Slope Tectonics (Geological Society Special Publication 351)

Slope Tectonics

The Geological Society of London Books Editorial Committee Chief Editor

Bob Pankhurst (UK) Society Books Editors

John Gregory (UK) Jim Griffiths (UK) John Howe (UK) Rick Law (USA) Phil Leat (UK) Nick Robins (UK) Randell Stephenson (UK) Society Books Advisors

Mike Brown (USA) Eric Buffetaut (France) Jonathan Craig (Italy) Reto Giere´ (Germany) Tom McCann (Germany) Doug Stead (Canada) Gonzalo Veiga (Argentina) Maarten de Wit (South Africa)

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk.

It is recommended that reference to all or part of this book should be made in one of the following ways: Jaboyedoff, M. (ed.) 2011. Slope Tectonics. Geological Society, London, Special Publications, 351. Jarman, D., Agliardi, F. & Crosta, G. B. 2011. Megafans and outsize fans from catastrophic slope failures in Alpine glacial troughs: the Malser Haide and the Val Venosta cluster, Italy. In: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 253–277.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 351

Slope Tectonics

EDITED BY

M. JABOYEDOFF University of Lausanne, Switzerland

2011 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 10 000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ 44 (0)20 7434 9944; Fax þ 44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ 44 (0)1225 445046, Fax þ 44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2011. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/11/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-324-0 Typeset by Techset Composition Ltd, Salisbury, UK Printed by MPG Books Ltd, Bodmin, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ 1 800-972-9892 Fax þ 1 802-864-7626 E-mail: [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ 1 918-584-2555 Fax þ 1 918-560-2652 E-mail: [email protected] Website: http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ 91 11 2327-9113/2326-4180 Fax þ 91 11 2326-0538 E-mail: [email protected]

Contents Foreword JABOYEDOFF, M., CROSTA, G. B. & STEAD, D. Slope tectonics: a short introduction

vii 1

EL BEDOUI, S., BOIS, T., JOMARD, H., SANCHEZ, G., LEBOURG, T., TRICS, E., GUGLIELMI, Y., BOUISSOU, S., CHEMENDA, A., ROLLAND, Y., CORSINI, M. & PE´REZ, J. L. Paraglacial gravitational deformations in the SW Alps: a review of field investigations, 10Be cosmogenic dating and physical modelling

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SAINTOT, A., HENDERSON, I. H. C. & DERRON, M.-H. Inheritance of ductile and brittle structures in the development of large rock slope instabilities: examples from western Norway

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HENDERSON, I. H. C. & SAINTOT, A. Regional spatial variations in rockslide distribution from structural geology ranking: an example from Storfjorden, western Norway

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BO¨HME, M., SAINTOT, A., HENDERSON, I. H. C., HENRIKSEN, H. & HERMANNS, R. L. Rock slope instabilities in Sogn and Fjordane County, Norway: a detailed structural and geomorphological analysis

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MARTINOTTI, G., GIORDAN, D., GIARDINO, M. & RATTO, S. Controlling factors for deep-seated gravitational slope deformation (DSGSD) in the Aosta Valley (NW Alps, Italy)

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BARON˘, I., KERNSTOCKOVA´, M., NOVOTNY´, R., BURIA´NEK, D., HRADECKY´, P., HAVLIC˘EK, P. & MELICHAR, R. Palaeostress analysis of a giant Holocene rockslide near Boaco and Santa Lucia (Nicaragua, Central America)

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JABOYEDOFF, M., OPPIKOFER, T., DERRON, M.-H., BLIKRA, L. H., BO¨HME, M. & SAINTOT, A. Complex ˚ knes rockslide, landslide behaviour and structural control: a three-dimensional conceptual model of A Norway

147

PEDRAZZINI, A., JABOYEDOFF, M., FROESE, C. R., LANGENBERG, C. W. & MORENO, F. Structural analysis of Turtle Mountain: origin and influence of fractures in the development of rock slope failures

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HENDERSON, I. H. C., LAUKNES, T. R., OSMUNDSEN, P. T., DEHLS, J., LARSEN, Y. & REDFIELD, T. F. A structural, geomorphological and InSAR study of an active rock slope failure development

185

BIANCHI FASANI, G., DI LUZIO, E., ESPOSITO, C., MARTINO, S. & SCARASCIA-MUGNOZZA, G. Numerical modelling of Plio-Quaternary slope evolution based on geological constraints: a case study from the Caramanico Valley (Central Apennines, Italy)

201

AMBROSI, C. & CROSTA, G. B. Valley shape influence on deformation mechanisms of rock slopes

215

GHIROTTI, M., MARTIN, S. & GENEVOIS, R. The Celentino deep-seated gravitational slope deformation (DSGSD): structural and geomechanical analyses (Peio Valley, NE Italy)

235

JARMAN, D., AGLIARDI, F. & CROSTA, G. B. Megafans and outsize fans from catastrophic slope failures in Alpine glacial troughs: the Malser Haide and the Val Venosta cluster, Italy

253

Index

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Foreword Landslides represent serious problems for humans and infrastructure throughout the world. In order to cope with these problems, knowledge about the fundamental processes leading to landslides is of major importance. This Special Publication Slope Tectonics introduces a new term dealing with structures and processes formed mainly as a result of gravitation. Tectonics is a term that is normally linked to major structural or deformational features, and their relation, origin and evolution. Research on tectonic processes has, however, not been of main focus for people working on landslide analysis, although it can be also perfectly transformed to structures formed by gravitation. For example, structural geology and tectonics has not been the main focus in disciplines such as engineering geology and rock mechanics. Geological features and processes formed by slope tectonics are often the most important controlling factor for slope instabilities and failures. A better understanding of slope tectonics and its features are of vital importance for constructing reliable geological models for landslides, which again is necessary for stability analysis and for design and implementation of reliable mitigation measures. Slope Tectonics stresses the need and importance of a better understanding of slope tectonics, and the controlling geological structures for hazard mapping and analysis of large landslides. It presents a variety of aspects, including regional hazard analysis, structural geology, the evolution of slope failures, rock mass characteristics and strength, site-specific geological failure models, numerical modelling and the use of new remote-sensing technologies. The 13 papers and the introductory paper are based on presentations given at a research workshop

at the University of Lausanne, and are written by scientists from active landslide research groups in Italy, Switzerland, Norway, France, Czech Republic and Canada. They present a series of excellent case studies from Norway, Italy, France, Central America and Canada. The presented cases demonstrate the need for integrating different disciplines for a better understanding of landslide processes and their hazards. Slope Tectonics is edited by Professor and Director Michel Jaboyedoff from the Institute of Geomatics and Analysis of Risk at the University of Lausanne, Switzerland. Michel Jaboyedoff has a wide and long experience in landslide research, and is particularly well known for his interdisciplinary skills, creativity, ability to use new technology and to implement scientific results into practical use. In addition, he has, over the last few years, built up an excellent landslide group at the University of Lausanne. As a scientist responsible for the monitoring and early warning of large rockslides in inhabited areas in Norway, I hope that this book, Slope Tectonics, will generate enthusiasm and increased research activity of importance for hazard and risk handling in mountainous areas and fjord regions throughout the world.

Lars Harald Blikra Professor and Chief Geologist A˚knes/Tafjord Early Warning Centre and Sogn and Fjordane University College Norway April, 2010

Slope tectonics: a short introduction MICHEL JABOYEDOFF1*, GIOVANNI B. CROSTA2 & DOUG STEAD3 1

2

Institut de Ge´omatique et d’Analyse des Risques, University of Lausanne, Amphioˆle, 1015 Lausanne, Switzerland

Dipartimento di Scienze Geologiche e Geotecnologie, Universita` degli Studi di Milano Bicocca, Milano, Italy

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Department of Earth Sciences, Simon Fraser University, Burnaby, B.C., Canada V5A 1S6 *Corresponding author (e-mail: [email protected])

Geomorphology, structural geology and engineering geology allow description of the main characteristics of a slope in distinct ways that can be combined to provide a complementary view of the operative slope processes. The subjects presented in this Special Publication include: slope morphology and evolution; mechanical behaviour of the material; modes of failure and collapse; influence of lithology and structural features; and the role played by controlling factors. This Slope Tectonics volume comprises a series of very different contributions that attempt to underline a multidisciplinary approach that should form the framework of slope instability studies. Slope Tectonics is adopted in this volume to mean deformation that is induced or fully controlled by the slope morphology and that generates features that can be compared to tectonic features. The stress field in a slope is the result of gravity, topography and the geological setting created by an ensemble of geodynamic processes. Active tectonics (also called neotectonics) generates a stress field that can control slope processes; a strong feedback existing between geological history, tectonics, lithology, geomorphological evolution and topography. As a consequence, a list of factors and their relative influence can be presented. (1)

Fabric induced by a local stress field within a slope: † discontinuities and local faults with cataclastic bands of variable thickness; † folds (Fig. 1), associated predominantly with brittle structures; † complex failure paths (stepped or multisurface); † local failures: rock bridge failures or extensional failures (graben-like or pseudograben-like); † subsidence due to weak or soluble materials causing complex sliding –toppling phenomena.

(2)

(3)

Reactivation of pre-existing faults, discontinuities, joints, foliations or rock anisotropies: † surfaces characterized by residual or lower than peak strength; † formation of composite failure surfaces. Regional tectonic movements inducing new slope morphologies: † uplift; † major fault movement; † pull-apart zones: † folding.

The boundary between classical tectonics and slope deformations, especially at a large scale, has always been indistinct as emphasized by Antoine (1988) (Schultz-Ela 2001). In his paper Antoine discussed mechanisms like ‘diverticulation’ defined by Lugeon (1943) and Badoux (1963), where part of Pre-Alpine nappes were reversed in geometry by huge landslides inverting the stratigraphy (Antoine 1988). A remarkable geometrical analogy exists between basin extensional tectonics (Wernicke 1981) and certain landslide spreading in clays (Voight 1973; Varnes 1978; Hutchinson 1988); this is despite the fact that Wernicke’s hypothesis demonstrated that Basin and Range regions were not produced by huge landslides, but by low-angle faulting induced by geodynamic processes. Regional extension has produced changes in topography and, as a consequence, significant gravity-induced deformations. Therefore, we suggest that the term ‘slope tectonics’ is justified and must be recognized as an important component in slope deformation. Slope instability implies movements driven by gravity that can produce irreversible deformations. In the past, slopes were viewed as privileged erosion zones (De la Noe & De Margerie 1888; Strahler 1950), whereas few erosion processes were attributed to landslides (Young 1972). Since the work of Simonett (1967) and Hovius et al. (1997), the link

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 1 –10. DOI: 10.1144/SP351.1 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Kink-fold in the front of a translational landside in foliated amphibolites of northern Norway. This shows a similar structure to folds in thin skin tectonics style except that the lithostatic stress is not sufficient to close the voids. (From Braathen et al. 2004; reproduced with permission from the authors and Norwegian Journal of Geology.)

between landslides and erosion processes has been emphasized by various authors (Korup et al. 2010). This linkage has been suggested both for shallow and deep-seated landslides. In the first case, shallow landslides frequently evolve into debris flows along the slopes, allowing a faster and more complete transport to the fluvial network. For deepseated landslides the processes act in a slightly different way causing a slow removal of the material from the slope. In these cases the erosion and the rate of erosion induced by landslides apply over a different timescale and very often can be considered as a means for both morphological stabilization, through a decrease in slope gradient, and the introduction of marked changes in valley morphology and river sediment yield. Static and dynamic equilibrium can be important controls on the evolution of very large slope deformations. In fact, rivers are linear elements incising the topography and which do not erode the entire slope topographic surface. The slope surface is affected by erosive processes that, in fact, affect the entire topographic surface. As a consequence, diffused erosion or mass movements are required to displace material away from slopes into the fluvial system. Small and large slope instabilities clearly act differently in degradation of slopes and mountainous topography according to their volume, density and temporal frequency, and to the degree of disturbance that they cause to the slope materials. Gravity plays an important role in large deformations of the Earth’s crust (De Jong & Scholten 1973; Schultz-Ela 2001). Molnar & Lyon-Caen

(1988, 1989) showed that during mountain belt building, above a certain thickness of the Earth’s crust, the gravitational stress may be higher than the compressive strength. As a consequence, the mountain belts cease to elevate and begin to spread laterally (Delacou et al. 2004). Theoretically, this phenomenon can be considered to be similar to cases studied by Savage et al. (1985) and Savage & Varnes (1987) (Fig. 2). This hypothesis is also supported by the results of analogue modelling (Bachmann et al. 2009). Lugeon & Oulianoff (1922) have described the role of gravity-induced deformations on rock outcrops (Fig. 3) and began the documentation of a considerable number of large slope movements (Heim 1932). Terzaghi (1963) identified deepseated landsliding from both a mechanical and kinematic point of view. Subsequently, it has been shown by analytical solutions that gravitational stresses can induce slope deformations (Savage et al. 1985; Savage & Varnes 1987) (Fig. 4). More recent models support these findings for more complex geometries and material properties (Kinakin & Stead 2005; Ambrosi & Crosta 2006). Nevertheless, new techniques available for the observation of ground movements (e.g. InSAR (Interferometric Synthetic Aperture Radar), LiDAR (Light Detection And Ranging) and highresolution satellite imagery) provide the opportunity to show that the valley flanks within major mountain belts are moving both slowly and continuously over large areas (Singhroy 1995; Carnec et al. 1996; Dehls et al. 2002; Colesanti et al. 2003; Ambrosi

INTRODUCTION

3

Fig. 2. (a) Extension and compression zones deduced from the analytical model (modified after Savage & Varnes 1987). (b) Illustration of a DSGSD that presents similar features to those predicted by the model shown in (a) (modified after Agliardi et al. 2009).

