se prêtent mieux que d'autres au 'coup de rabot' marin. Par exemple les séries sédimentaires paléozoïques à faciès rythmique (schistes et grès en disposition alternée), qui affleurent largement le long de la côte du Chili semi-aride, permettent de plus beaux aplanissements que les effusions volcaniques mésozoïques, probablement parce qu'elles sont de résistance moyenne à l'abrasion et qu'elles donnent aux vagues les instruments d'attaque nécessaires à leur action mécanique. De même les granités du batholite côtier, presque toujours réduits à l'état d'arène sur de grandes épaisseurs, n'ont pas offert une grande résistance à l'érosion littorale. Les séries anciennes, touchées par les

98 / Geomorphology tectoniques successives qui ont affecté le vieux socle, offrent des pendages très inclinés, voire parfois verticaux. Or de tels redressements se prêtent tout particulièrement au nivellement marin. Les roches granitiques, même là où elles n'ont pas été décomposées par la météorisation, permettent également l'élaboration de plates-formes parce qu'elles sont toujours densément fissurées; cette fracturation, souvent orthogonale, prépare le délogement des blocs par les vagues de tempête. La tendance au soulèvement épéirogénique de la côte du Chili semi-aride pendant le Quaternaire, plus ou moins marquée selon

les secteurs, a favorisé de son côté l'inscription des différents cycles marins. Ainsi les terrasses d'abrasion successives sont-elles plus clairement séparées et les falaises qui les limitent plus nettes que si le continent avait été stable. Herm, D., 1969 Marines Pliozân und Pleistozân in Nord- und Mittel-Chile, unter besonderer Berucksichtigung der Entwicklung der Mollusken-Faunen, Zitteliana 2. Paskoff, R., 1970 Recherches géomorphologiques dans le Chili semi-aride (Bordeaux).

P0153 Erosiona! features due to piping in Venezuela LEO PEETERS Berckem, Belgium Erosional features due to piping were studied on the southern border of Lake Valencia, Venezuela. They developed by combined action of human activity (a drainage canal), piping in heterogeneous material of a huge alluvial cone (gravel, sand, clay), and a tropical climate with a severe dry season of five to six months. The drainage canal was first deepened by erosion. Afterwards sections of the canal were widened by sapping and collapsing of

the walls due to piping until the original canal of width 1m and depth 0.5m became a deep canyon. The canal was constructed some 40 years ago. The zone of active piping is clearly visible on aerial photos taken in 1954, but at that time the canyon did not exist. In 1968 the canyon was noticed in the field. Hence, it takes, under these circumstances, less than half a century to transform parts of a small ditch into a canyon.

P0154 A simulation model for landslide prediction: an example from the coast of Normandy, France CHARLES L. ROSENFELD Centre de Géomorphologie, France Recent recreational development has occurred along the picturesque coastline of Normandy and has posed several problems analogous to those associated with other developing coasts in humid temperate regions. As along the southern California or New England coasts of the United States, construction has been increasing in areas of active sea cliffs where conventional soil mechanics or foundation engineering techniques have failed to predict the long-range frequency and distribution of hazardous landslides and slumps. In many cases damage and destruction of life and property have resulted from these unpredicted occurrences. In this paper, we shall attempt to demonstrate

a technique employing a simulation model of the coastal morphology of a region to delimit zones having a high hazard potential. It is common in the geophysical sciences to describe nature through idealized physical and mathematical models. The important factors in a particular situation are isolated for study, and associated phenomena which make it impossible to describe the relationships mathematically are excluded from the analysis. Simulation models serve three main purposes: (1) they develop a qualitative understanding of and appreciation for the physical mechanisms; (2) they can be tested experimentally, so that consistency of results with 'real world' data leads to greater con-

Géomorphologie I 99 fidence in the model and measurement techniques; (3 ) they can be used to predict the behaviour of variables. Mass movement may be related to any one or a combination of 'internal' or 'external' causes, or 'triggers.' Thus, the model must be adjusted by measurable parameters to environmental characteristics in such a way as to account for all possible determinants. Following a damaging landslide at the Semaphore near Longues-sur-Mer, France, in the spring of 1970, the Centre de Géomorphologie made a study of the causes of the mass movement. The results of this investigation provided the basis for the construction of this simulation model. The primary cause of the Tongues' slide was the saturation of the basal marls beyond their plastic limit, thus triggering their deformation under the weight of 40m of massive limestone overburden. The mechanics of this saturation process, however, involves a complex system of tectonic fracturing, desiccation fissuring, basal sapping by marine action, solution of overburden, and gelifraction. The resulting occurrence depends on a variety of interactions and 'feedback' effects among the geomorphic processes. A major obstacle to the formulation of quantitative simulation models in geomorphology has been the failure of conventional mathematical models to incorporate fully the quantitative effects of interactions between processes. Ahnert (1971) has described a computer simulation and thereby demonstrated the ability of the high-speed computer to incorporate 'feedbacks' from quantitative field measurements into a dynamic model of slope formation. Full quantification of such a model requires that even processes for which actual rates are not directly known be quantitatively defined. Thus the model is modified to incorporate quantitative measurements designed to provide accurate estimates of certain process rates. Also incorporated is a subroutine which replaces the action of stream downcutting with one of marine erosion that simulates basal sapping along a wave cut platform. In order to describe the static properties of the geologic materials (rri) the following relationship is quantitatively examined : m — f (Wc, Wm, Dp, Fc), where Wc = rate of chemical decomposition, W^ — rate of mechanical disintegration, Dp = diameter of

residual particles, Fc = cohesive force between particles (after Peltier 1954). Large unweathered samples were collected from each lithologie unit in the study area. Quantitative indices were derived through the following series of tests. The shear strength of each sample was measured with a pneumatic shear press. Since the major action of chemical weathering is the dissolving of calcium carbonate, the solubility coefficient of each sample was measured through controlled calcimetry. The sensitivity of the samples to gelifraction was used as the index of mechanical disintegration. Each sample was subjected to 100 rapid freeze/thaw cycles in the cryoclastic laboratory and the granulometry of the debris produced was plotted against iterations. These measurements were related to the porosity/ permeability characteristics determined by vacuum immersion and a mercury porosimeter. Finally, the residual materials were assigned Atterburg limits for liquidity and plasticity as a measure of cohesion. It was assumed that these measurements represented reasonable estimates of the static characteristics of the rock materials and therefore the relative resistance of each lithologie unit was adjusted in the model accordingly. Local adjustments were made to reflect the intensity of fissurization due to tectonic warping. Following field examinations of eight characteristic profiles made at selected points along the coast (after Elhai 1965), adjustments of parameters controlling the mode of mass movement were made in the model, based on: (1) the shear strength of the supporting material; (2) the amount of basal sapping (marine erosion), reflected in the amplitude of the tides and the exposure to direct wave action; and (3) the amount and characteristics of residual material. Thus, the COSLOP model (Ahnert 1971) was qualitatively adjusted to varying environmental circumstances by providing multiple options for combinations of the process modes and by 'tuning' the model quantitatively by the determination of relative values of parameters which control the rates of processes. In this manner it is possible to test field hypotheses of slope development. Correct evaluations produce synthetic slope profiles that agree in all essential characteristics with the actual landform the hypotheses seek to