& Crosta 2006; Colesanti & Wasowski 2006; Troisi 2007; Crosta et al. 2008b; Osmundsen et al. 2009) in a form of dynamic equilibrium that persists over thousands of years. The present landslides ‘paradigm’ did not predict this because most of the studies of mass movements were dedicated to areas displaying clearly present or past slope movements that were identified based on morphology. However, unfortunately, a lot of structures were either not identified or were confused with regional tectonic features. Again, the spatial and temporal scale typical of such phenomena suggest a strong link between gravity and tectonics, as well as between local morphological changes and the tectonic evolution of mountain belts (Crosta et al. 2009; see also Ambrosi & Crosta et al. 2011).

Structures associated with slope deformation Large slope movements have been encountered during dam projects (Lugeon 1933; Gignoux & Barbier 1955; Desio 1961) creating problems for dam stability and in meeting the impermeability requirements for foundation materials (Fig. 5). Drilling and excavation of hard rock masses at the slope toe has revealed that they can overlie recent alluvial and glacial deposits (Desio 1961; Terzaghi 1963), emphasizing that slow large rock mass movements can involve exceptional displacements (Figs 2b & 5). Lugeon & Oulianoff (1922) have shown the importance of toppling movements affecting entire

Fig. 3. Cross-sections showing three examples of ‘toppling’ in a slope near Martigny Switzerland or ‘Trois exemples de balancement superficiel des couches dans un versant’. (From Lugeon & Oulianoff 1922; reproduced with permission from the Socie´te´ Vaudoise des Sciences Naturelles.) The thickness of the affected zone can vary from a few metres to hundreds of metres.

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Fig. 4. Deep-seated gravitational slope deformation proposed by Terzaghi (1963).

slopes through the analysis of Alpine structures (Fig. 3). As previously noted, the observed deformational structures associated with slope movements are well documented, and have been variously described as ramps, folds thrusts, horst, graben and cataclastic bands (Choffat 1929; Terzaghi 1950; Voight 1973; Varnes 1978; Hutchinson 1988; Chigira 1992). Cruden (2000) made an attempt to classify slope-bedding attitude and potential failure mechanisms. Deformations at the level of large Alpine slopes were first observed during the 1940s (Ampferer 1939; Stini 1941; Terzaghi 1963), and the deep-seated gravitational slope deformations (DSGSD) were studied in detail by Zischinsky (1966, 1969), Nemcˇok et al. (1972) and Mahr & Nemcˇok (1977). Radbruch-Hall (1978) and Zaruba & Mencl (1982) summarized most of these features, providing an excellent illustration of uphill-facing scarps that was subsequently studied in detail by Bovis (1982). Recent work by Agliardi et al. (2001) and Ambrosi & Crosta (2006) shows the

importance of progressive slope failure mechanisms, and the close link between pre-existing tectonic structures and glacial unloading. They also showed that such slope deformations are the preferential site for the development of large landslides. A link between regional structural features and DSGSD is strongly suspected (Agliardi et al. 2001, 2009; Ambrosi & Crosta 2006) but remains enigmatic (Crosta & Zanchi 2000). It should be noted that in many slides the deformation is localized within a weak zone that is frequently coincident with bedding planes or foliation (Heim 1932; Cruden 1976; Braathen et al. 2004) or pre-existing lithological and tectonic features (Ambrosi & Crosta 2006). In addition, Agliardi et al. (2001) showed how active landslides can be found at the toe of slopes in extremely developed DSGSDs. These phenomena have been described as mass rock creep by Chigira (1992) and Chigira & Kiho (1994), who illustrated how many tectonic-like features (brittle faults and folds, cataclastic bands) can

Fig. 5. Cross-section showing the presence of an alluvial deposit below the sliding mass in the French Alps (modified after Gignoux & Barbier 1955).

INTRODUCTION

be produced at the slope scale by mass movement. These deformations lead to the degradation of the rock mass and are indicators of pre-failure and failure mechanisms. Chigira (1992) described these features as occurring in slopes of very different size, from a few tens of metres to thousands of metres in height. Depending on the material, the morphology and the presence of weakness zones or other external factors, the scale of gravitational deformations can vary from small to large sections of slope to entire mountain flanks and ridges (Beck 1968; Massironi et al. 2003) or multiple ridges and catchments (Giannini 1951; Argnani et al. 2003; Crosta et al. 2008a). Many ‘fossil mass movements’ can be recognized in sedimentary sequences in the form of olistostromes (Stow 1986), large slumps and chaotic deposits. The study of submarine landslides and of sedimentary formations show that such phenomena affect continental margins over distances of up to several hundreds of kilometres (Hampton et al. 1996; Locat & Meinert 2003), and that shear bands, folds and chaotic structures can exist at different scales. Finally, the limit between gravitational processes and purely tectonically induced movements can be quite difficult to identify (Fig. 6).

Mechanisms and processes The slope stability analyses performed for these types of phenomena require specific comment owing to the varied mechanical processes, and the variations in properties in space and time. Creep, progressive failure and degradation, spatial extent and continuity of the failure surface, and its changes with time (e.g. from a confined condition, as stated by Hutchinson 1988, to fully developed failure surfaces), together with the changes in type and intensity of perturbations with time, all need to be considered when dealing with DSGSD phenomena.

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The role of geological structures and the importance of pore fluid pressure in the generation and formation of otherwise inexplicable structures was investigated in early work carried out by Hubbert & Rubey (1959) and Rubey & Hubbert (1959). Inherited structures like bedding, joints, faults, folds and weak rocks control many slope instabilities (Agliardi et al. 2001; Brideau et al. 2009), but an important issue that must be considered to further the understanding of gravity-driven deformations is time-dependency characteristics and the occurrence of pre-failure mechanisms (Petley & Allison 1997; Sjo¨berg 2000; Leroueil 2001). It is clear that the failure of most slopes is, in fact, time dependent (Agliardi et al. 2001; Eberhardt et al. 2004; Stead et al. 2006; Eberhardt 2008); but the true nature of this time-dependent failure is not yet well understood, sometimes requiring an assumed mode of progressive shear strength decrease in order to simulate such phenomena (Eberhardt et al. 2004). The failure of rock bridges (Scavia 1995) is not yet used routinely in slope stability studies owing to difficulties in determining in the field the percentage of intact rock bridges within a slope. Most work to date has considered two-dimensional rock ridges; yet, the three-dimensionality of rock bridges adds a further degree of complexity to rock slope analysis as noted by Elmo et al. (2007). Nevertheless, the concept of progressive failure is interesting (Bjerrum 1967) because the use of localized failure is suitable for explaining many observed phenomena (Eberhardt et al. 2004) both in soil and in rock. Recent discrete element modelling approaches have emphasized the importance of considering fracture propagation in the failure of large rock slopes and the effect of progressive rock mass damage (Stead et al. 2006; Alzo’ubi 2009; Brideau et al. 2009; Lorig et al. 2009; Vyazmensky et al. 2010). The geometry of observed deformations and structures both within and at the surface of the examined slopes is not always scale dependent. In fact, the rheology of a material makes it possible

Fig. 6. Sketch illustrating landslides affecting valley slopes and regions indicating the different orders of landslide deformation.

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for a similar structure to be present at the centimetre level or at the mountain chain level, as demonstrated for faults by Riedel (1929) and by Skempton (1966). The geometry of the failure surface is strongly dependent on the boundary condition of the system. For instance, the La Clapie`re landslide (France) developed a failure mechanism involving toppling failure during its first evolutionary phase (Giraud et al. 1990) followed by the development of a continuous sliding surface. The driving mechanism could have been the dissolution and deformation of the gypsum and anhydrite (Fig. 7) lying underneath the slope (Guglielmi et al. 2000). Rock slope weakening and the landslide activity can be linked to tectonic activity, such as tectonic uplift and uplift gradient, or to seismic activity (Keefer 1984; Burbank et al. 1996; Crosta et al. 2008a, 2009; Hovius et al. 2009), and to water-table changes and cycles (Zangerl et al. 2003; Jaboyedoff et al. 2009). The importance of considering three-dimensional (3D) influences on large-scale rock slopes and landslides, particularly when modelling slope deformations, includes not only kinematical release provided by structures but the influence of 3D slope geometry, water pressures and effects on seismicity. The importance of considering 3D effects in the modelling of the kinematics of large rock slopes is demonstrated by Brideau & Stead (2010), Brideau (2010) and Ambrosi & Crosta (2011).

In summary, the slope tectonics approach proposed herein consists of the location and identification in the field of features associated with pre-slope failure and failure deformation phases, and features that provide information on the transition from one phase to the other, as well as the influence of controlling factors.

Volume contents This volume includes a selection of contributions presented at the Slope Tectonics Conference, which was held at the University of Lausanne (Switzerland) on 15– 16 February 2008. These contributions provide an indication of the current state of knowledge and of the most recent research progress in the field we refer to as ‘Slope Tectonics’. The following provides a brief overview of the various manuscripts included. The first paper presents an overview of the problem of understanding time-dependent slope deformation. El Bedoui et al. propose to study the incipient deformation history of a slope after deglaciation. Using their multidisciplinary approach, starting from field investigation to modelling, they show that the slopes are progressively destabilized along pre-existing structures but also through progressive failures. They support the hypothesis of the development of secondary landsliding within DSGSDs.

Fig. 7. Illustration of different potential interaction mechanisms between weak zones in a rock mass with resulting slope instability and slope deformation. The slope can be stressed by (a) undercut, (b) weak zone or (c) material deformation located below the slope.

INTRODUCTION

The following papers are dedicated to reactivation of ancient structures. Saintot et al. demonstrate the importance of pre-existing structures, such as faults filled with gouges developed along the foliation planes. In a regional study in Norway, Henderson & Saintot show how pre-existing structures control the slope susceptibility to rockslides in the Storfjorden area (western Norway). Bo¨hme et al. show how the development of rock instabilities promoted by several different mechanisms in Sogn and Fjordane County (Norway) are often linked to the knick point along a slope profile. Martinotti et al. suggest that the dissolution and deformation of rocks can induce progressive slope failures (Aosta Valley, NW Alps). Several papers deal with the development of landslides promoted by pre-existing structures that control the formation of a series of induced structures. Baron et al. show how palaeostress at the regional scale can be controlled by landsliding and gravitational deformation in the volcanic province of Boaco and Santa Lucia (Nicaragua, Central ˚ knes America). A 3D conceptual model of the A (Norway) landslide is proposed by Jaboyedoff et al. This paper shows that several pre-existing structures control the stepped failure surface, the lateral release plane and the movement direction. The failure surfaces activated by the slope movement follow weaknesses zones within the rock mass. Pedrazzini et al. describe the influence of pre-existing discontinuities linked to regional folding in slope stability of the Frank Slide scar (Turtle Mountain, Canada). Henderson et al., who show that landslide activity can develop along active faults, support the important role played by active pre-existing structures. The InSAR technique is adopted for this purpose at a site in Norway. The role of tectonic activity is also illustrated for the case study of the Caramanico area of the Central Apennines, as discussed by Bianchi Fasani et al. The use of numerical models for the understanding of the role played by morphology and preexiting structures is presented in two papers. Ambrosi & Crosta, using 3D mechanical modelling, investigate the influence of the orientation of weak elements within the rock mass (e.g. jointing, foliation, layering) and of slope morphology on the development of unstable sectors. Their results are supported by observations made in the field at different sites in the European Alps. In a case study Ghirotti et al. demonstrate the agreement between mechanical modelling and field observations at the Celentino deep-seated gravitational slope deformation (Peio Valley, NE Italy). Finally, Jarman et al., by inspection of anomalous fans and of surrounding slope morphologies in Val Venosta (east-central Alps, Italy), suggest how

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these could be reconstructed to demonstrate the occurrence of very large rock avalanches. The Slope Tectonics 2008 Congress received financial and logistical support from the Faculte´ des ge´osciences et de l’environnement of the University of Lausanne. The editor is grateful to the Advisory Committee, which supported this initiative: L.H. Blikra, Ch. Bonnard, A. Braathen, M. Chigira, J. Coe, R. Couture, M.-H. Derron, S.G. Evans, C.R. Froese, I. Henderson, V. Labiouse, J. Locat and A. Saintot, who gave their scientific support to the organization. We also thank the local group that helped us with the organization: J.-L. Epard, F. Baillifard, T. Oppikofer and A. Pedrazzini (IGAR). M. Jaboyedoff is also grateful to his colleagues at the Institute of Geomatics and Analysis of Risk at UNIL: P. Eraso, P. Horton, A. Loye, L. Perozzi, C. Schrocker and I. Spinello. We thank the reviewers without whom no publication would have been possible: F. Agliardi, A. Braathen, M.-A. Brideau, M. Chigira, J. Coe, R. Couture, E. Eberhardt, I. Evans, M. Geertsema, M. Giardino, J.S. Griffiths, A. Gu¨nther, R. Hermanns, M. Hurlimann, J. Hutchinson, O. Korup, C. Longchamp, A. Pedrazzini, G. Scarascia Mugnozza, A. Saintot, C. Squarzoni, and K. Sudmeier. Finally, we thank all the participants.