100 / Geomorphology explain. Subsequent iterations of the computer simulation predict the future rate and form of coastal landform development, thereby indicating areas with high landslide frequency and potential. Efforts are continuing to verify the statistical accuracy of the simulation. Where synthetic profiles approximate natural slopes, a multiple regression program is used to determine the degree to which the model 'varies' from reality. Using a program written by Ongley (1970) the profiles are being broken down into slope components (i.e. crest, free face, debris slope) which are independently described as a series of rectilinear segments. Thus the model profiles are quantitatively contrasted with reality to determine their degree of correspondence so that existing variance may be located and eliminated by adjustment. There has, unfortunately, been a tendency for physical models to be applied quite out

of context. It must be emphasized, therefore, that a physical model is not a law of nature. The basic assumptions must be clearly recognized before intelligent use may be made of the resulting predictions. Andre Journaux; L.C. Peltier; Jacob Aghassy; Oswald Schmidt; University of Pittsburgh. Ahnert, F., 1971 A General and Comprehensive Theoretical Model of Slope Profile Development (Occasional Papers in Geography, U. Maryland, No. 1 ). Elhai, H., 1965 La Normandie occidental entre la Seine et le golfe normande-breton, étude morphologique (Bordeaux). Ongley, E., 1970 Determination of rectilinear profile segments by automatic data processing, Zeit. fur Geomorphol. 14, Heft 4. Peltier, L.C., 1954 Quantitative geomorphometry (unpublished manuscript).

P0155 Complex rapids (sula complexes) in tropical rivers J.l.s. ZONNEVELD Rijksuniversiteit Utrecht, Netherlands The rapids in many tropical rivers are intrinsic parts of labyrinthic systems; they occur together with anastomozing river branches and islands. In the Guy anas the individual rapids are called 'sulas'; for the labyrinthic systems the term 'sula complexes' could be used. Several factors play a role in the formation of sula complexes : ( 1 ) differential chemical weathering, giving rise to an irregular weathering front in the subsoil over which the river is flowing, followed by subsequent epigénesis; (2) the comparative poverty of abrasive materials in tropical rivers, due to the same chemical weathering and the dense vegetation cover protecting the hillslopes; (3 ) the alternation of savannah and humid forest conditions, causing changes in debris load and discharge and thus giving the river

alternative possibilities in terms of eroding hard rock or removing soft material; (4) the precipitation of ferromanganese compounds, indurating sandstones and causing the sula conglomerates and breccias. The phenomenon 'sula complex' is not restricted to tropical rivers. Morphographically the St Lawrence River in the area of Montreal, for instance, shows the same characteristics. In this case the pattern originated as a result of the (relative) lowering of the water level of the former Champlain Sea. In other cases (for instance, the rapids near Bradley Airport in the Connecticut River, USA) the river hit hard rock masses when lowering its channel into the soft bottom sediments of a former lake (in this case, Lake Hitchcock).

P0156 Effects of lithology and time on slope characteristics NICHOLAS BARiss University of Nebraska at Omaha, USA The purpose of this study is to investigate statistically the effects of lithology and time

on certain slope development tendencies in small stabilized valleys, using measurements

Géomorphologie I 101 of selected profile characteristics. Because research is still in progress, only part of the results are presented here. Valley profiles were measured in the field in Wisconsinan loesses and in non-loessic sediments which are Upper Miocene or younger. The non-loessic sediments are geologically unconsolidated terrestrial deposits composed of sand and gravel which arc clayey, partly cemented, and calcareous (mainly O gállala and Sheep Creek-Marsland formations; Condra and Reed 1959). Slope units (Young 1964, 17), comprising a profile, were measured with Abney hand level and steel tape as rectilinear portions between terrain breaks (1° or larger); this method was adopted because over 80 per cent of the profiles are angular rather than smoothly curved. Only the convex parts (Young 1964, 19) of the profile, measured as rectilinear units, are considered for analysis. Depending upon whether the measured valley has been subjected to one or two phases of erosion, each profile has two (2 single) or four (2 lower plus 2 upper) convexities. Similarly, each half profile, i.e., one side of the valley, has one or two convexities respectively. Single and lower convexities are grouped together in the analysis. Within each convexity, the following units were selected for analysis: 1. Maximum slope represents the beginning of stabilization (see criteria of sampling below). This rectilinear unit is considered to be affected by recent basal removal in the upper convexity and existing basal removal in the lower. (Basal concavities from slope deposits are absent or poorly developed in most of the small valleys of the sampling area. ) Minor slumps or landslips appear on the maximum slope, particularly in the case of loess. 2. Stable slope (term from Young 1970, 588) is the rectilinear unit immediately above the maximum break representing a more advanced stage of stabilization. This unit has no slumps or landslips. (The maximum break is defined here as the maximum difference in the angles of inclination between two adjacent rectilinear units within the conyexity.) By the use of the SPSS system of computer programs (Nie et al. 1970), frequency distributions of the maximum and stable slopes were determined whose arithmetic means, standard deviations, and distributional

tendencies (i.e. the degree of best fit of the normal curve to the observed distributions as indicated by the chi-square test) were included in the analysis. Slope development in loess versus non-loess was investigated by a quantitative comparison of the lower convexities in terms of the above parameters; a similar comparison of lower and upper convexities in the same sediment indicates the effects of time. Since slope development is extremely complex because of the large number of controlling factors, an attempt was made to reduce the number of variables as much as possible. Also, this study is concerned with slope characteristics after stabilization rather than regional slope values. As a result of these considerations, the following criteria were used in the selection of the valleys: 1. The areas of sampling, all covered by 7Î2 foot topographic maps, are located west and south of the Sand Hills, in the dissected plains and hills of western Nebraska and northwestern Kansas; the climate of this grassland environment is semiarid with 350500mm of annual precipitation. 2. Land use is almost exclusively pasture for the profiles involved. 3. Topographic criteria: (a) Valleys included in sampling are cut into uplands, either dissected or undissected, and have a definite linear shape, ( b ) Small valleys of first or second order without permanent streams were included. Channel incision into the flat bottom is either small (less than one metre) or absent. Sections with evidence of undercut slopes caused by a meandering intermittent stream were avoided; thus in the case of valley asymmetry, the cause should be other than unequal basal removal, (c ) In order to avoid large variations in the base level control, a local relief of 15-50 metres was established for consideration. (Local relief is defined here as the difference in elevation between the upland above the measured valley and the confluence of it to the next valley. ) Bluff topography of a major stream was avoided by the exclusion of those valleys which are located within 800 metres of the edge of the floodplain. 4. Since the purpose of this study is to analyse slope characteristics in stabilized valleys which originated from the last or last two phases of erosion, valleys whose maximum slopes exhibit scarplets higher than

TABLE 1

Slope angle (in degrees)

LOESS

Upper convexity Lower convexity NON-LOESS

Lower convexity

Log-tan of slope angle

Sample size

Mean

Standard dev.

Distr. tendency

Mean

Standard dev.

Distr. tendency

Maximum slope Stable slope Maximum slope Stable slope

68 42 147 146

19.0 7.7 29.0 9.1

5.77 3.76 5.99 3.60

Normal Right-skewed Normal Right-skewed

-0.48 -0.92 -0.26 -0.83

0.14 0.24 0.11 0.19

Normal Close to normal Left-skewed Normal

Maximum slope Stable slope

97 90

22.5 8.7

5.39 3.42

Normal Right-skewed

-0.39 -0.85

0.12 0.18

Normal Normal

Géomorphologie I 103 TABLE 2 Slope units compared

Difference between means

Effect of lithology: loess versus non-loess (lower convexities only) 1. Maximum slopes 2. Stable slopes (in log-tangents)

Significant Not significant

Effect of time: lower versus upper convexities in loess 1. Maximum slopes 2. Stable slopes in (log-tangents)

Significant Marginally significant

0.5m were excluded as unstable. Broad, 'graded' valleys which probably do not represent the last phase of erosion were also excluded by establishing a minimum value of 10° for the maximum slope and 5° for the maximum break. The above values, admittedly arbitrary, were estimated from the former field work of the author. (The prestable phases of valley formation in loess have previously been investigated by the author; see Bariss [1968, 1971].) From a population which met the above criteria, valleys were selected by random sampling. Results of the frequency distribution analysis of the maximum and stable slopes in 312 convexities are given in Table 1 in terms of arithmetic means, standard deviations, and distributional tendencies. In order to provide the basis for a meaningful /-test, the logarithms of the slope tangents (Speight 1971 ) were also calculated, thereby facilitating the determination of normal distributions for the stable slopes which, in terms of angles, display right-skewed distributions. In all cases, distributional tendencies were checked by the chi-square test. For twenty-six upper convexities and eight lower convexities, the stable slope was not measured because of the difficulties in determining the maximum break in the field. The statistical significance of the differences between the means, estimated by the Mest at the 1 per cent level of significance, is summarized in Table 2. Concerning the designation 'marginally significant,' in the case of the difference between the means of the stable slopes of the upper and lower convexities in loess, the observed value of '/' (2.5) is between the one and two per cent levels of significance (2.58 and 2.33) but closer to the one per cent.

The regular, mainly imimodal frequency distributions and the small variations in the standard deviations within the populations of the maximum and stable slopes seem to reflect the relatively short period of time between stabilization and advanced gradation in valley development considered in the analysis. The standard deviations of the slope angles are dependent upon slope steepness and stability but unaffected by time and lithology. The low values and narrow range of all logtangent standard deviations seem to support Speight's findings (Speight 1971, 309) and to de-emphasize again the effect of time and lithology upon this parameter. The association of normal distributions with the maximum slopes (early stage of stabilization), together with the poor development of the basal slope deposits, suggests some 'regulating' effect of the basal removal which might have caused a symmetrical scatter of the values about the mean. Rightskewed distributions, however, are associated with the stable slopes, which tend to maintain their slope angles and are subjected to less change. Distributions of the log-tangent values tend to be more uniform, i.e., all but one are normal. Time and lithology do not seem to influence distributional tendencies. Standard deviations and distributional tendencies of the slope angles might be considered useful indicators of stability. Whereas the arithmetic means of the stable slopes (in terms of log-tangents) do not appear to be affected by lithologícal differences, the significantly higher values of the maximum slopes of the lower convexities in loess might be attributed to ( 1 ) the greater diversion of surface waters by piping and a high degree of permeability, (2) higher cohesion, and (3 ) the role of deep gullies associated with steep, debris-controlled slopes at the

104 / Geomorphology pré-stable phase of valley formation in loess (Bariss 1971). Comparison of the arithmetic means between the lower and upper convexities in loess suggests that both stable and maximum slopes are affected by time, the latter ones much more significantly. The model of reclining slope retreat is also confirmed here. It might be concluded that whereas the maximum slopes seem to be controlled both by time and lithology, the stable slopes (expressed in log-tangents) appear to be affected, to a lesser extent, by time only. Senate Research Committee of the University of Nebraska at Omaha; Vera Bariss; Lee C. Bush; John F. Shroder; John Sumstad. Bariss, N., 1968 A comparative landform study of selected loess areas in the Missouri

River basin, chap. 6 in Loess and Related Eolian Deposits of the World, ed. by C.B. Schultz and J.C. Frye (Proc. vnth INQUA Cong.), 81-99. - 1971 Gully formation in the loesses of central Nebraska, Rocky Mi. Social Sci. J. 8 (2), 47-59. Condra, G.E., and B.C. Reed, 1959 The Geological Section of Nebraska (U. Nebraska Conservation and Survey Div., Lincoln, No. 14A). Nie, N.H. et al., 1970 Statistical Package for the Social Sciences (New York). Speight, J.G., 1971 Log-normality of slope distributions, Z. Geomorphol. 15 (3), 290311. Young, A., 1964 Slope profile analysis, Z. Geomorphol., Suppl. 5, 17-27. - 1970 Concepts of equilibrium, grade, and uniformity as applied to slopes, Geog. J. 136, 585-92.

P0157 Principaux résultats géomorphologiques du projet Hudsonie ANDRÉ CAILLEUX Université Laval Canada Mis sur pied par M. Louis-Edmond Hamelin, ce projet a pour objet l'étude multidisciplinaire coordonnée d'aires de recherches bien délimitées, échelonnées sur 15 degrés de latitude, en Ontario, Québec et Territoires du Nord-Ouest, le long de la rive est de la Mer d'Hudson et au-delà. La concentration de recherches variées sur une même aire permet une économie de moyens et une meilleure étude des interactions entre milieux physique, biologique et humain. A Poste-de-la-Baleine (55°N, température moyenne annuelle —4.6°c) Palbedo de la neige compactée est de 77 pour cent et passe à 85 pour cent quand s'y ajoute une couverture de neige fraîche (C. Wilson). Le socle archéen granitogneissique arasé est recouvert par du Protérozoïque - calcaire à Stromatolithes et basalte - en pente douce (5 à 6° ) vers la mer d'Hudson (S. Biron). L'inlandsis labradorien a recouvert la région au Quaternaire, mais il ne reste que très peu de till (J.-P. Portmann) parce qu'après le retrait du glacier la mer de Tyrrell a envahi et délavé la région, encore déprimée par la surcharge glaciaire (glacio-isostasie). Puis, comme suite à la décharge, les terres se sont soulevées, portant jusqu'à près de 300

mètres d'altitude les traces d'action marine, notamment les cordons littoraux faits de blocs arrondis et homométriques. Au cours de ce retrait, le type de côte a changé: au début, à l'est, sur TArchéen, côte à skjaer avec des centaines d'îlots; puis, plus à l'ouest, à l'approche de la couverture protérozoïque, côte rectiligne soulignant la pénéplanation; enfin emersion progressive des cuestas protérozoïques (L.-E. Hamelin, etc. ). Des hommes préhistoriques ont occupé les champs de blocs soulevés (traces d'habitations, outillage). Les sables apportés par la Grande rivière de la Baleine forment des cordons littoraux soulevés et de belles terrasses qu'entament des glissements spectaculaires favorisés par les argiles sous-jacentes et peutêtre par le dégel de la glace du sol. En tout cas, sous la terrasse de l'aéroport, le sol gelé permanent a été rencontré et dans certaines tourbières il y a des países atteignant 7m de haut, et des mares rondes résultant de leur fusion. Des concrétions, les unes ferrugineuses (alios), les autres calcaires, se sont formées dans les sédiments postérieurs au départ de l'inlandsis, donc sous un climat analogue à l'actuel, qui est nettement froid» Dans les concrétions calcaires, J.P. Adolphe