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Paraglacial gravitational deformations in the SW Alps: a review of field investigations, 10Be cosmogenic dating and physical modelling S. EL BEDOUI1, T. BOIS2*, H. JOMARD3, G. SANCHEZ4, T. LEBOURG2, E. TRICS2, Y. GUGLIELMI5, S. BOUISSOU2, A. CHEMENDA2, Y. ROLLAND4, M. CORSINI4 & J. L. PE´REZ6 1

LRPC Nancy, 71 rue de la Grande Haie, 54510 Tomblaine, France

2

GEOAZUR CNRS-UNS-IRD-UPMC, UMR 6526, Nice Sophia-Antipolis University, 250 Avenue de Albert Einstein, 06560 Sophia-Antipolis, France

3

Institute of Radioprotection and Nuclear Safety (IRSN), Fontaine-aux-Roses, France

4

GEOAZUR CNRS-UNS-IRD-UPMC, UMR 6526, Nice Sophia-Antipolis University, Avenue de Valrose, 06000 Nice, France 5

Centre de Se´dimentologie – Pale´ontologie, Universite´ de Provence Aix-Marseille 1, 3 Place Victor Hugo, 13331 Marseille Cedex 03, France

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Centre des Etudes Techniques de l’Equipement Me´diterrane´e, 52 Bd Stalingrad, 06520 Nice, France *Corresponding author (e-mail: [email protected]) Abstract: Catastrophic deep-seated landslides (DSL) are generally considered to be the result of large slope deformations also known as deep-seated gravitational slope deformation (DSGSD). This paper aims to build a synthesis of multiple studies made in the Tine´e Valley (southern French Alps) to assess the geometrical, kinematical, mechanical and chronological relationships between these two gravitational processes. At the scale of the valley, data issued from geological, geomorphological and 10Be dating indicate a clear geometrical link between DSGSD and DSL occurring at the base of the slope and suggest that gravitational slope evolution began after the glacial retreat (13 ka BP). This is supported by the example of the well-documented La Clapie`re slope. A continuous evolution process is characterized geometrically and temporally from geomorphic observations and analogue modelling. Coupling structural, geomorphological, physical and chronological studies allowed us to propose a four-dimensional (4D) deformation model mechanically correlated with progressive failure concept. The validity and variability of this reference site are discussed at the valley scale (taking Isola and Le Pra slope deformation as examples). It allows a rough estimation of the state of slope deformation at the valley scale to be constructed and the slope evolution with time to be considered. This 4D model could then be considered as a reference for other deep-seated gravitational slope deformations in comparable Alpine valleys.

Gravitational slope deformation plays an important role in relief evolution of mountain ranges (Jarman 2006); however, the interconnected processes leading from large spatial and timescale deep-seated gravitational slope deformation (DSGSD) (Dramis & Sorriso-Valvo 1994) to catastrophic rock slope failure remain poorly understood (Agliardi et al. 2001; Ballantyne 2002). Glacier retreat in Alpine valleys is often considered to be a major conditioning factor in slope destabilization (Evans & Clague 1994; Ballantyne 2002; Tibaldi et al. 2004; Bigot-Cormier et al.

2005; Hippolyte et al. 2006; Apuani et al. 2007). The main deglaciation effects on gravitational motion include topographic change of valleys (Savage & Varnes 1987; Augustinus 1995) and/or debuttressing of slopes leading to tensile stress state (Hutchinson 1988; Apuani et al. 2007). Both effects could strongly influence in situ stress conditions and rock strength parameters at the slope scale (Bachmann 2006). Structural heterogeneities, such as inherited tectonic faults and fractures, are also assumed to play a dominant role in gravitational slope failure

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 11–25. DOI: 10.1144/SP351.2 0305-8719/11/$15.00 # The Geological Society of London 2011.

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processes (Scavia 1995; Kaneko et al. 1997; Hermanns & Strecker 1999; Bachmann et al. 2004; Brideau et al. 2005; Jomard 2006; Bois et al. 2008). It is also commonly admitted that progressive failure within a rock slope can initiate and propagate through preferential weakened fault zones (Sartori et al. 2003; Bachmann et al. 2004; Willenberg 2004). This paper focuses on a post-glacial Alpine area: the Tine´e Valley (southern French Alps), which was affected by consecutive glaciations during the Quaternary period. Slopes located in this area are affected by recent DSGSD and deep-seated landslides (DSL) (Follacci 1987; Julian & Anthony 1996; Jomard 2006). This review paper is the synthesis of different approaches: field investigations, cosmonuclide dating of gravitational events and physical modelling that were performed at the valley scale and at the slope scale (the La Clapie`re slope) in order to: (i) establish chronological and physical links between DSGSD and DSL; and (ii) propose a model of slope deformations including some geomechanical considerations. To this end, the La Clapie`re slope is described in this paper as an observatory site used to understand the slope deformation evolution processes linking DSGSD and catastrophic rock collapses.

Regional-scale gravitational deformation in the upper Tine´e Valley General settings The Argentera –Mercantour massif is the southernmost external crystalline massif of the Western Alps. It is characterized by a polyphased deformation evolution from Variscan to Alpine orogenies (Corsini et al. 2004). The basement consists of highgrade metamorphic and intrusive rocks of late Carboniferous age (Ferrara & Malaroda 1969). It is unconformably covered by a marine sedimentary succession of Late Carboniferous –Cenozoic age, partly detached at the level of the Triassic evaporites (Faure-Muret 1955) and overthrusted by Penninic clastic units during the Late Eocene –Early Oligocene (Autapie nappe and Parpaillon nappe) (Tricart 1984). The upper Tine´e Valley represents the western boundary between the basement with its Permo-Triassic tegument and the detached Mesozoic sedimentary cover (Figs 1 & 2). The basement rocks are migmatitic paragneisses with meta-granodioritic intrusions (Faure-Muret 1955). Ductile fabrics present a global N1308 dipping to the NE foliation. The main slope directions are collinear to the major N1108 –1408 trending fault set of the massif and make a 708 angle towards a secondary N0008 – 0308 fault system (Fig. 1).

The Argentera –Mercantour massif and its foreland has been a high relief area since the Early Pliocene (Fauquette et al. 1999). Altitudes range between 400 and 3143 m a.s.l. (metres above sea level) (at Mt Ge´las) in the gneissic bedrock and 3051 m (at Mt Pelat) in the sedimentary cover (Fig. 1). The combined glacial and river network have deeply eroded and incised the massif, resulting in slope heights of 2000 m. The glaciations history of the area is complex and poorly documented but as in many other Alpine valleys, the morphology of the upper Tine´e area was strongly influenced by Rissian and Wu¨rmian glacial ages. Most of the current glacial landforms come from the Wu¨rmian glaciation. They are characterized by glacial polished surfaces of glacial deposit (Julian 1980). From 10Be dating of polished glacial surface and radiocarbon analysis, Bigot-Cormier et al. (2005) and Sanchez et al. (2009) reported that the Tine´e Valley glacier was totally deglaciated at 12 ka BP whereas higher slope parts were deglaciated around 8 ka BP (Julian 1980). Besides, successive positive pulses since the last main deglaciation have been evidenced in the massif, highlighting a discontinuous deglaciation since the Wu¨rm (Julian 1980). However, glaciers have not been as important as in the other northernmost massifs in the Alps with a southern extension of the ice to a 500 m a.s.l. altitude and a maximum thickness of 500 m in the valley. The summits were also not covered by any ice sheets. As a consequence, in the morphology of the Tine´e Valley the mean slope angles vary from 358 up to 1800 m to 258 above this altitude.

Large gravitational deformation inventory An inventory was established by Jomard (2006) (see also Bois et al. 2008). A distinction is made between deformations that are clearly linked to gravity (bounded by a failure surface and a large mobilized volume that correspond to DSL), and deformations that appear diffused in very large volumes of the massif. Such deformations display extensive graben-like features and correspond to large sagging zones of the slopes. Zischinsky (1966) first proposed the term sackung for those surface manifestations of deep-seated rock creeps of foliated bedrocks in the Alps. This kind of movement was then observed in almost all mountain ranges and most authors today use the generic term deep-seated gravitational slope deformation (DSGSD) introduced by Dramis & Sorriso-Valvo (1994) to name the landforms and geomorphic evidence, such as double crested ridges, troughs, antislope scarps and ridge depressions, associated with those deformations (Agliardi et al. 2001). The origin of these features (tectonic, gravitational or both) are still poorly understood even if recent

GRAVITATIONAL DEFORMATION IN THE SW ALPS

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Fig. 1. Major tectonics features and gravitational deformations mapped in the Tine´e Valley (modified after Bois et al. 2008).

results strongly tend to demonstrate a gravitational origin (Bachmann 2006). These characteristic morphostuctural features were mapped at the Tine´e Valley scale (Jomard 2006). Field investigations within the Tine´e slopes showed that most of Wu¨rmian glacial morphologies

and deposits are affected by gravitational deformations, suggesting gravitational motions since the last deglaciation (Jomard 2006). Most of these observations have been made in the metamorphic basement, in the left bank of the valley. Eight DSGSD zones were recognized containing 20 DSL (Fig. 1).

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Fig. 2. The La Clapie`re rockslide embedded in the Colle Longue DSGSD, showing the geometrical relationship between the inherited tectonic structures, DSL and DSGSD (modified after Jomard 2006).

A regional map representing DSGSD and DSL deformations (morphostructures) was compared to the structural map of the massif (Fig. 1). † The orientations of DSGSD’s morphostructures are strongly influenced by both the inherited tectonic framework and the mean N1308-trending foliation. In particular, the major N1108– 1408E fault set explains the localization of most of the internal deformations of DSGSD within the valley. † The other N0008–0308E fault set is also expressed in some cases through deformations secant to the main direction. † The most developed DSL occurred at the intersection of two fault set, as the La Clapie`re DSL between the N1108 –1408E and N0008–0308E fault set. The density of morphostructures within DSGSDs appears to be closely dependent on the angular relation between slope direction, foliation and faults orientations. Morphostructures are much more developed when this angle is close to zero. The most developed Colle Longue DSGSD could then be explained by the narrow angle existing between the N1408E main direction of the crests/ valley and the N1308E main faults and foliation direction. The Colle Longue DSGSD represents an area of about 45 km2. This zone, which presents extensive deformations spreading from the foot to the crests, has been chosen for a more accurate description. Counterscarps are recognized in the upper slope part down to 1800 m a.s.l., although troughs are observed in the lower slope part up to 2100 m a.s.l. Counterscarps can reach 20 m high and always are guided by the foliation planes whose main orientation is N1308–0408 NE. It also appears that counterscarps connect at depth to inherited tectonic

faults that guide the gravitational deformations. Troughs display 15 m wide apertures that are geometrically associated to downward-dipping (SW) normal and strike-slip fault zones. Both structures affect the slope down to the tributary valleys floor indicating that deformations should be deeper than these valley incisions that can reach 1000 m high. No failure surface enveloping the DSGSD was observed. Compressive features were only observed in some DSL feet encased in the DSGSD zone (like the large active La Clapie`re DSL). A long-term deformation of the slope is then characterized after the deglaciation. From the chronological point of view, two main observations are made: † In upper slope parts (.1800 m), gravitational deformations affect Wu¨rmian glacial morphologies of high-altitude tributary glaciers, rock glaciers and active screes. Glacial sediments filling cracks are also observed in a number of counterscarps. Deformations are mostly represented by counterscarps and scarps. † In lower slope parts, deformations affecting Wu¨rmian deposits and morphologies are represented by scarps. Troughs are filled with colluviums and no glacial sedimentary fillings are observed. On the La Clapie`re slope (Jomard 2006), a trough located on a down-slope normal fault was detailed (Fig. 3). The fault, characterized by its gouge and slickenslide planes, was recently toppled and created a large tension aperture (10 m) infilled by regular post-glacial colluviums sedimentation. The toppling of rock columns allowed an intense fracturing of the downward wall of the fault. Finally, this intense fracturing gives the possibility for sliding surfaces to develop and cross-cut the overall structure (Fig. 3). Chronological constraints based on 10Be dating of gravitational scarps, troughs and landslide surfaces in the Tine´e area show three successive periods of gravitational instabilities (Bigot-Cormier et al. 2005; Sanchez et al. 2009). The first closely post-dates the last deglaciation event (12–13 ka) with an age of 10–11 ka, a second destabilization occurs at 7 –9 ka BP and a third at 2.5– 5.5 ka BP. Thus, gravitational slope deformations have been effective at least since 12 ka BP, leading to the present active large deep-seated landslides such as La Clapie`re.

Slope-scale analysis of La Clapie`re Field analysis The La Clapie`re slope (Fig. 2) is one of the most active DSLs of the valley. This DSL affects 60  106 m3 of rock of the metamorphic basement.

GRAVITATIONAL DEFORMATION IN THE SW ALPS

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Fig. 3. Trough cross-section (modified after Jomard 2006). Tilted normal fault gouge infilled with post-glacial colluvium affected by recent failure. (a) Photograph and (b) interpretation.

The first activity reported dates from the early 1930s, with a peak evolution between 1960 and 1990 leading to a 130 m-high scarp development in the middle of the slope. Thus, since the 1990s, structural investigations (Follacci 1987; Ivaldi 1991), geomechanicals triaxial experiments on rock samples, hydrogeological studies (Cappa et al. 2004; Guglielmi et al. 2005), numerical modelling and geophysical surveys (Lebourg et al. 2005; Jomard et al. 2007, 2010) have been performed underlining a complex post-failure behaviour. The La Clapie`re DSL is geometrically bound by N1108–1408E and N0108–0308E fault sets (Fig. 1). Other morphological signs of gravitational destabilization are guided by those tectonic orientations, and are represented inside and outside the DSL boundaries, mainly consisting of troughs aligned on tectonic fault scarps (Fig. 3). These troughs show, respectively, a trace linear far from the active DSL, curved close to the DSL and dislocated within the landslide body. Fifteen troughs were mapped with an average N1208 direction and a 100 –5000 m length (Fig. 4) that indicated a deep slope deformation in agreement with the observations made in the valleys bounding the slope (Fig. 2). Troughs were mainly observed from the toe to the middle of the slope (1500–2100 m in elevation) and their evolution is clearly linked to the initiation of the La Clapie`re DSL. Their average orientation is parallel to the slope, although troughs closed to the western boundary of the currently active DSL have an orientation close to N1308 (Fig. 4). Troughs 1– 4 that are the closest to the active DSL scarp are highly deformed (and even cut by the scarp) while torsion progressively vanishes from troughs 5 –15. Troughs 1, 6 and 14, and the west

lateral scarp propagation of La Clapie`re active DSL, were dated using an in situ produced 10Be cosmogenic approach (Bigot-Cormier et al. 2005; Sanchez et al. 2009). The result suggest that the troughs become progressively younger from trench 1 (10 ka BP), 6 (7.2 ka BP) to 14 (5.6 ka BP), meaning that a deformation of the slope propagated from the toe to the top of the slope at about 4.4 ka BP. The upper lateral scarp of the currently active DSL was dated at 3.6 ka BP, showing that after the up-slope deformation propagation a deep failure initiated in the middle part of the slope that ultimately bounds the currently active DSL. Two-dimensional (2D) physical modelling experiments were performed reproducing a NNE– SSW cross-section of the La Clapie`re slope in order to analyse the links between superficial and deep-seated deformations (Figs 5 & 6).