Géomorphologie I 105 a décelé des bactéries vivantes, qu'on voit remuer sous le microscope. Certains fragments plats de calcaire protérozoïque ont subi, sur leur face exposée à l'aire, une corrosion qui leur a donné un aspect vermiculé rappelant à s'y méprendre celui que, dans la bordure du Sahara, on attribue à la rosée. Près du littoral, le vent a remanié les sables des terrasses ou des cordons littoraux, d'où des cuvettes de déflation et des dunes paraboliques et à caoudeyres. Les dépôts nivéo-éoliens - accumulation de neige et de sable mélangés ou en lits alternants - donnent lieu à des formes annuelles, donc éphémères, mais très remarquables: bordillons littoraux, reposant sur des glaçons et atteignant 3m de haut; puis, quand la neige fond et que le sable restant se concentre pardessus ce qui en reste, cônes pointus et mamelons doux crevassés (Cl. Rochette et A. Cailleux), boulettes de sable, micromoraines de derivation et, sous les surplombs, microcratères d'impact de gouttes d'eau de

fonte, pastilles de sable et curieuses stalactites de sable (A. Jahn et A. Cailleux). Des flaques de sable, sur les chaumes couchés d'Elymus de l'année précédente, attestent que les cordons littoraux soulevés sableux s'engraissent par apport nivéo-éolien. Poste-de-laBaleine est un lieu éminemment favorable pour toutes ces études. Au Lac Guillaume-Delisle (= Richmond Gulf), outre les países, des champs de polygones de toundra (5 à 7m de diamètre) et d'ostioles, sont parmi les effets du froid sur le terrain (D. Lagarec). A Puvirnituq (60°N), de loin en loin, des dalles d'Archéen ont été redressées en monticules ou pyramides par la congélation (M. Bournérias). Enfin, l'étude statistique de 15,000 photos aériennes montre qu'entre 45° et 60°N c'est vers 55 °N que se trouve le maximum des tourbières structurées et probablement aussi celui des tourbières en général (E. Thibodeau).

P0158 Observations sur l'exportation de produits en solution par certains cours d'eau appalachiens de la région de Sherbroooke, Québec, Canada PIERRE CLEMENT Université de Sherbrooke, Canada Des observations sur les concentrations en produits dissous ont été effectuées de 1969 à 1972 sur divers cours d'eau des Appalaches dans la région de Sherbrooke, Québec: la rivière Eaton et ses affluents, soit un bassin de 650km2 environ, est à l'étude pour la Décennie Hydrologique Internationale, ainsi que les écoulements d'un petit bassin versant de 80ha, équipé pour l'étude expérimentale de l'érosion. Ces cours d'eau coulent en général sur des terrains sédimentaires (schistes ardoisiers, quartzites et calcaires impurs) ou faiblement métamorphisés, recouverts par les dépôts de la dernière phase de la glaciation wisconsinienne, dérivés du substratum. La couverture végétale est variée, allant des forêts secondaires à feuilles caduques et mixtes aux champs cultivés. Par comparaison le long des profils des cours d'eau et selon l'évolution dans le temps, l'origine de certains éléments dissous est discutée: solum, formation superficielle, roche en place, apports artificiels. Ainsi, le

fer total et l'aluminium, entraînés surtout lors des crues, sont arrachés aux berges et à la roche en place; les sulfates sont abondants dans les eaux circulant sur la roche riche en pyrite; le calcium et le magnésium obéissent à la dilution en hautes eaux et proviennent du sous-sol, la décarbonatation ayant atteint une profondeur de 1 à 2m, même dans la moraine de fond dérivée de calcaires impurs; les phosphates sont entraînés depuis les champs cultivés; les concentrations en nitrates sont commandées par le rythme saisonnier de la végétation dont le prélèvement fait tomber la courbe d'évolution dans le temps, au printemps. Le cheminement de ces éléments vers les talwegs se fait surtout par écoulement hypodermique et au contact de la roche en place, et peu par ruissellement de surface, ainsi que le montrent les résultats obtenus en parcelles expérimentales. Le rapport entre le tonnage évacué en surface et celui observé dans le cours d'eau varie de 2 à 4 en faveur du second. Une estimation de '

106 / Geomorphology l'érosion chimique pour les trois années considérées montre la part majoritaire de cette action dans l'évolution géomorphologique; l'essentiel du tonnage est exporté lors des crues, notamment celles de printemps, non seulement par suite du volume d'eau écoulé, mais aussi par suite de la préparation du matériel par les actions chimiques et biochimiques durant les périodes de basses eaux. Les nombreuses courbes d'évolution dans le temps montrent en effet des décalages par rapport aux débits. Note : Cette étude se proposant d'inclure la crue de printemps de 1972, il nous est impossible de donner ici les valeurs chiffrées qui l'accompagnent. Comité consultatif national de la Recherche géographique; Conseil national de Recherche du Canada; Ministère de l'Education du

Québec; Ministère des Richesses naturelles du Québec; l'Université de Sherbrooke. Carson, M.A., and E.A. Sutton, 1971 The hydrological response of the Eaton River basin, Quebec, Can. J. Earth Sci. 8,10215. Gadbois, P., 1970 Contribution à l'étude de l'érosion d'un petit bassin versant (environs de Sherbrooke, Québec, Canada), Mémoire de maîtrise, Dépt. géographie, U. Sherbrooke. Liekens, G.E., F. H. Bormann, N.M. Johnson, D.W. Fisher, and R.S. Pierce, 1970 Effects of forest cutting and herbicide treatment on nutrient budgets in the Hubbard Brook watershed-ecostystem, Ecal. Monograp/z,y,40(i),23-47. Ministère des Richesses naturelles du Québec, Publications DHQ 2-3-5 et HG 2.

P0159 Une forme originale de dissection des dépressions fermées en milieu aride : les kaluts du désert du Lut (Iran) ROGER COQUE Université de Paris /, France Localisé dans le sud-est de l'Iran, le désert du Lut constitue un vaste bassin endoréique de 200 sur 300km et de 205 à 500m d'altitude, bordé par de hautes chaînes montagneuses dépassant souvent 3000m. Malgré une pluviosité annuelle de l'ordre de 50mm, il apparaît totalement dépourvu de végétation sur une bonne partie de son extension (200 X 150km). Il doit son originalité géomorphologique à l'existence de formes de dissection remarquables appelées localement kaluts. C'est cet aspect du désert iranien qui fera l'objet de cette étude. Au point de vue géologique, le Lut représente un bassin molassique intra-alpin. Une puissante sédimentation synorogénique s'y est effectuée principalement au Tertiaire. Elle comporte deux séries détritiques continentales de plusieurs milliers de mètres d'épaisseur, séparées par une discordance angulaire. Ces deux séries sont bien visibles dans les coupes du piémont occidental. On y observe: une série oligo-miocène (3500m), à bancs de gypse, argiles rouges et vertes, grès rouges et conglomérats, poudingues à ciment gréseux rouge très dur; une série mio-pliocène (3000m), à poudingues mal consolidés, sa-

bles, grès et argilites sableuses, gypse et anhydrite en gros bancs. Au sommet subsistent, localement, des vestiges d'un conglomérat légèrement consolidé (kechit). Les deux séries sont énergiquement plissées en anticlinaux déversés vers l'est, accompagnés de failles longitudinales récentes. Des coulées signalent l'existence d'un volcanisme quaternaire. Dans la partie centrale de la cuvette n'affleure que la seconde série, avec des faciès plus fins et nettement évaporitiques. Elle y présente des sables, des limons et des argilites salées, des lentilles de gypse et d'anhydrite. Dans sa partie supérieure on y observe aussi des lentilles de graviers et de conglomérats. Trois grandes unités géomorphologiques se partagent la cuvette du Lut. Au nord et à l'est il s'agit d'un plateau caillouteux (reg) couronné par un encroûtement calcarogypseux, qui se termine par un escarpement à corniche bien marqué au-dessus de la dépression centrale. Un massif dunaire (erg) de quelque 9,000 à 10,000km2, dénommé Rig-e-Lut, le recouvre au sud-est. Enfin une dépression centrale allongée du NNO au SSE