Two-dimensional physical modelling of the ‘La Clapie`re’ slope A complete description of the analogue material (called Slope1) and the loading device developed to perform our scaled physical models is available in Bachmann et al. (2004), Chemenda et al. (2005) and in Bois et al. (2008). A short description is available in the Appendix of this paper. For these models a scaling factor of 1/50 000 was chosen. The vertical faults have been numbered from F1 near the valley toe to F6 near the crest. Two distinct configurations of the slope were tested: † a homogenous model (Fig. 5) that must be considered as a massive homogeneously fractured massif without any major localized weak zone;

16 S. EL BEDOUI ET AL. Fig. 4. Evolution of the La Clapie`re slope deformation for the last 10 ka BP. (a) and (b) Troughs opening from 1600 to 2250 m elevation a.s.l.; (c) surface shearing close to the future rockslide area; (d) rockslide collapse and sliding (modified after El Bedoui et al. 2009).

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Fig. 5. Homogeneous model of the La Clapie`re slope: (a) the non-deformed model; (b) the first deformation stage and its interpretative sketch; and (c) the last deformation stage and its interpretative sketch (modified after Bois et al. 2008).

† a model considering the N1408 fault zones previously presented (Fig. 6). The listric geometry of those faults was deduced from field investigations (Fig. 2). Homogeneous model. On the initial deformation stage (Fig. 5b), a deep sliding surface was formed inside the model.

Its maximum depth was equivalent to 1500 m. This sliding surface bounded a large unstable volume involving the entire massif. A 100 m- high scarp was formed behind the topographic crest. On the final deformation stage, the sliding surface becomes a more complex fracture network. Its width increased with the displacement of the

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Fig. 6. A slope model of La Clapie`re cut by six inherited normal listric faults: (a) the non-deformed model; (b) the first deformation stage and its interpretative sketch; (c) the second deformation stage and its interpretative sketch; and (d) the last deformation stage and its interpretative sketch (modified after Bois et al. 2008).

sliding unit (Fig. 5c). This complex failure zone reached a maximum depth of 2200 m and was associated with two antithetic newly formed faults. The escarpment located behind the topographic crest kept growing to reach a maximum of 400 m. A slope cut by six inherited normal listric faults. During the initial deformation stage (Fig. 6a), most of the superficial non-elastic deformation

was localized on three inherited normal faults and on a newly formed gravitational fault. On the inherited fault close to the topographic crest (F6) a 100 m-high escarpment appeared (Fig. 6b). On the inherited faults close to the slope toe (F2 and F3) two troughs appeared (Fig. 6b). Finally, a newly formed gravitational fault formed at the slope toe that propagated inside the massif rather like a subhorizontal thrust fault

GRAVITATIONAL DEFORMATION IN THE SW ALPS

(Fig. 6b). Internal irreversible deformations took place along this newly formed thrust fault and also along an antithetic normal fault initiating from the inflection point of fault F6 towards the topographic surface (Fig. 6). During the second deformation stage, the deformation increased (Fig. 6c): another normal trench appeared at the top of the fault F5. The newly formed thrust fault propagated inside the model and a relatively small superficial DSL was finally triggered between the first normal fault and the slope toe. In the last deformation stage (Fig. 6d), the sliding surface propagated inside the model through the connection of the inflection points of faults F4– F6, delimiting a deep gravitational moving zone. The first antithetic fault propagated and reached the topographic surface. The hillside was affected by a second DSL, which was a retrogressive one. The fractured model deformation can then be summarized as follows. † The normal troughs are formed on the faults close to the valley immediately followed by the appearance of the first normal shifting on faults close to the topographic crest. † The first antithetic faults are formed, and a sliding surface (i.e. thrust fault) propagates from the toe. † New troughs are created higher on the slope due to a retrogressive deformation process. The sliding surface propagates and a small-scale DSL is triggered. † The sliding surface is connected to the F6 fault delimiting the total moving mass. The deformation of this part of the slope leads to the formation of successive landslides that affect the slope. Modelling major results. On one hand, the deformation pattern obtained in the homogeneous configuration showed that even if the kinematics of the rupture is coherent, the localization of the deformation is not evident. Indeed, it seems obvious that the ‘La Clapie`re’ slope cannot be considered to have fractured homogeneously. Some localized weak zones have to be taken into account, such as the N1408 faults zones. This is confirmed by the second model configuration for which it appears clearly that the deformation is mainly localized on pre-existing faults. However, on the other hand, those models are 2D models developed according to a NNE–SSW cross-section of the ‘La Clapie`re’ slope, and owing to this simplification some structures have not been considered (e.g. other inherited faults such as the N0308 ones). It is reasonable to say that those structures must also have had an influence on the

19

localization of the deformation and, at a larger scale, on the global deformation pattern of the massif.

From DSGSD to catastrophic failure at the slope scale Based on field investigations and absolute ages (10Be), El Bedoui et al. (2009) calibrated a model of the slope evolution for the last 10 ka. Physical modelling shows a very good agreement with this proposed model based on field and dating work underlining the link between those superficial deformations and the failure propagation at depth. From 10 to 5.6 ka BP (Fig. 4) extensional structures (troughs) spread from the toe to the top of the slope, showing a good correspondence with physical modelling (Fig. 6b, c) (phase I). The trench-like morphology is explained by the reactivation of inherited structures that could have been induced by the stress release effects in the slope related to glacial retreat of the last glaciation. This phase I could be related to the deep fracture retrogressive propagation showed by physical modelling (Fig. 6c, d). From 5 to 3.6 ka BP (Fig. 4) a shearing of troughs occurred in the western lower part of the slope that also displayed a high vertical displacement (phase II). This 3D twisting of surface troughs, which is related to shear deformations developing deep in the slope, could be induced by preferential tangential movements along a major N0308 reactivated vertical fault zone located in the east part of the slope. At 3.6 ka BP (Fig. 4) a failure initiates, wrapping this eastern lower part of the slope. In the last 50 years this failure evolved in a large-scale failure surface bounding the currently active La Clapie`re DSL (phase III). 10 Be ages indicate that phase I (troughs opening – DSGSD) propagates over a very long period (6 ka BP). Phase II (troughs shearing) is only constrained by two dated events: the trench opening (10 ka BP) and the first failure associated with the rockslide (3.6 ka BP). If physical modelling does not allow a kinematic model of the slope evolution, results underline the fact that the first rockslide event was synchronous with the upper and younger trench (dated at 5.2 ka BP). It could indicate that the trench shearing more probably occurred between 5.2 and 3.6 ka BP rather than immediately after the trench opening at 10 ka BP. The kinematics of slope deformation have been characterized in two dimensions and extended in three dimensions, and suggest a non-linear creep-like phenomenon. The significance of the deduced model needs to be discussed in a mechanical way and compared to other places in the valley to assess its validity.

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Discussion Mechanical processes at the slope scale Rock slope failure in general results from many different processes and among these the influences of the pre-existing heterogeneities (such as bedding planes, foliation and fault zones) (Terzaghi 1962; Kato & Hada 1980; Chigira 1985; Agliardi et al. 2001) the mountain height and slope gradient appear to be predominant. Rock slope stability is thus highly dependent on the specific characteristics of the fault zones (e.g. density of the fault network, persistence at depth, the geometry of the faults, etc.) (Hermanns & Strecker 1999). Our study underlines the major role played by regional scale weak planes that correspond to inherited major fault zones (Bois et al. 2008). The temporal evolution of slope deformation at La Clapie`re is in good accordance with laboratory experiments that display highly non-linear creeplike phenomena even for hard rocks. In previous work conducted on the gravitational evolution of rock slopes (Bru¨ckl & Parotidis 2005; Petley et al. 2005; Apuani et al. 2007) the evolution from DSGSD to a localized catastrophic failure is physically and mechanically regarded as a creep-like phenomenon. Bru¨ckl & Parotidis (2005) numerically studied the slope instabilities resulting from deep-seated gravitational creep and the transition from a slow evolution phase to rapid sliding. The process of subcritical growth was considered to explain the primary phase of deep-seated gravitational creep because it allowed the progressive damage of the rock mass at a lower stress than the rock strength. It could correspond to the progressive failure growth at depth characterized by troughs opening and twisting at the surface. The total failure at depth then outlines the currently active rockslide, acceleration of the movements being related to the ‘smoothing of the basal surface’. Another first-order process that has an influence on strength reduction, and thus on damage process, is weathering controlled by climatic and fluids circulations (Hill & Rosenbaum 1998; Hall & Andre´ 2001; Pellegrino & Prestininzi 2007). It also has a strong control on the failure process, especially in the case of granitic rocks in the Alps (Girod 1999; Jaboyedoff et al. 2004). Indeed, the alteration of such rocks has two main influences: on the one hand, it leads to the formation of clays; thus reducing the fluid circulation and increasing the pressure (Girod 1999). This is particularly true in faulted zones where alteration is concentrated basically along localized weak zones: fractures and faults (Migon & Lidmar-Bergstro¨m 2002; Wyns 2002). On the other hand, weathering causes a progressive strength reduction of the rock material, which is

stress dependent. This softening is generally maximal at the surface and diminishes with depth (Chigira 2001). Even if this contributing factor has not been taken into account by the physical models, Chemenda et al. (2009) used a 2D finite-element numerical model to show that, in the case of the La Clapie`re slope, a progressive reduction in the mechanical resistance (due to a progressive reduction in the model cohesion), combined with the particular geometry of glacial Alpine valleys, can lead to the DSL. A time-dependent model (Fig. 7) calibrated on the La Clapie`re slope failure was proposed in order to fit the observed surface displacements as function of time (El Bedoui et al. 2009). This model presents three phases: (1) very slow displacements over a long time period (several mm year21); (2) an increase in surface displacements related to a deep slide plane (crack coalescence); and (3) a catastrophic evolution over a very short time period.

Extrapolation of the model at the valley scale The La Clapie`re slope is part of the Colle Longue DSGSD. Looking at the Tine´e Valley scale, different evolution stages were observed, from pre-failure stages characterized by troughs dislocation, to active and fossil DSL. Two representative localities were studied: the Isola slope and the Pra slope (Figs 1 & 8). The Isola slope is located downstream of the upper Tine´e Valley. This slope is also embedded in the Colle Longue DSGSD (Figs 1 & 8). The average slope is 358 in the basal part (from 850 to 1700 m elevation) and 258 in the upper part. A large series of N1208 troughs crop out in the basal part, extending laterally over distances of 500 m. A surface shearing of these troughs was located at elevations of between 1000 and 1500 m, which is similar to the La Clapie`re case and indicates a contrast of surface velocities between the upper stable part of the slope and the lower more intensely deformed area. The lower slope part shows signs of shearing that are as developed as those on the La Clapie`re slope (Fig. 8). It indicates that this slope can correspond to the second stage of the evolution model of slope deformation (Figs 6 & 7). This advanced stage of slope deformation is further exacerbated by very active rockfall activity and may indicate a future landsliding activity during the next century.

Conclusion Coupling the different approaches presented in this paper has allowed the elaboration of a 4D evolution

GRAVITATIONAL DEFORMATION IN THE SW ALPS

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Fig. 7. Curve of time-dependent progressive failure calibrated on the La Clapie`re slope (modified after El Bedoui et al. 2009).

model of rock slope destabilization linking failure propagation at depth and subsequent deformations at the surface. The model seems to be well constrained for the La Clapie`re slope and seems also to be valid at the Tine´e Valley scale. Furthermore, as the spatial correlation between gravitational deformation and weakened zones of inherited faults suggest, the major structural framework has a key role in such processes. Thus, by comparison with other slopes, we suggest that kinematics of slope evolution are controlled by a critical spacing/ arrangement of major inherited faults. De facto, the deglaciation and the consecutive slope angle increase of the valley is supposed to have a strong influence on slope destabilization due to modifications of in situ stress condition and rock strength parameters. The main implication of these results is that the model proposed based on the La Clapie`re slope can be considered as a reference at the valley scale and, probably more widely, at the massif scale as proposed by Jomard (2006). Large-scale geomorphological studies (Jomard 2006) should be coupled with instrumental surveys (GPS) performed at the valley scale and mechanically scaled models; this

will, respectively, allow: (i) an estimation of volumes potentially mobilized during the destabilization (based on deformation of morphological signs; i.e. troughs); and (ii) a rough estimation of time prior to failure by comparison between instrumental velocities and the reference calibrated on the La Clapie`re slope. Authors sincerely express their gratitude to the GIS CURARE who financed this project and to the reviewers for their useful remarks.

Appendix Slope1 is a low frictional elasto-brittle– plastic analogue material with strain softening (Chemenda et al. 2005). This material represents a compositional system based on liquid and solid hydrocarbons. To create a model, the melted analogue material Slope1 is moulded into a rigid box at a temperature of 50 8C. In order to create the faults a series of openings cut in the two opposite lateral sides of the model box are used to position taut strings. After cooling to a temperature of 20 8C, at which the crystallized material is strong enough to be easily

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Fig. 8. Morphostructures of the Isola slope.

handled without damage and easily cut with the strings without being damaged in areas other than the cuts, strings are translated along the slots to generate the faults and then removed. The model surface was then shaped to the desired topography (Figs 5a & 6a). The length of the model is thus equal to 14 cm and the width (third dimension) is equal to 30 cm. The third dimension has been chosen large enough to prevent any edge effects. Once the model is prepared it is loaded into a vertical accelerator table. The latter consists of a mobile platform that can be lifted up to 2 m and then released. During

its free fall the model reaches a maximum velocity of 6 m s21. The platform is then rapidly but smoothly decelerated to zero velocity when it comes into contact with a progressive shock absorber of 5 cm stroke. During this phase the model undergoes a strong vertical deceleration (up to 500 m s22). This deceleration acting in the same direction as gravity is repeated until failure develops, usually at approximately 100 cycles. Preliminary calibration tests are needed to determine which acceleration must be imposed onto a model for a given configuration (geometry, prefracturing state, etc.) in order to observe

GRAVITATIONAL DEFORMATION IN THE SW ALPS failure for a number of loading cycles ranging from 100 and 150. Model deformation can be observed accurately after each acceleration cycle. This discrete loading technique has proved to be equivalent to a continuous quasi-static loading (Chemenda et al. 2005). The main similarity criterion is:

sco scm ¼ ro go H o rm gm H m

(1)

where r is density and g is gravity acceleration, sc is the strength under uniaxial compression and H is the spatial scale of the phenomenon (the mountain height H, for example). The superscripts ‘o’ and ‘m’ mean original and model, respectively. To ensure that the deformation will be localized along brittle structures, as in nature, Slope1 has to exhibit a high degree of softening. Hence, the experiments were carried out at a fixed temperature of 20 8C. This mechanical behaviour is comparable to the strength degradation behaviour introduced into some numerical models. At this temperature the coefficient of friction measured on the pre-existing fractures is m ¼ 0.2. Cross-sections were made at the end of each experiment by cutting the model at various positions after cooling it to 10 8C in order to increase its strengths. Some experiments were stopped in the early stages of model deformation to analyse the corresponding evolution of internal slope deformation.

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Jomard, H., Lebourg, T. & Tric, E. 2007. Identification of the gravitational boundary in weathered gneiss by geophysical survey: La Clapie`re landslide (France). Journal of Applied Geophysics, 62, 47–57. Julian, M. 1980. Les Alpes Maritimes Franco-Italiennes. Etude Ge´omorphologique. PhD thesis, Aix-Marseille II University, Marseille. Julian, M. & Anthony, E. 1996. Aspect of landslide activity in the Mercantour Massif and the French Riviera, southeastern France. Geomorphology, 15, 275–289. Kaneko, K., Otani, K. J., Noguchi, Y. & Togashiki, N. 1997. Rock fracture mechanics analysis of slope failure. Deformation and Progressive Failure. In: Asaoka, A., Adachi, T. & Oka, F. (eds) Geomechanics. Elsevier, New York, 671– 676. Kato, J. & Hada, S. 1980. Landslides of the Yoshino– Gawa water system and its geological aspects. Research Reports of the Kochi University. Natural Science, 28, 127– 140. Lebourg, T., Binet, S., Tric, E., Jomard, H. & El Bedoui, S. 2005. Geophysical survey to estimate the 3D sliding surface and the 4D evolution of the water pressure on part of a deep seated landslide. Terra Nova, 17, 399– 406. Migon, P. & Lidmar-Bergstro¨m, K. 2002. Europe: problems of dating and interpretation of geological records. Catena, 49, 25– 40. Pellegrino, A. & Prestininzi, A. 2007. Impact of weathering on the geomechanical properties of rocks along thermal-metamorphic contact belts and morphoevolutionary. Geomorphology, 87, 176 –195. Petley, D. N., Mantovani, F., Bulmer, M. H. & Zannoni, A. 2005. The use of surface monitoring data for the interpretation of landslide movement patterns. Geomorphology, 66, 133–147. Sanchez, G., Rolland, Y., Corsini, M., Braucher, R., Bourles, D., Arnold, M. & Aumaıˆtre, G. 2009. Relationships between tectonics, slope instability and climate change: cosmic ray exposure dating of active faults, landslides and glacial surfaces in the SW Alps. Geomorphology, 117, 1– 13; doi: 10.1016/ j.geomorph.2009.10.019. Sartori, M., Baillifard, F., Jaboyedoff, M. & Rouiller, J. D. 2003. Kinematics of the 1991 Randa rockslides (Valais, Switzerland). Natural Hazards and Earth System Sciences, 3, 423– 433. Savage, W. Z. & Varnes, D. J. 1987. Mechanics of gravitational spreading of steep-sided ridges (sackung). Bulletin of Engineering Geology and the Environment, 35, 31–36. Scavia, C. 1995. A method for the study of crack propagation in rock structures. Ge´otechnique, 45, 447–463. Terzaghi, K. 1962. Stability of steep slopes in hard unweathered rock. Ge´otechnique, 12, 251– 270. Tibaldi, A., Rovida, A. & Corazzato, C. 2004. A giant deep-seated slope deformaion in the Italian Alps studied by paleoseismological and morphometric techniques. Geomorphology, 58, 27– 47. Tricart, P. 1984. From passive margin to continental collision; a tectonic scenario for the western Alps. American Journal of Sciences, 284, 97–120.

GRAVITATIONAL DEFORMATION IN THE SW ALPS Willenberg, H. 2004. Geologic and kinematic model of a complex landslide in crystalline rock (Randa, Switzerland), PhD thesis, Swiss Federal Institute of Technology Zurich. Wyns, R. 2002. Climat, eustatisme, tectonique: quels controˆles pour l’alte´ration continentale? Exemple des se´quences d’alte´ration ce´nozoı¨ques en France.

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Bulletin d’Information Ge´ologie des Bassins, Paris, 39, 5 –16. Zischinsky, U. 1966. On the deformation of high slopes. In: Proceedings of the First International Conference of the International Society of Rock Mechanics, Lisbon, Volume 2. International Society of Rock Mechanics, Lisbon, 179– 185.

Inheritance of ductile and brittle structures in the development of large rock slope instabilities: examples from western Norway A. SAINTOT1*, I. H. C. HENDERSON1 & M.-H. DERRON1,2 1

Geological Survey of Norway, Leif Eirikssons Vei 39, N-7491 Trondheim, Norway

2

Lausanne University, IGAR, Quartier UNIL-Sorge, CH-1015 Lausanne, Switzerland *Corresponding author (e-mail: [email protected]) Abstract: The high density of slope failures in western Norway is due to the steep relief and to the concentration of various structures that followed protracted ductile and brittle tectonics. On the 72 investigated rock slope instabilities, 13 were developed in soft weathered mafic and phyllitic allochthons. Only the intrinsic weakness of such rocks increases the susceptibility to gravitational deformation. In contrast, the gravitational structures in the hard gneisses reactivate prominent ductile or/and brittle fabrics. At 30 rockslides along cataclinal slopes, weak mafic layers of foliation are reactivated as basal planes. Slope-parallel steep foliation forms back-cracks of unstable columns. Folds are specifically present in the Storfjord area, together with a clustering of potential slope failures. Folding increases the probability of having favourably orientated planes with respect to the gravitational forces and the slope. High water pressure is believed to seasonally build up along the shallow-dipping Caledonian detachments and may contribute to destabilization of the rock slope upwards. Regional cataclastic faults localized the gravitational structures at 45 sites. The volume of the slope instabilities tends to increase with the amount of reactivated prominent structures and the spacing of the latter controls the size of instabilities.

In hard rock masses, the steepness of high slopes may not be the unique parameter to explain the development of large gravitational instabilities. Many studies document the role of pre-existing structural grain and tectonic weakening in the development of large instabilities in hard rock, and it is widely accepted that the failure surface(s) may be partly or fully structurally controlled (e.g. Nemcok et al. 1972; Cruden 1976; Varnes 1978; Giraud et al. 1990; Chigira 1992; Scavia 1995; Guzzetti et al. 1996; Julian & Anthony 1996; Sauchyn et al. 1998; Agliardi et al. 2001; Kellogg 2001; Krejcˇ´ı et al. 2002; Bachmann et al. 2004; Eberhardt et al. 2004; Brideau et al. 2005; Vilı´mek et al. 2007; El Bedoui et al. 2008; Brideau et al. 2009; Jaboyedoff et al. 2009). Worldwide, the structural development of many large rock slope instabilities may be controlled by regional faults (Kaneko et al. 1997; Sartori et al. 2003; Ambrosi & Crosta 2006; Bois et al. 2008). A large number of studies showed the link between high frequency sets of persistent joints and slope susceptibility to rockfalls (and especially discrete rockfalls as defined in Varnes 1978). Indeed, most of the pre-existing structures described in these studies belong to the domain of brittle deformation, and few publications demonstrate the influence of the pre-existing ductile fabrics and folds in the development of slope instabilities (Jones 1993; Badger 2002; Coe & Harp 2007). In Norway, the bedrock

geology reveals an extremely complex and protracted ductile and brittle tectonic history that may intuitively have a strong influence on the spatial distribution of instabilities along the steep slopes of the Norwegian coastal fjord areas. An intense collection of 5000 structural measurements allowed us not only to place 72 rock slope instabilities in a regional structural framework but also to precisely investigate the role of geology and specifically structural geology in the development of the rock slope instabilities. These were not only the gravitational structures, thus related to the slope instabilities themselves (open large fractures, basal sliding planes, transfer structures, etc.), but also the older structures (exfoliation, foliation, folds, joints, fault zones, etc.) that belong to the prolonged geological history of the area. With the field observations at these 72 sites, we established the clear relationships between the location of the failure surface(s) and pre-existing geological structures. We present in detail in this paper the sites that best exemplify the relationships between slope instabilities and inherited structures. We describe case studies that illustrate: (1) how the ductile structures (metamorphic foliation and folds) may focus the formation of gravitational structures; and (2) reactivation of brittle structures to produce gravitational failures. We also appraise the relationship between the estimated unstable volume, the number of involved preexisting structures and the type of the gravitational

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 27–78. DOI: 10.1144/SP351.3 0305-8719/11/$15.00 # The Geological Society of London 2011.

28

A. SAINTOT ET AL.

slope deformation. Using field examples, we attempt to show how a combination of large and favourably orientated inherited structures (as fold hinges, brecciated/cataclastic tectonic faults) in a slope may specifically determine the final volume of the gravitationally deformed slope. We focus on identifying the geological factors that may render the steep slopes of western Norway more susceptible to instability development. Besides lithological factors (weak v. hard rocks) and the abundance of pre-existing planar structures (favourably orientated with regard to the slope), we have the opportunity in the Storfjord area to determine how high-grade metamorphism polyphase ductile folding may enhance slope susceptibility. We also examine the stability of some of the studied rockslides by a quantification of cohesion values and friction angles of shallow-dipping basal sliding planes.

An overview of the geology and rock slope instabilities in western Norway A large variety of mostly Lower Palaeozoic and Precambrian metamorphic rocks cover western Norway, and the area under investigation is known as the Western Gneiss Region (Fig. 1). The latter refers to the setting of the early Palaeozoic Caledonian Orogeny where the abundant gneissic rocks are found in both autochthonous and allochthonous positions. During the Caledonian Orogeny, the rocks were severely reworked by a general NW– SE-oriented bulk crustal shortening that resulted in a SE-directed nappe transport, in regimes ranging from the deepest ductile domain to the brittle– ductile transition (Roberts & Gee 1985; Hossack & Cooper 1986; Roberts 2003). Poorly documented Precambrian events (Tucker et al. 1990; Ska˚r & Pedersen 2003; Røhr et al. 2004) remain recorded in the protoliths of the Caledonian metamorphic rocks. In Phanerozoic times, after the Devonian semi-ductile syn- or post- orogenic Caledonian collapse (Hossack 1984; Fossen 2000), intense brittle tectonic events affected western Norway including the Permo-Triassic and Jurassic rifting phases and the opening of the North Atlantic Ocean (Torsvik et al. 1997; Valle et al. 2002; Mosar 2003). However, the stratigraphic records of these tectonic phases are lacking onshore western Norway and the major faults that pertain to the listed brittle tectonic events are far from being accurately identified. Started from a recent but still debated time, an ongoing high-rated uplift brings the intensively tectonized rocks to the surface. Norway also underwent significant glacial and post-glacial tectonic activity including large-magnitude earthquakes at the time of ice-cap melting with triggered rockfalls and possibly avalanches, isostatic rebound and further

neotectonic activity (Olesen et al. 2000, 2004). The deep incision of glaciers produced steep slopes along typical U-shaped valleys (Fig. 2). Etzelmu¨ller et al. (2007) summarized well the resulting relief of western Norway with these words: alpine, steep slopes, heavily over-deepened glacial valleys. The sudden unloading of the steep slopes due to a rapid ice-cap melting associated with a large discharge of water into the slope fractures generated a large amount of gravitational structures in Norway (Blikra et al. 2002, 2006b). Blikra et al. (2006b) invoked frost wedging (i.e. opening of preexisting fractures due to the ice expansion) as a primary factor of slope instability structural development. In the recent past, permafrost covered most of the slopes of western Norway and, during a Younger Drias dramatic cooling event, it even occurred down to sea level (Blikra & Longva 1995) causing ice wedging at the largest possible scale [the lower limit of permafrost today is at c. 1500 m a.s.l. (metres above sea level) in western Norway]. Most of the rock slope instabilities in Norway are believed to have been structured during those two periods and to be presently dormant or inactive (Table 1). The extreme Alpine relief of the coastal fjord areas of Norway combined with the high concentration of pre-existing geological structures, seasonal heavy precipitation, snow melting and frost periods makes these areas prone to catastrophic slope activity (Blikra et al. 2006b). As such, the two counties of Møre og Romsdal and Sogn og Fjordane of western Norway are characterized by a high density of past large rockfalls and avalanches (Anda & Blikra 1998; Blikra et al. 2000, 2002, 2006b; Braathen et al. 2004; Furseth 2006; Bo¨hme et al. 2011), with some 340 historical events recorded over an area of 90 000 km2 and by 72 potential large rock slope failures (Fig. 1 & Table 1). In Sogn og Fjordane County, some hundred historical events are identified and some 20 locations show potential for slope failures (Fig. 3) (see Bo¨hme et al. 2011). In Møre og Romsdal County, three areas were carefully studied. The first two investigated areas are the Romsdalen and Sunndalen valleys because of their concentration of past events. Romsdalen is a 30 km-long, typical U-shaped glacial valley of western Norway (Fig. 4a). Seventeen past large slope failures are recorded along the Romsdalen Valley. Five large –very large instabilities are identified along the southern slope of the valley (Fig. 4a). Along the 30 km-long Sunndalen Valley, 10 past events are identified and five instabilities are currently under examination (Fig. 4b). The third investigated area in Møre og Romsdal County is the 360 km2 Storfjord region (Henderson et al. 2006; Henderson & Saintot 2011) and its 42 potential large rock slope failures (Fig. 5). Field investigation