Géomorphologie I 107 se caractérise par les formes de dissection appelées kaluts. Les kaluts se localisent dans une zone d'environ 150km de long sur une cinquantaine de km de large. Ils consistent en crêtes séparées par des couloirs systématiquement orientés NNO-SSE. Leur importance relative varie selon les secteurs. A l'ouest, des crêtes massives de 1 à 3km de large séparent des couloirs étroits de 0.2 à 0.3km de large. A cette partie compacte succèdent, vers le nord et vers l'est, des secteurs aérés par suite du développement des couloirs aux dépens des crêtes. Ces dernières finissent même par se réduire à des échines effilées de quelques centaines de mètres de largeur seulement, disposées en quinconce. Puis il ne subsiste plus que des cloisons, des tours et des piliers, de quelques mètres de haut, disséminés dans la plaine constituée par la coalescence des couloirs. Dans les secteurs où ils atteignent leur développement maximum, les kaluts s'allongent sur plusieurs kilomètres et leur dessin en plan dessine de larges sinuosités. Ils dominent alors les couloirs de plusieurs dizaines de mètres, jusqu'à 70-80 mètres. Avec la réduction de leurs dimensions, la forme générale tend à devenir aérodynamique comme dans le cas des yardangs, le profil longitudinal opposant une proue au vent à un effilement vers l'aval-vent. Sur les versants raides (40° ) et convexes on relève les marques nombreuses et diverses de l'érosion actuelle. Les plus importantes sont représentées par des ravins régulièrement espacés de 10 à 15m. Dans les intervalles s'inscrivent les innombrables rigoles du rill, de 5 à 6cm de profondeur, dans une pellicule solifluée d'une vingtaine de centimètres d'épaisseur de la formation des kaluts. A intervalles irréguliers se développent de grandes niches d'arrachement au-dessus de puits localisés à la base des versants, qui servent occasionnellement d'exutoires aux eaux circulant dans la masse des kaluts. Les surfaces sommitales des kaluts les plus massifs présentent, en fait, des séries de puits d'absorption alignés dans les thalwegs qui les dissèquent en lanières. Lorsque le matériel présente un faciès plus nettement lagunaire, la cohésion assurée par l'abondance des sels s'exprime par le dégagement d'entablements et de replats à corniches dus à l'érosion différentielle. Dans certains cas, des profils à

brisures multiples donnent un aspect ruiniforme typique aux crêtes. C'est à cette circonstance qu'elles doivent leur nom de kalut, qui signifie ruine en persan. Ces formes étranges posent de difficiles problèmes génétiques. Ils se situent dans deux perspectives de temps. Les uns concernent la dynamique actuelle; les autres se rapportent à l'origine même des kaluts. De toute évidence, les premiers sont les moins délicats. Leurs solutions reposent sur l'interprétation des particularités des modelés de détail des versants. Selon les cas, ces particularités révèlent des actions hydriques ou éoliennes. L'action de l'eau s'effectue selon des modalités diverses. Il y a d'abord celle du ruissellement, soit concentré dans les ravins, soit diffus dans les intervalles. Les premiers exercent une dissection active, qui peut aboutir au tronçonnement des crêtes; les seconds décapent les versants de la pellicule solifluée, après imbibation, lors des pluies les plus importantes. Mais une partie de l'eau s'infiltre à la faveur des chapelets de puits et d'une multitude de conduits plus modestes. Cette circulation interne, de type karstique, donne lieu aux grosses résurgences qui jalonnent le contact entre les crêtes et les couloirs. Leurs débits intermittents se manifestent par des creusements notables développés à partir des exutoires. Ainsi minés de l'intérieur par l'élimination des sels et des éléments détritiques les plus fins, les crêtes et leurs versants connaissent des tassements, des éboulements et des arrachements localisés. Finalement concentrées dans les couloirs, ces eaux de ruissellement et d'infiltration peuvent se mêler à celles des écoulements provenant des reliefs périphériques. Elles finissent par y disparaître, par infiltration et par evaporation, en abandonnant leurs sels sous la forme d'efflorescences et de croûtes débitées par des réseaux de fentes de dessication. Comparativement, la participation du vent au façonnement actuel des kaluts apparaît singulièrement plus modeste. Armé de sable, il exerce une corrasion dont les traces visibles à la base des versants contribuent à raidir leurs profils. On peut aussi constater les effets de la déflation dans les couloirs, notamment sous l'aspect de cuvettes hydro-éoliennes de forme ovale, creusées à la faveur de la floculation des particules argileuses et limoneuses, par les cristallisations salines, dans les creux

108 / Geomorphology où se concentrent les eaux. Ailleurs le vent pousse vers le Sud les nappes d'eaux salées qui les inondent épisodiquement. Leur evaporation laisse des franges salines aux festons caractéristiques. Enfin le vent entraîne les trains de barkhanes qui finissent par submerger les couloirs dans leurs sections méridionales. Mais l'observation montre également que les parts respectives de l'eau et du vent, dans le façonnement actuel des kaluts, varient en fonction de leurs dimensions. Sur les plus importants d'entre eux les actions hydriques prédominent nettement, favorisées par les impluviums qu'ils constituent. Les attaques éoliennes l'emportent, au contraire, sur les plus modestes, d'où leur aérodynamisme, puis sur les parois verticales des buttes résiduelles marginales. Les vrais yardangs que l'on rencontre dans certaines marges, au nord notamment, représentent, peut-être, un stade très avancé d'une dissection qui aboutit à un approfondissement de la cuvette. Bien des arguments nous invitent à envisager de telles modifications dans le rôle de l'eau et du vent au cours du Quaternaire, à la faveur des variations climatiques classiques que connaissent alors les déserts tropicaux. Ici, également, des étagements de glacis d'érosion et leurs couvertures de débris en témoignent dans les piémonts. De toute évidence, les Pluviaux intensifient les actions de l'eau aux dépens de celles du vent, des situations inverses se réalisant au cours des Interpluviaux. Cependant, en dépit des apparences, le vent constitue sans nul doute l'agent le plus efficace de la morphogenèse. La dissec-

tion en lanières de la formation mio-pliocène suit une orientation systématique NNO-SSE qui est la sienne. Mais cette direction est aussi celle des accidents tectoniques plioquaternaires du piémont occidental. On peut penser (G. et J. Conrad) que l'ablation éolienne a pu exploiter des diaclases majeures parallèles aux accidents bordiers, phénomène classique en milieu désertique accusé, même dans des roches plus cohérentes comme les grès (M. Main guet-Michel). Enfin, seul le vent reste susceptible d'exporter du matériel d'une dépression fermée. Le creusement de la cuvette centrale du Lut, expression du bilan de son évolution morphologique au cours du Quaternaire, établit, en définitive, la primauté de la participation éolienne. Conrad, G., and J. Conrad, 19700 Le Tertiaire continental des monts de Kerman et du Lut (Iran oriental), C.R.Acad. Sci., Paris 270, 1421-3. - 19706 L'évolution quaternaire de la dépression du Lut (Iran oriental), ibid., 1672-4. Dresch, J., 1968 Reconnaissance dans le Lut (Iran), Bull. Assoc. Geog. fr., 362-3, 143-53. Gabriel, A., 1938 The southern Lut and Iranian Balutchistan, Geog. J. 102, 193210. - 1964 Zum Problem der Formenschatzes im extrem-ariden Raiimen, Mittel. Osterr. Geog. Ges., Wien 106, 3-15. Mainguet-Michel, M., 1968 Le Borkou, aspects d'un modelé éolien, Ann. Géog., Paris 421, 296-322.