ROCK SLOPE INSTABILITIES IN NORWAY

29

Fig. 1. Lithological map of western Norway with the location of past and potential large rock slope failures (Geological Survey of Norway database: http://www.ngu.no/kart/skrednett/ and http://www.ngu.no/kart/bg250/; UTM co-ordinates, zone 32N). The frames limit the four areas shown in detail in Figures 3– 5.

30

A. SAINTOT ET AL.

Fig. 2. Maps showing (a) the high relief of western Norway with the present-day contours of uplift and (b) values of slope gradients along the fjords and valleys (derived from a 25-m digital elevation model; UTM co-ordinates, zone 32N).

Table 1. Summary of the field observations made at the 72 investigated sites of western Norway. A black line splits the table according to the lithology at the slope instability (soft v. hard rocks, in the top and bottom parts of the table, respectively) Studied sites (located/ described on)

Lithology

State*

Rock topple, rockfall Rockfall

Dormant?

Strandanipa (Fig. 3) Storfjord site 17b (Figs 5 & 6)

Micaschists

Steiggjeberg (Figs 3 & 28)

Pyroxene-granulites, Rockfall mafic gneiss, syenites, granodiorites and cataclasites Pyroxene-granulites, Rockfall mafic gneiss

Activity

Byttejuv (Fig. 3)

Pyroxene-granulites, Rock topple mafic gneiss

Dormant?

Nærøydalen (Fig. 3)

Mangerite, gabbro, gneiss, anorthosite, syenite Phyllite allochthon over gneiss autochthon Micaschists

Translational rockslide, rock topple, rockfall

Rockfall activity

Translational rockslide

?

Phyllites/gneiss

Several translational rockslides

Kloppuri (Fig. 3)

Vik (Midtfjellet and Hesten) (Fig. 3) Vollan (Figs 4, 18 & 19)

Hjellane (Figs 3 & 15)

Ultramafics

Exfoliation

Foliation

Activity

Activity

Regional fractures (faults, lineaments)

Reactivated east –west contact between gneiss and ultramafics Cataclasic fault zone

Unstable north– south vertical foliation NNE –SSW steep and opened Horizontal sole detachment crossing the slope (anorthosites over mafics) Sliding plane

? Translational (þrotational) rockslide, flexural toppling

?

Fold

Toppling along foliation

Large regional tectonic contacts between squeezed Caledonian allochthonous units and trending parallel to the slope Gneiss oversthrusting phyllites- thrust plane 308 dipping toward the fjord

Maximum No. of volume inherited (Mm3) prominent structures

Easting UTM 32

Northing UTM32

Reference

0

0.05

309472

6836201

Henderson et al. (2008)

1

0.001

398636

6905544

1

0.001

433412

6792621

Selmer-Olsen (1972), Henderson et al. (2008)

1

0.01

430926

6787851

1

0.02

431642

6795996

1

0.5

378839

6749543

Lied (1984, 1987), Henderson et al. (2008) Jørstad (1957), Henderson et al. (2008) Henderson et al. (2008)

1

1

370600

6774600

Henderson et al. (2008)

1

100

505863

6938057

Saintot et al. (2008)

1

100

416513

6811642

Henderson et al. (2008)

31

(Continued)

ROCK SLOPE INSTABILITIES IN NORWAY

Type*

Studied sites (located/ described on)

Lithology

32

Table 1. Continued Type*

State*

Phyllites

Rockslide, complex Active

Viddalen (Figs 3 & 14)

Allochthonous phyllites over autochthonous gneiss

Translational rockslide, rock topple

?

Tussen (Figs 3 & 15)

Phyllites

Translational rockslide

?

Storfjord site 6b (Figs 5 & 29)

Micaschists, garnet-rich micaschists, mica-rich gneiss

Translational rockslide

Dormant

Tirskardskreda (Fig. 3) Katlenova (Fig. 3)

Gneiss

Rockfall

Foliated quartz diorite Conglomerates

Migmatitic gneiss

Foliation

Fold

Regional fractures (faults, lineaments)

No. of Maximum inherited volume prominent (Mm3) structures

Easting UTM 32

Northing UTM32

Reference

Domaas et al. (2000, 2002), Blikra et al. (2000, 2002, 2006b), Braathen et al. (2004), Olesen et al. (2000, 2004), Henderson & Blikra (2008) Domaas et al. (2002), Henderson et al. (2008), Bo¨hme et al. (2011) Henderson et al. (2008), Bo¨hme et al. (2011)

1

200 (900 – 1500)

399394

6747916

NW –SE lineament þ horizontal sole detachment

2

1

405509

6743020

NNE– SSW lineaments as back-crack and structures associated and parallel to sole detachment NE– SW and north–south regional fractures as back-crack and NW – SE fractures as transfer

2

30

429274

6822214

5

15

389618

6887976

Henderson et al. (2006)

Active

0

0.001

329764

6789038

Rockslide

Activity

0

0.04

299673

6789681

Rock topple

Active?

0

0.1

295095

6790289

Rock topple

Active?

0

0.25

406596

6749619

W-Hyllestad, Conglomerates Lifjellet (Fig. 3)

Rockfall

Active?

0

0.25

294234

6790808

Henderson et al. (2008) Henderson et al. (2008) Henderson et al. (2008), Bo¨hme et al. (2011) Lied (1989), Henderson et al. (2008) Henderson et al. (2008), Bo¨hme et al. (2011)

Storfjord site 15 (Fig. 5)

Rockfall

?

1

0.001

398950

6894113

E-Hyllestad (Fig. 3) Terakamben (Fig. 3)

Gneiss

Gentle fold with axis azimuth parallel to slope dip direction.

Sliding plane

Fold axes parallel to sliding vector

Unstable fold hinge (crest)

A. SAINTOT ET AL.

Fla˚m (Fig. 3)

Exfoliation

0.001

398848

6900405

1

0.001

397444

6892294

1

0.001

406004

6898608

Sliding plane

1

0.001

416847

6898988

Sliding plane

1

0.001

418820

6900727

1

0.001

419260

6900300

1

0.005

382731

6785994

1

0.03

312548

6820914

1

0.05

343411

6790016

Sliding planes

1

0.5

422135

6913043

Sliding planes

1

0.5

423030

6913169

1

2

501134

6941198

Henderson & Saintot (2007), Saintot et al. (2008)

1

3

445567

6921062

Henderson & Saintot (2007)

1

7

439875

6920534

1

12.5

341573

6789323

1

40

437833

6924487

Henderson & Saintot (2007) Henriksen (2005), Eiken (2007), Henderson et al. (2008) Braathen et al. (2004), Blikra et al. (2006b)

Gneiss

Rockfall

?

Storfjord site 28 (Fig. 5) Storfjord site 31 (Fig. 5) Storfjord site 41b (Fig. 5) Storfjord site 42 (Fig. 5) Storfjord site 42b (Fig. 5) Hallandshammaren (Fig. 3) Gjøringbøfjellet (Fig. 3)

Gneiss

Rockfall

?

Gneiss

Rockfall

?

Gneiss

Translational rockslide Translational rockslide Rockfall

? ?

Translational rockslide Rock topple

Dormant

Gra˚berget (Fig. 3)

Gneiss

Rockslide, rock topple

Dormant

Storfjord site 39 (Fig. 5) Storfjord site 39b (Fig. 5) Ottem (Fig. 4)

Gneiss

Translational ? rockslide Translational ? rockslide Rockslide, complex ?

Flatmark (Figs 4 & 9)

Gneiss

Rockslide, rock topple

Dormant

Svartinden (Figs 4 & 8) Stopelen (Fig. 3)

Gneiss

Translational rockslide Rockslide

Dormant

Parallel to slope, steep and reopened as back-crack Sliding plane

Børa (Figs 4 & 9)

Gneiss

Gneiss Gneiss Gneiss Gneiss

Gneiss Gneiss

Foliated quartz monzonite

NNE– SSW regional fractures as open cracks Cracks along north–south lineaments Fault cutting the ridge

Active

Dislocation of fold hinges Sliding planes

Dormant?

Active?

Rockslide, complex Active

Folded gneiss

Sliding plane-opened steep slope-parallel exfoliation

Back-crack developed where small-scale vertical folds

Sliding planes

Parallel to slope, steep and reopened as back-crack

Reactivation of gneiss– granite contact parallel to cliff

Henderson et al. (2006)

Henderson et al. (2008) Moxnes (2002), Henderson et al. (2008) Henderson et al. (2008)

33

(Continued)

ROCK SLOPE INSTABILITIES IN NORWAY

1

Storfjord site 16 (Fig. 5)

Studied sites (located/ described on)

Lithology

34

Table 1. Continued Type*

State*

Gneiss

Rockfall

?

Foliated quartz monzonite

Translational rockslide

Dormant

Storfjord site 3 (Fig. 5) Storfjord site 38 (Figs 5 & 8) Storfjord site 1 (Fig. 5) Storfjord site 2 (Figs 5 & 22)

Gneiss

Rockslide

Gneiss Gneiss Gneiss

Easting UTM 32

Northing UTM32

Reference

396937

6884863

Sliding planes

North– south lineaments

2

0.1

383401

6843354

?

Sliding planes

2

0.5

396159

6898227

Translational rockslide Translational rockslide Translational rockslide

?

Sliding planes

2

0.5

419049

6914939

?

Sliding planes

Back-crack along NNW – SSE regional fractures North– south brecciated fault as transfer Back-crack NNE –SSW

2

1

396005

6901730

2

1

396090

6899135

Henderson et al. (2006)

Translational rockslide Translational rockslide Translational rockslide Translational rockslide Translational rockslide

?

Sliding planes

2

1.5

348366

6785839

?

Sliding plane

2

2

406765

6888339

Henderson et al. (2008) Henderson et al. (2006)

?

Sliding plane

North– south transfer

2

2

406224

6888422

?

Sliding plane

NNE– SSW transfer

2

2

400972

6889600

North– south transfer

2

2

403072

6890107

Crack along north–south regional fractures

2

2

409826

6906792

Large regional tectonic contacts between squeezed Caledonian allochthonous units and trending parallel to the slope Transfer zones along north– south brecciated faults

2

5

506485

6936915

Henderson & Saintot (2007), Saintot et al. (2008)

2

5

397952

6869539

Henderson et al. (2008)

NW –SE transfer

2

5

399662

6906560

Henderson et al. (2006)

?

Gneiss

Rockslide

?

Ivasnasen (Figs 4, 18 & 19)

Augen gneiss

Translational rockslide

Active?

Oppiga˚rdshyrna (Fig. 3)

Foliated quartz monzonite

Rockslide, rock topple

Dormant

Storfjord site 18 (Figs 5 & 9)

Gneiss

Translational? Rockslide

Dormant

Gneiss

Maximum No. of volume inherited (Mm3) prominent structures 0.001

Storfjord site 23 (Fig. 5)

Gneiss

Sliding plane

Regional fractures (faults, lineaments)

2

Gneiss

Gneiss

Fold

North– south lineaments

La˚nefjorden (Fig. 3) Storfjord site 11 (Figs 5 & 12) Storfjord site 11b (Figs 5 & 12) Storfjord site 12 (Figs 5 & 7) Storfjord site 12b (Figs 5 & 12)

Gneiss

Foliation

?

Main sliding plane N0408/0508 þ north– south fractures Slope-parallel trending lineaments Localization of back-crack

Localization of back-crack at the hinge Shalow fjord-dipping sliding planes Sliding planes

Localization of cracks along axial fold planes Steep opened plane

Lied (1975), Henderson et al. (2008)

Henderson et al. (2006)

Henderson et al. (2006) Henderson et al. (2006)

A. SAINTOT ET AL.

Storfjord site 8 (Figs 5 & 23) Rustøyane (Fig. 3)

Exfoliation

Gneiss

Rockslide, rock topple

Rockfall activity

Fold axes parallel to sliding vector

Storfjord site 14b (Fig. 5)

Gneiss

Dormant

Sliding plane

Storfjord site 2b (Fig. 5) Storfjord site 21 (Figs 5 & 10)

Gneiss

?

Sliding planes

Contact gneiss/ ultramafic rocks

Translational (þrotational?) Rockslide Translational rockslide Translational rockslide

Oppstadhornet (Oterøya) (Fig. 1)

Gneiss

Rockslide, complex Active

Hegguraksla (Fig. 5)

Banded mica-gneiss Rockslide, rock topple

Active

Sliding planes

Large-scale folds spatially constrain the gravitational deformation

Storfjord site 20 (Figs 5 & 11)

Gneiss

Translational rockslide

?

Vertical cracks

Sliding plane guided by fold geometry

Storfjord site 22 (Figs 5 & 26)

Gneiss

Translational rockslide

?

Opened vertical foliation perpendicular to slope

Storfjord site 26 (Figs 5 & 27)

Contact between two gneiss units

Translational rockslide

?

Sliding plane

Storfjord site 27 (Figs 5 & 27)

Contact between two gneiss units

Translational rockslide

?

Sliding plane

Storfjord site 10 (Figs 5 & 7)

Gneiss

Translational rockslide

?

Dormant

Sliding plane guided by fold geometry Sliding plane

Steep opened foliation as cracks Steep opened foliation as cracks Sliding planes

Large back-crack along north– south regional fractures Back-crack NNE –SSW

2

5

393283

6893897

Henderson et al. (2006)

2

7.5

395731

6891907

Henderson et al. (2006)

Main sliding plane N0308/458 North– south regional fractures as open cracks NW –SE regional fractures as transfer

2

10

395242

6900486

2

15

409275

6903646

Henderson et al. (2006) Henderson et al. (2006)

2

20

388951

6953563

2

10 to 60

415455

6907550

North– south cracks

3

0.1

405511

6904418

Back-crack along vertical north– south regional fault and sliding along west-dipping north – south faults North– south regional fractures as open cracks North– south regional fractures as open cracks North– south regional faults as transfer and NW –SE regional fractures as back-crack

3

1

409319

6906098

3

2

415515

6904360

3

2

416537

6902887

Henderson et al. (2006)

3

2

403639

6888157

Henderson et al. (2006)

Robinson et al. (1997), Blikra et al. (2001, 2002), Braathen et al. (2004), Derron et al. (2005) Furseth (1985, 2006), Harbitz et al. (2003), Braathen et al. (2004), Blikra et al. (2006a), Oppikofer et al. (2008a), Oppikofer & Jaboyedoff (2008) Henderson et al. (2006) Henderson et al. (2006)

35

(Continued)

ROCK SLOPE INSTABILITIES IN NORWAY

Storfjord site 4 (Figs 5 & 10)

36

Table 1. Continued Studied sites (located/ described on)

Lithology

Type*

State*

Exfoliation

Sliding plane

Foliation

Gneiss

Translational rockslide

?

Ha˚rstad (Fig. 4)

Gneiss

Translational rockslide

Dormant

Vertical, reopened, transfer Sliding planes

Storfjord site 10b (Figs 5 & 25)

Gneiss

Translational rockslide

?

Sliding plane

Storfjord site 13 (Fig. 5)

Gneiss

Translational rockslide

?

Sliding planes

Storfjord site 9 (Figs 5 & 24)

Gneiss

Translational rockslide

?

Sliding planes

Mannen (Figs 4, 13 & 21)

Gneiss

Translational rockslide

Active

Parallel to slope, steep and reopened as back-crack

Storfjord site 21b (Fig. 5)

Amphibolites and augen gneiss overthrusting gneiss

Translational rockslide

Dormant?

Sliding planes

Regional fractures (faults, lineaments)

No. of Maximum inherited volume prominent (Mm3) structures

Easting UTM 32

Northing UTM32

Reference

Back-crack along north– south regional faults

3

4

393396

6892440

Back-crack along 508 east-dipping brecciated fault and cracks along NW –SE regional fractures Back-crack along north– south lineamentsNW –SE regional fractures as transfer North– south regional faults as transfer and east –west lineaments as incipient back-crack North– south regional faults as transfer and east –west/NE– SW regional fractures as back-crack North– south faults as transfer zones

3

5

475507

6945271

Henderson & Saintot (2007), Saintot et al. (2008)

3

5

405423

6886351

Henderson et al. (2006)

3

5

394750

6889093

Henderson et al. (2006)

3

7.5

399845

6885400

Henderson et al. (2006)

3

10

436542

6925595

Henderson & Saintot (2007), Dahle et al. (2008)

3

20

409421

6903574

Localization of back-crack following an east– west recumbent fold Fold axial planes North– south þ NE– SW lineaments parallel to slope

A. SAINTOT ET AL.

Storfjord site 5 (Figs 5 & 20)

Fold

Gneiss

Translational rockslide

Dormant

Sliding planes

Storfjord site 7 (Figs 5 & 23)

Gneiss

Translational rockslide

Dormant

Sliding plane

˚ knes (Fig. 5) A

Gneiss

Translational rockslide

Active

Sliding plane

Gikling (Figs 4, 16 & 17)

Gneiss/ fault rocks

Active Translational rockslide (with a slight spreading of the allochthonous above the sole detachment)

Hinge of fold (bulge) limiting two blocks. Back-crack along tight fold hinges Fold axial planes localized part of the back-fracture

NE– SW þ north –south regional fractures as transfer – NW –SE fractures as back-crack NE– SW þ north –south regional fractures as transfer – NW –SE fractures as back-crack North– south regional fractures, large brecciated fault as eastern border /large fault as western border

4

5

393938

6886845

4

12

394063

6886297

5

80

395600

6895500

Sole detachment, large steep faults as limits of instability

5

100

492671

6945891

Henderson et al. (2006)

Blikra et al. (2006a), Venvik Ganerød et al. (2008), Oppikofer et al. (2008b), Jaboyedoff et al. (2011) Henderson & Saintot (2007), Saintot et al. (2008)

*Type of movement in bedrock according to Varnes (1978) and Cruden & Varnes (1996), and state of activity as in Cruden & Varnes (1996). Given measurement of planar structures: strike and dip angle (right-hand rule).

ROCK SLOPE INSTABILITIES IN NORWAY

Storfjord site 7b (Figs 5 & 23)

37

38 A. SAINTOT ET AL.

Fig. 3. The 22 studied slope instabilities in Sogn og Fjordane County with the distribution of historical slope failures reported on a geological map (Geological Survey of Norway database: http://www.ngu.no/kart/skrednett/ and http://www.ngu.no/kart/bg250/; legend of the lithologies as in Fig. 1; UTM co-ordinates, zone 32N).

ROCK SLOPE INSTABILITIES IN NORWAY

39

Fig. 4. (a) The past gravitational failures and the four studied slope instabilities in the Romsdalen Valley and a photograph illustrating the typical U-shape overdeepened glacial Romsdalen Valley (view to the east). (b) The past gravitational failures and the five studied slope instabilities in the Sunndalen Valley. (Background: scale 1/250 000 geological map; UTM co-ordinates, zone 32N: Nilsen & Wolff 1989; Tveten et al. 1998; Geological Survey of Norway Database: http://www.ngu.no/kart/skrednett/ and http://www.ngu.no/kart/bg250/).

40

A. SAINTOT ET AL.

Fig. 5. Map of the Storfjord area showing the historical slope failures and the 40 unstable rock slopes including the two ˚ knes and Hegguraksla (background: scale 1/250 000 geological map on a 25-m permanently monitored rockslides, A digital elevation model; UTM co-ordinates, zone 32N: Tveten et al. 1998; Geological Survey of Norway Database: http://www.ngu.no/kart/skrednett/ and http://www.ngu.no/kart/bg250/).

at the regional scale of Storfjord was decided not only because this area is characterized by the largest number of past events but also because two of the permanently monitored active large rock˚ knes and Hegguraksla rockslides of Norway, A slides, are located there (Blikra et al. 2006a)

(Fig. 5). Extensive fieldwork allowed a detailed examination of the known rock slope instabilities, for new ones to be identified and mapped (Henderson et al. 2006, 2008; Henderson & Saintot 2007; Saintot et al. 2008). It resulted in the total of 72 current slope instabilities. Most of the studied

ROCK SLOPE INSTABILITIES IN NORWAY

slopes show a high susceptibility to develop rockfalls of limited volumes (Derron et al. 2005; Crosta et al. 2007; Derron 2009) not only because of the steepness of the slopes but also because of the high density of interconnected tectonic joints in the rocks. The rockfalls that: (1) may threaten the community; (2) have a significant volume of 1000 m3 and more; and (3) show a high frequency of occurrence are included in the 72 analysed instabilities of this paper. Also, we encountered during our field studies 43 unstable volumes equal to or above 1 Mm3 (see Table 1) prone to evolve to rock avalanches (with a threshold value as in Evans et al. 2006). These are exclusively rockslides (deep-seated rockslides, and translational or rotational rockslides) or display complex slope movements combining rock topples and slides (according to Varnes 1978 and Cruden & Varnes 1996) (Table 1).

Influence of bedrock geology and the reactivation of pre-existing ductile and brittle structures in slope instability development The lithological factor in slope instability development The lithological map (Fig. 1) shows mainly the predominance of hard gneissic rocks in western Norway. They form the autochthonous basement but also most of the allochthons of the Caledonian tectonic units. However, the same map (Fig. 1) also displays some patches of other lithologies that are intrinsically weaker than the surrounding gneisses as amphibolites, schists and micaschists, and weathered mafic and ultramafics. Out of the 72 studied slope instabilities in western Norway, 13 developed in such rocks assumed to be softer than the gneisses. The ultramafics that we observed during the field campaign are strongly weathered and highly fractured compared to the host gneissic rocks. In detail and for example in Storfjord area, the frequency of fractures in the ultramafics is higher than 4 per m and three different orientations of interconnected joints are common (Fig. 6). Such brittle deformation leads to a mechanical weakening and enhances the chemical weathering. Consequently, the ultramafics can no longer be considered as hard rocks. When lying on steep slopes, the patches of ultramafics form very unstable packages that disaggregate and generate rockfalls. The failure surfaces generally coincide with the contact between the gneiss and the ultramafics. For the only Sogn og Fjordane County, 11 of the 20 slope instabilities developed in such soft rocks. Nine of them are specifically concentrated to the east

41

along deep narrow fjords that installed on a micarich phyllitic and mafic–ultra-mafic allochthonous belt (Fig. 3) (see Bo¨hme et al. 2011). This zone is assumed to be of high susceptibility to large gravitational slope deformation. The largest unstable slope of Norway developed along this belt at Fla˚m (location shown in Fig. 3) (Henderson & Blikra 2008). Its maximum volume may reach 200 Mm3. Therefore, the lithology appears to be an important factor controlling the development of unstable masses along the steep sides of the fjords and the failure plane(s) often reactivate(s) the contact between the different lithological units.

The role of pre-existing ductile structures The bedrock of Norway is of particular interest when investigating how the ductile fabrics influence slope stability. From Sogn og Fjordane to Møre og Romsdal counties and further north, to the Møre Trøndelag Fault Complex, the Caledonian ductile strain increases together with the regional metamorphism that ranges from low– middle to ultrahigh grades (with a recorded peak metamorphism of 800 8C at 100 km depth in the eclogites of the coastal area of Møre og Romsdal County: Tveten et al. 1998). The overall metamorphic foliation is rather flat or gently dipping along open folds in the whole Sogn og Fjordane County. Towards the north (i.e. towards the internal high–ultra-highgrade metamorphic core of the Caledonides) and in the Møre og Romsdal County, close –tight folds become more abundant and polyphase folds are common. In the Storfjord area, the ductile deformation of the basement is subsequent to at least three folding phases of the Caledonian Orogeny. At the Møre Trøndelag Fault Complex (Fig. 1), the ductile fabric tends to be parallel to the fault segments, that is, foliation and axial surfaces of close – tight folds are steep –vertical and ENE –WSW trending. It is assumed that the Møre Trøndelag Fault Complex was the locus for a very high shear strain that largely overprinted the older trends of ductile deformation. The following field case studies exemplify how the ductile grain may decrease the slope stability. The role of foliation in slope instability development. Studies of rock slope failures and associated discontinuities in intensively foliated rocks are documented; for example: in British Colombia, Canada, by Evans & Clague (1998, 1999); in the Central Italian Alps by Govi et al. (2002); in the Andes, Argentina, by Hermanns & Strecker (1999); as well as in Norway by Blikra et al. (2000, 2001, 2002, 2006b), Braathen et al. (2004), Henderson et al. (2006, 2008) and Venvik Ganerød et al. (2008) among others. In western Norway the

42

A. SAINTOT ET AL.

Fig. 6. (a) Aerial orthophotograph showing a lens of ultramafics (dunites) prone to rockfalls in the gneissic basement of Storfjord area (site 17b, located in Fig. 5). (b) Photograph of a NW–SE-orientated wall in the ultramafics with the NE– SW joints, gently NW-dipping joints and the steep foliation dividing the rock in blocks. (c) Stereonet of steep structures: joints in black and foliation in grey. The high frequency of joints in the ultramafics contributes to a mechanical weakening and enhances the chemical weathering.