P0160 Le système morphogénétique 'géodynamique' JEAN DEMANGEOT Université de Paris X, France i. On sait que la glyptogénèse résulte non d'une 'érosion', au sens vague du terme, mais de l'action de 'systèmes morphogénétiques' bien définis: système périglaciaire, système semi-aride, etc. Tous ces systèmes ont en commun de résulter d'actions exogènes bioclimatiques: ruissellement de l'eau, gélivation, etc. Or il existe un système provoqué par des actions endogènes, que nous avons déjà proposé d'appeler système géodynamique (Demangeot 1968), et qu'il conviendrait peut-être d'ajouter à la liste des systèmes

classiques. Comme eux, il procède d'abord à une microfragmentation des versants, comparable par bien des aspects au 'weathering', puis à un déplacement des particules, comparable à l'érosion stricto sensu. il. Nous appellerons 'géodynamique' tout mouvement récent ou actuel d'origine endogène et susceptible de modifier la topographie: déformations néotectoniques classiques, volcanisme superficiel ou laccolithique, diapirisme salifère. La plupart de ces mouvements sont accompagnés de tremblements de

Géomorphologie I 109 TABLEAU 1. Les enchaînements du système 'géodynamique' Conséquences statiques A Macrostructuration

Microstructuration

Conséquences dynamiques B

Effets morphologiques2

Néotectonique

Abaissements et soulèvements du sol; gradins de faille

Indurations, broyages, mylonitisations, laminages

Secousses, effet 'table vibrante'

Volcanisme

Bombements laccolithiques ; coulées; failles

Cuisson, fendillement

Secousses

Primaires : tassements, écroulements, éboulis, éboulements, coulées, glissements, etc.

Tectonique salifère Séismes profonds

Bombements du sol; petites failles ?

Broyages, fendillements Trituration

9

Détente post- tectonique1

?

Roche 'expansée'

Causes géodynamiques

Secousses et effet 'table vibrante' Projections directes, secousses ?

Secondaires : désorganisation hydrologique et hydrographique, réadaptations, épandages, etc.

1 Et détente post-érosion (vallées déglacées, cavernes, pains de sucre ...). 2 Par combinaison avec les systèmes bio-climatiques. terre à foyer superficiel. Il faut y ajouter les effets des séismes profonds, explicables ou non par le 'mécanisme au foyer' (Symposium 1957) : dans ces derniers cas les séismes ne sont plus conséquences mais causes. m. Ces mouvements endogènes variés, moins indépendants les uns des autres qu'ils n'en ont l'air, aboutissent à des résultats que l'on peut subdiviser en 'résultats statiques' et 'résultats dynamiques'. Les résultats statiques A (voir tableau) consistent en une néostructuration de la surface du sol, mais à des échelles différentes. On peut appeler 'macrostructuration' l'ensemble des déformations d'échelle métrique ou décamétrique : bombements, gradins de failles. Les exemples en sont innombrables, en Italie centrale, pour ne citer qu'un cas (Demangeot 1965). On peut, ensuite, appeler 'microstructuration' la fragmentation des roches à l'échelle décimétrique, centimétrique ou millimétrique. La mylonitisation ou le broyage des plans de faille est l'exemple le plus connu de cette microstructuration géodynamique. Mais on sait, depuis quelques années seulement, que la détente mécanique des roches après leur cisaillement provoque également une fragmentation : d'où la roche 'expansée' des ingénieurs. Il est à remarquer que la poussée au vide engendrée par le creusement rapide des versants, voire des cavernes souterraines, entraîne également une microstructuration de détente (Galibert

1965; Renault 1970). Les effets de la microstructuration géodynamique sont extrêmement voisins de ceux que provoquent les agents classiques du weathering, et il n'est pas toujours facile de distinguer une brèche de détente d'une brèche de gélifraction par exemple. Les mouvements endogènes ont également des résultais dynamiques (B) puisqu'ils communiquent aux fragments rocheux une énergie cinétique propre, qui ne doit rien à l'entraînement par les agents exogènes. On distinguera: a) les secousses classiques (quelle qu'en soit l'origine) : les débris rocheux, déjà placés en porte-à-faux, se mettent à glisser ou à tomber, d'où des chutes de pierres, des éboulements accompagnés de nuages de poussière, des glissements de terrain même. Il est probable que les tremblements de terre aident au façonnement des cavités karstiques. b) les déplacements horizontaux de paquets de terrain entiers, par effet 'table vibrante', sur plusieurs centaines de mètres ou plusieurs kilomètres. Ce singulier mécanisme a été pressenti par Montessus de Ballore à propos des séismes d'Assam (Montessus de Ballore 1924, 44) et décrit au Chili plus récemment (Tazierî 1960). c) les projections de fragments expansés d'une roche en état de détente : tel miroir de faille frais, telle paroi de grotte, explosent en surface et projettent leurs éclats latéralement,

110 / Geomorphology sans intervention exogène. Le phénomène est connu des ingénieurs et il a causé beaucoup d'accidents lors du percement du tunnel sous le Mont-Blanc par exemple (Janin 1962). iv. Cette production et, parfois, cette projection de débris rocheux sont évidemment liées à la lithologie (préfiguration des fragments). Mais si elles sont, dans leur essence, indépendantes des conditions exogènes, elles se combinent avec elles dans la réalité géographique, puisqu'elles se produisent en un point donné du Globe. Il y a donc nécessairement interférence d'une part avec les systèmes bio-climatiques, d'autre part avec les conditions topographiques (accélération des processus en montagne). D'où des résultats morphologiques très variés. La Cordillère des Andes nous offre un large échantillonnage de ces effets: colmatage des dépressions intraandines tropicales (Tricart et Michel 1965), grands versants réglés et talus d'éboulis sous climat désertique (Dollfus, Gabert et Laharie 1970), topographie 'cataclysmique' avec loupes de glissement et coulées en climat tempéré humide (Borde 1966), etc. Le domaine d'extension du 'système géodynamique' est immense. Il englobe non seulement toutes les régions de tectonique récente (arcs alpin et circum-pacifique), mais aussi les régions anciennes et stables où l'érosion a creusé de grandes incisions génératrices de détente mécanique (plateaux tropicaux, chaînes Scandinaves, etc.). Un immense

champ de recherches s'ouvre aux géomorphologues, si ces prémisses sont exactes. Borde, J., 1966 Les Andes de Santiago et leur avant-pays (Bordeaux). Demangeot, J., 1965 Géomorphologie des Abruzzes adriatiques (Paris). - 1968 Mouvements du sol et morphogenèse (Lisbonne), 182-208. Dollfus, O., P. Gabert, et R. Laharie, 1970 Les problèmes morphologiques du piémont désertique des Andes péruviennes, Rev. Géog. alpine 265-300. Galibert, G., 1965 La haute montagne alpine (Toulouse). Janin, B., 1962 Les tunnels routiers du Mont-Blanc etc, Rev. Géog. alpine 87-120. Montessus de Ballore, B. de, 1924 La Geologie seismologique (Paris). Renault, Ph., 1970 La formation des cavernes (Paris). Symposium, 1957 The Mechanics of Faulting (Ottawa, Publ. Dominion Observatory, 20). Tazieff, H., 1960 Interprétation des glissements de terrain accompagnant le grand séisme du Chili, Bull. Soc. belge Géol. iv, 1-11. Tricart, J., et M. Michel, 1965 Monographie et carte géomorphologique de la région de Lagunillas (Andes vénézuéliennes), Rev. Géomorphol. dynamique, xv, 1-14.

P0161 A propos des limites inférieures des phénomènes cryonivaux et glaciaires quaternaires aux marges méridionales et orientales du domaine méditerranéen JEAN DRESCH Université de Paris VU, France Des progrès sont faits dans la recherche des changements de climat pendant le Quaternaire dans le domaine méditerranéen. Et pourtant, encore, des opinions très diverses sont avancées. Certains prétendent que le gel fut intense jusqu'au niveau de la mer au point que le permafrost explique certaines figures sur les rivages du Liban et d'Israël (De Vaumas 1963, etc.). Ils s'appuient sur les estimations de température des eaux de la Méditerranée par Emiliani. Mais ces estimations sont très contestées et beaucoup de formes décrites ne peuvent être expliquées par une cryergie

intense au niveau de la mer. Il a pu geler et neiger aux bords méridionaux et orientaux de la Méditerranée, puisqu'aujourd'hui le thermomètre peut, sous abri, descendre à 0° et que des flocons de neige ne sont pas très rares en hiver. De la sorte, sur des versants de calcaires gélifs, des dépôts de pente lités ont pu se former localement jusqu'à proximité du niveau marin. On ne saurait en conclure que le sol pût être gelé durablement. Il n'en existe aucune preuve sûre, même en haute plaine, ou haute montagne. S'il est très généralement admis que les températures au niveau de la mer furent

Géomorphologie I 111 sensiblement supérieures sur les rives méridionales et orientales de la Méditerranée à celles des rives européennes, les observations conduisent à des conclusions divergentes au Maghreb, au Levant et au Moyen-Orient continental. Au Maghreb, les observations faites au Maroc ont permis d'établir et d'étendre au reste de la région une chronologie devenue, dans sa précision croissante, quelque peu schématique. Au cours de pluviaux plus ou moins froids, des phénomènes périglaciaires, nivaux et glaciaires se sont manifestés à des altitudes inférieures aux altitudes actuelles. L'action du gel et de la neige a laissé des traces au dessus de 800m dans le Rif marocain et le Tell algéro-tunisien et jusque dans le nord de la Dorsale tunisienne, et a exercé une influence importante dans les formes héritées au dessus de 1300m. Des cirques s'observent en contrebas des plus hautes crêtes (2500 dans le Rif, 2200 en Grande Kabylie). Ces limites approximatives concernent les versants où l'orientation est favorable. Elles s'élèvent vers le sud. Au Maroc, dans le Moyen Atlas, la limite inférieure des phénomènes périglaciaires généralisés, pendant le dernier pluvial, peut être suivie vers 2200m, et la limite inférieure des neiges persistantes à environ 3000m (massifs du Bou Iblane et du Guèberaal). Dans le Haut Atlas, des cirques glaciaires s'observent dans les massifs dont les crêtes dépassent 3700m. Le fond de la plupart est à 3500m. Dans le seul massif culminant du Toubkal, des appareils glaciaires ont sculpté des vallées sans dépasser 5km de long et des glaciers rocheux ont glissé jusqu'à 1900m. Mais Pinfluence morphologique de la neige fut importante jusqu'à 1700m. Ces limites sont celles de la dernière glaciation. Une glaciation antérieure a laissé des traces rarement observables. Elle dut être plus importante. Vers l'est, en régime plus continental, l'action du gel a pu s'exercer plus bas, à moins de 1400m. Au Hoggar, au cœur du Sahara, des moraines de névé ont été signalées au dessus de 2200m. Dans les montagnes du Levant méditerranéen, si dans le bas pays les fouilles archéologiques ont permis d'apporter des précisions souvent remarquables sur la paléoclimatologie d'un Quaternaire qui, comme en Afrique du Nord, ne fut jamais ni très froid, ni très aride, les opinions divergent davantage sur les changements climatiques survenus en

montagne. En Grèce centrale, la limite se serait élevée de 1700 à 2200m d'ouest en est (Messerli 1967). On la retrouverait entre 2100 et 2300m selon diverses observations en Anatolie occidentale, beaucoup plus haut, à 2900, en Anatolie centrale. Elle était supérieure à 2400 en Crète. Au Liban, au flanc du plateau sommital de Qornet es Saouda, sous le vent, côté Békaa, des cirques portent encore des taches de neige plus ou moins permanente et sont fermés, en aval, par des moraines de névé vers 2800-2900m. On peut faire des observations comparables sur le versant sud-oriental de l'Hermon (2814) où des moraines de cirque sont situées à un peu plus de 2500m. Mais dans ces deux cas les conditions d'accumulation et de conservation de la neige sont et étaient exceptionnellement favorables. Aussi bien les moraines signalées par K. Kaiser (1961) dans le cirque des Cèdres, au dessus de Faraya, et dans la vallée du Ouadi Berdaouni, au dessus de Zahlé, ne sont que des coulées de solifluction liées à l'importance des marnes, dans la série de l'Albien, de l'Aptien et plus encore du Cénomanien, et des failles dans la vallée du Berdaouni où ces formations tendres sont en outre écrasées. Les coulées de solifluction, si fréquentes à toutes altitudes dans la région, ont été favorisées par un climat beaucoup plus humide, neigeux, que dans les bassins intérieurs, toujours beaucoup plus secs. Les hauts plateaux du Sannine et plus encore du Qornet et Saouda furent à coup sûr recouverts par des névés ou même des glaciers de plateaux au dessus de 2500-2600m, soit environ 7 00-8 00m au dessous de la limite probable actuelle des neiges persistantes. S'il n'y eut pas de vraies langues glaciaires, un vaste glacier rocheux, au pied sud du Qornet es Saouda, témoigne à la fois de l'importance de la couverture nivo-glaciaire dans les vallées karstifiées de plateau, et de l'importance de la cryoclastie dans la série du Cénomanien très gélive. Des actions cryonivales, figures de sols 'polygonaux' ou striés, dépôts de pente lités, s'observent jusqu'à des altitudes de 1800m (Klaer 1957). Elles ont fourni le matériel qui explique les terrasses remarquables de la Bekaa. Ces contrastes entre les climats des montagnes et des plaines au Quaternaire se trouvent en Iran plus évidemment encore qu'aujourd'hui. L'Elbourz, le massif du Tacht Soleiman du moins, possède encore des

112 / Geomorphology glaciers. L'action nivo-glaciaire s'est manifestée jusque dans les massifs culminant à environ 3000m, altitude du reste parfois contestée, et des actions cryo-nivales quaternaires sont évidentes 700-800m plus bas. Mais le piémont de Téhéran fut toujours beaucoup plus sec. Les limites actuelles ou passées sont encore mal connues dans le Zagros. Les montagnes de la région de Kerman qui dépassent 3500 ou même 4000m ne semblent pas porter la marque d'anciens glaciers, sauf peut-être le Kuh-e-Taftan (4042), au sud de Zahedan, vraisemblablement plus arrosé. Mais des phénomènes cryo-nivaux y sont évidents et il est certain qu'ils furent plus importants au Quaternaire, parce que le climat fut non seulement plus froid, mais aussi plus humide, contrairement à l'hypothèse de H. Bobek (1963). On ne saurait expliquer autrement les dépôts de versants qui se ravinent, ainsi que les glacis d'ablation en roche tendre, les kewir ou dacht, et les terrasses, étages dans les piémonts, quand des abaissements des niveaux de base l'ont per» mis. Ils sont généralement au nombre de trois, parfois davantage. Ils témoignent de l'importance des précipitations, vraisemblablement neigeuses surtout, comme aujourd'hui, en haute montagne, et d'une cryoclastie qui fournissait la charge des torrents. La cuvette du Lut a toujours été, au Quaternaire, très aride au point d'être restée azoïque. Mais les eaux apportées par lès hautes montagnes de l'ouest, la chaîne de Kerman, y ont modelé, avec le vent, des formes remarquables (Dresch 1968). Tel dut être le cas dans tout le domaine méditerranéen semi-aride et aride. Les mon-

tagnes y furent, à plusieurs reprises, au cours du Quaternaire, plus fraîches et plus humides. Dans les bas pays plus secs, les changements de climat furent beaucoup plus atténués. Mais les conditions régionales, latitude, continentalité, importance relative des massifs montagneux plus ou moins élevés, volumineux, bien exposés aux vents humides, capables d'accumuler de la neige, ampleur des versants, volume des dépressions ou plateaux, organisation du drainage endoréique ou exoréique, expliquent la complexité des relations entre hauts et bas pays et les difficultés de préciser les limites en altitude des phénomènes déterminés par les changements de climat. Bobek, H., 1963 Nature and implications of Quaternary changes in Iran, Changes of climate, Colloque UNESCO, Rome. De Vaumas, 1963, 1964, 1970 Nombreuses publications, Rev. Géog. phys. et Géol. dyn. Dresch, J., 1968 Reconnaissance dans le Lut, Iran, Bull Assoc. Géog. fr. 362-3. Kaiser, K., 1961 Die Ausdehnung der Vergletscherungèn, 6e Congrès int. du Quaternaire, Varsovie, t. m. Klaer, W., 1957 Beobachtungen zur rezenten Schure — und Strukturbodengrenze im Hochlibanon, Z. Geom. 1. Messerli, B., 1967 Die eiszeitliche und die gegenwârtigen Vergletscherung im Mittelmerraum, Géog. Helvetica 3. Péchoux, P.Y., 1970 Traces d'activité glaciaire en Grèce centrale, Rev. Géog. alpine.

P0162 Etude de la denudation chimique des pays montagneux H.K. GABRiELiAN Université d'Etat de Yerevan, URSS Parmi les processus fluviaux dans le développement exogène du relief, l'écoulement des matières dissoutes - l'érosion chimique de la terre ferme - occupe une place à part. Le transport des matières dissoutes par les fleuves est le meilleur critère pour l'expression de l'intensité de cette érosion chimique. Mais l'aspect géomorphologique en est encore très mal connu. Les premières tentatives visant à l'évaluation de l'écoulement fluviatile

chimique ont été faites par I. Murray (1887) et A. Penck (1894). L'évaluation globale la plus connue, établie à partir des sources les plus certaines, est l'œuvre de F. Clarke (1924); elle porte principalement sur les données concernant l'Europe et les EtatsUnis 24mph) az(x,L) = vertical standard deviation of the pollutants as a function of distance (x in m) from the source and the atmospheric stability class (L). Five stability classes are defined according to Turner (1964, 90-91 ) ; H = effective emission height (in m) ; T1/2 — time in which one-half of a pollutant has been removed (for SO2 about 4 hours) Equation (1 ) is applicable when the vertical distribution of the pollutants is not restricted. Under inversion situations, however, with a shallow mixing layer (M in m) the groundlevel concentration is calculated from (2)

X(*, L,

S, M) = [Q/Mu(S)(2n/l6)] x exp[-(0.639*/«CS))/r1/2]. The long-term average concentration xav °f a receptor is then calculated from either equation (1 ) or (2) by summing the individual concentrations for all sources (N), stability (L), and wind speed classes (5) : (3) %av= ZNXLZsF(Dn,L,S)"x,(Xn,L9S) where Dn = the wind direction sector in which transport from a particular source n to the receptor occurs Xn = the distance from a particular source n to the receptor F(Dn,L,S) — relative frequency of winds blowing into the given 22.5° wind direction sector (Dn), for a given stability class (L), and wind speed class (5 ). Martin's model has, however, one significant drawback in that it assumes a constant wind speed and wind frequency within the 22.5° sector. Daniels (1971) has shown that this produces errors in the concentrations of up to ±40%. These serious errors in diffusion computations can be avoided by applying harmonic analysis which relates the pollutant calculations to angular intervals of one degree. The Gifford-Hanna model is a relatively

130 / Climatology simple model for estimating long-term average pollutant concentrations from urban sources which has been used by the authors as a baseline for comparison with estimates from more sophisticated diffusion models (Gifford and Hanna 1970; and Hanna 1971 ). They contend that for many operational uses such as legislative and prompt enforcement action the computation time for the model should be short, and for engineering design the computations should be essentially simple. Only in special research which aims to unravel the complexities of urban air pollution would the development and use of the most sophisticated models be justified. For each grid square the source strength (Q in g /sec) is calculated from emission source inventories, which are usually presented on a checkerboard pattern with a grid spacing (Ax) of one to ten kilometers. Assuming a constant wind direction with speed (u in m/sec), the ground-level concentration (x in ¿ig/m3) in grid number '0' is calculated from

SO2 concentrations as predicted by Miller and Holzworth (1967, 49-50) for Nashville (AM and PM) and Los Angeles (PM) were 0.84, 0.84, and 0.88, respectively. Using equation (5) and including the 95 per cent confidence limits, Hanna's calculations were 0.73 (0.50-0.86), 0.73 (0.50-0.86), and 0.75 (0.58-0.86). Since the 95 per cent confidence limits overlap, it can be concluded that the more sophisticated Miller/Holzworth model does not yield significantly (a = 0.05) better results than the simple equation in (5). For the estimation of annual average pollutant concentrations for urban areas the Environmental Protection Agency (Federal Register, 1971, 6686) suggests using a model developed by Holzworth (1970). The model input requires the average emission rate for the city area (2av in g sec-1 m-2), the size of the city, i.e. the downwind distance across the city (S in m), and the wind speed (u in m sec-1) through the mixing layer (M in m). Assuming ground-level and uniform emissions and neglecting the reactivity and the 1 lateral diffusion of pollutants, the average (4) x = (2/7t)*l/«[(A*/2)/