ROCK SLOPE INSTABILITIES IN NORWAY

foliation is a secondary metamorphic foliation with, for gneissic rocks, a differentiated compositional banding between felsic and mafic minerals, and, for phyllitic rocks, cleavage and schistosity from platy mineral planar alignments (e.g. Twiss & Moore 1992; Passchier & Trouw 2005). It is actually the main penetrative structural grain, and the first and inherent factor of rock strength anisotropy to consider when studying rock slope instabilities (Braathen et al. 2004; Blikra et al. 2006b). Several field examples given below focus on the reactivation of foliation planes as gravitational structures on overdip –underdip cataclinal slopes (with anaclinal and cataclinal slopes as defined by Powell 1875: see in Cruden & Hu 1996; Cruden 2000). Foliation reactivated as basal shear planes of rockslides on overdip cataclinal slopes (i.e. with respect to foliation). Thirty of the 72 studied slope instabilities correspond to rockslides with a basal shear plane parallel to foliation on overdip cataclinal slopes (Table 1). Most of them are located within the dioritic –granitic gneiss of the Caledonian autochthonous basement of western Norway. Compositional variation in this host rock is such that some layers are extremely mafic and consist of almost 100% biotite and/or hornblende. We observed that the shear planes at the base of the rockslides are commonly and preferentially developed along the hornblende –biotite-rich layers and not along feldspar – quartz-rich layers. The inherently weak properties of mica and amphiboles (their low shear strength compared to quartz and feldspar) and their platy nature focus strain at the base of the sliding block. Subsequently, gouges and breccias developed along these weak planar structures attesting to shearing along the plane (see also in Henderson & Saintot 2011). Figures 7 and 8 illustrate four localities where rockslides developed with the reactivation of the foliation as basal shear surfaces. In these four cases rockslides occurred in the past and the current slope instabilities are propagating from the previous failure surfaces. Site 10 of the Storfjord area (Fig. 7a–c; location in Fig. 5) displays foliation that dips 208 –358 towards the fjord. A back escarpment produced by the opening of a steep fracture delimits the potential unstable block with an estimated volume of 2 Mm3. Site 10 is one of the potential rockslides lying on the approximately 558 dipping southern face of Geiranger Fjord, along which the foliation dips towards the fjord and is systematically reactivated as basal shear planes (sites 7, 7b, 9, 10 and 10b located in Fig. 5). The same setting is observed at site 12 in Storfjord (Fig. 7d, e; located in Fig. 5) (Henderson et al. 2006) with, in addition, an exfoliation that occurred along the foliation planes and that contributed to

43

their weakening by opening. Such a process is also observed all along the slope in the vicinity of site 38 in Storfjord (Fig. 8a) (Henderson et al. 2006). Exfoliation developed parallel to the 308 dipping foliation and may have facilitated the development of shear planes at the base of the large unstable slices at sites 38, 39 and 39b (located in Fig. 5). Figure 8b, c displays a 7 Mm3 unstable block, called Svartinden, in the Romsdalen Valley (located in Fig. 4a) that slid on a foliation-parallel plane which moderately dips towards the valley (Henderson et al. 2004; Henderson & Saintot 2007). The friction created a 20 cm-thick microbreccia along the sliding plane. This foliationparallel basal sliding plane is also the westwards prolongation of a previous failure surface (Fig. 8b). Propagation of failure surfaces along steep – vertical foliation. Eleven investigated instabilities of large volumes (sites 5, 18, 20, 22, 26 and 27 in Storfjord, see Fig. 5; Mannen, Flatmark, Børa in the Romsdalen Valley, see Fig. 4; Kloppuri and Byttejuv in Sogn og Fjordane, see Fig. 3 & Table 1) are located on slopes where the foliation is steep – vertical, and 10 of them are found on underdip– dip cataclinal slopes (with respect to foliation). Site 18 of Storfjord (Fig. 9a –c) accurately illustrates how the steep –vertical foliation along the cliffs can promote large columnar failures. This site has foliation that steeply dips towards the fjord. Large extensional structures developed parallel to the foliation. A down-dip displacement along these opened fractures resulted in the formation of several metre-high steps on the top of the unstable block. There is, therefore, evidence of a combination of shear and tension on the foliation planes. However, it is not clear what is controlling the sliding of the block at its base (Fig. 9b, c). The southern face of the Romsdalen Valley is characterized by a steep –vertical cliff parallel to foliation. The two large unstable fields of Flatmark and Børa (Fig. 9d, e; located in Fig. 4a) show the opening of foliation under gravitational forces to form large cracks at the back of downwardsdisplaced blocks. At Børa, on the plateau, a palaeoglacial valley trends parallel to the cliff (and therefore to the foliation; Fig. 9e). The ice-river is assumed to have partly contributed to the destabilization of the subglacial bedrock (at the back of the present-day cliff; Fig. 9e) by a partial opening of the foliation (probably by hydraulic fracturing, cf. Boulton & Caban 1995). The final destabilization at the edge of the plateau was coeval with debuttressing subsequent to ice retreat in the Romsdalen Valley. Folds as a factor for slope destabilization. Field observations at 17 localities allow identification of

44

A. SAINTOT ET AL.

Fig. 7. (a) Aerial photograph of a potential rockslide of 2 Mm3 at site 10 in Storfjord (location in Fig. 5) showing the presence of back escarpments (dashed line). A previous slide failed from the same structure (solid line). Additional scarps are marked in dotted lines. (b) Down extension of the back-crack structure (white dashed line) to a foliation parallel sliding plane (black line) (view to east; site 10 in Storfjord). (c) Close view of opened relay-stepped structures subsequent to the shear along the foliation-parallel sliding plane (site 10 in Storfjord). (d) Photograph of 2 Mm3 remaining instability west of a past failure at site 12, 900 m a.s.l., in Storfjord (located in Fig. 5) and (e) stereonets of structures (Schmidt’s projection, lower hemisphere). Exfoliation occurred along the foliation planes, and probably contributed to their loss of strength and to their reactivation as basal sliding planes. Note the open vertical joints acting as ‘back-crack’ type fractures and a conjugate system of joints with an intersection normal to the shear plane. These sets of joints are limited to the block and particularly consistent with newly developed structures related to downwards movement of the unstable block.

the role of fold geometries in the slope destabilization. Besides the intrinsic weakening of the rock volume during folding under ductile conditions (due to the flattening and stretching of minerals, in particular along the limbs), the fold geometries certainly guided the development of new failure surfaces. As mentioned earlier, the interferences of several phases of Caledonian ductile folding is particularly well preserved in inner Møre og Romsdal County where three phases can be distinguished. The Caledonian ductile folding tends to attenuate southwards in Sogn og Fjordane County to lead to gentle open folds. To the north it was totally overprinted by a late high shear strain along the Møre Trøndelag Fault Complex that developed a parallel-to-fault vertical foliation. Fourteen of the 17 sites where we observed the gravitational structures guided by pre-existing folds are therefore in

the inner parts of Møre og Romsdal County and in the Storfjord region, and exemplify the loss of slope stability at a regional scale owing to interferences of several ductile folding phases. Chosen field examples show the influence of centimetre- to kilometre-scale amplitude folds in the development of gravitational structures. At sites 4, 11, 11b, 12b, 20, 21 and 21b in the Storfjord area (located in Fig. 5), basal shear planes developed according to the geometrical constrains of the preexisting folds. The attitude of folds is described according to McClay (1991, table 3.2, p. 49). At site 4 (Fig. 10) the instability consists of at least seven column-shaped blocks with a potential total volume of 2–5 Mm3. These are currently detached from the north –south vertical cliff by a set of north–south deeply eroded regional faults from which extensional cracks propagate. A survey

ROCK SLOPE INSTABILITIES IN NORWAY

45

Fig. 8. (a) Structures observed from the helicopter at site 38, 1000 m a.s.l., in Storfjord (located in Fig. 5). The foliation dips moderately towards the valley along which sliding planes may have developed (dashed line). A weakening of the foliation surfaces occurred by opening when exfoliation superimposed. Along the potential shearing plane a blocky talus can be observed and can be the result of rock disintegration subsequent to movement along the shear plane. There is some evidence for steep, extensional fractures (dotted line) and the presence of an old sliding plane (from Henderson et al. 2006). The volume of the unstable slice is estimated to 0.5 Mm3. (b) Photograph of the remaining 7 Mm3 unstable volume called Svartinden, Romsdalen Valley, at 1300 m a.s.l. (located in Fig. 4a). The previous slide occurred along a foliation-parallel surface that propagated westwards and defined the brecciated basal shear plane of the remaining volume (Henderson et al. 2004; Henderson & Saintot 2007). (c) Same unstable block, Svartinden in Romsdalen Valley, viewed to the east.

by helicopter identified shallow fjord-dipping basal planes at the foot of the columns. A possible contributing factor to the basal shear plane development here is the presence of fjord-plunging inclined small-scale folds in the gneiss (Fig. 10b). These have a trend of approximately 1008N, exactly parallel to the extensional direction of the column failure (i.e. perpendicular to the main set of opened joints: Fig. 10c), and also a plunge of 308 –358ESE (Fig. 10d), which is parallel to the supposed dip of the low-angle sliding planes. The favourably orientated pre-existing ductile fabric in this area therefore promotes a slippage of material at a moderate – shallow angle eastwards towards the fjord. At site 21 in Storfjord (located in Fig. 5: Henderson et al. 2006) the slope displays an impressive rockfall (Fig. 10e), which threatened the small community

of Norddal village some years ago. Examination of the mountainside shows that instabilities still remain present. Two basal surfaces limit an estimated 15 Mm3 unstable volume and define a SW-facing wedge (Fig. 10e, f ). The 458 SE-dipping basal shear plane is subparallel to a lithological contact between mafic rocks at the base and a grey dioritic gneiss above. The gneiss unit displays steep open cracks (Fig. 10f). In addition, the two units display upright 458 plunging microfolds, and the shear plane is perpendicular to the axial planes of the microfolds and contains the fold axes (Fig. 10f, g). The second basal shear plane was not accessible and could only be observed from the helicopter. The rock slope instability at site 20 (Storfjord; located in Fig. 5: Henderson et al. 2006) lies on

46

A. SAINTOT ET AL.

Fig. 9. (a) and (b) Structures observed at site 18 of Storfjord (located in Fig. 5): steep foliation parallel to slope and fractures opened along the foliation and (c) sketch of hypothetical rock failure process (in Henderson et al. 2006). (d) and (e) Photographs showing the opening of the foliation at the edge of the plateau, southern side of the Romsdalen Valley (Fig. 4a). (d) The unstable block collapsing towards the valley at Flatmark along the nearly vertical foliation (Henderson & Saintot 2007). (e) View to the WNW of Børa slope instability (from www.norgei3d.no). Cracks developed parallel to steep foliation planes trending parallel to the cliff of the Romsdalen Valley. The palaeoglacial valley trends parallel to the cliff. The ice loading is assumed to have partly opened the steep foliation and to have therefore contributed to the destabilization of the edge of the plateau. The total destabilization occurred when the ice melted in the Romsdalen Valley.

an approximately 508 west-facing slope and is restricted to the Caledonian allochthonous augen gneiss units above the thrust surface (Fig. 11a). The southern side of the slope instability is observed in a deep gully that is a free boundary with respect to the unstable zone, frequently affected by rockfalls (Fig. 11a). At approximately 250 m altitude along the gully, the steeply SSE dipping foliation is clearly reactivated and promotes large open fractures (Fig. 11b, c). The large NNE –SSW joints are also reopened (Fig. 11b). Further up the gully the gneiss is largely folded and, at the base of what we

define to be the uppermost unstable block, we observed a large 408 dipping shear plane parallel to the axial planes of the small-scale folds (Fig. 11b, d). The instabilities at this site are quite small (0.1 Mm3 maximum) and consist of detached slabs from the southern free border. The limits of the slabs are the steep foliation, or the axial fold planes, and the NNE–SSW fracture sets (Fig. 11b). Sites 11, 11b and 12 in Storfjord (located in Fig. 5) lie on the approximately 608 northern side of Geiranger Fjord, where listric structures dipping towards the fjord are observed (Fig. 12). Some

ROCK SLOPE INSTABILITIES IN NORWAY

47

Fig. 10. Folds controlling the geometry of the basal sliding planes at sites 4 and 21 (Storfjord, located in Fig. 5: see Henderson et al. 2006). (a) Large detached column-shaped blocks at site 4 on a 1-m digital elevation model (UTM co-ordinates, zone 32N). North– south regional faults are deeply eroded and the resulting canyons are free boundaries at the back of the blocks. Cracks opened parallel to the fault zones. (b) Photograph from the helicopter of a basal shear plane (dashed line) at the foot of a block and traces of metre-scale folds (underlined by dotted lines). The inferred dip direction along the shallow fjord dipping basal shear planes is parallel to the fold axes at site 4. (c) Photograph from the helicopter of north– south cracks that detached columns from the north–south cliff and traces of metre-scale folds (dotted lines). The trend of fold axes (arrows) is parallel to the cracks. Stereonet of field data with north– south cracks as black dashed lines, foliation as grey lines and fold axes as dots. (d) Photograph of one of the inclined folds outcropping on the detached columns. (e) Orthophotograph of site 21 (UTM co-ordinates, zone 32N) showing the rockfall deposits and the unstable wedge (limited by the dashed line). (f ) Photograph of the unstable wedge of site 21 (from the helicopter). Stereonet of field data: the NW basal shear plane of the wedge of site 21 (dotted lines) contains the fold axes (dots) and developed perpendicular to the fold surfaces (full lines). (g) Sketch of the geometrical relationships between basal shear planes and inclined or upright gently plunging folds as observed at sites 4 and 21, respectively.

48

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Fig. 11. Unstable slope at site 20 in Storfjord (UTM co-ordinates, zone 32N; located in Fig. 5: see Henderson et al. 2006). (a) The gravitational deformation is restricted to the overlying Caledonian allochthonous augen gneiss unit. The present-day activity is along the free boundary on the northern side of a gully. Traces of foliation underlined with dotted lines: foliation is undulating and steep in the lower part of the slope and displays gently inclined folds in the upper part of the slope (b) Stereonets of structures measured on the field along the northern active edge of the gully. (c) Photograph of the structures that detach large slabs from the wall: NNE–SSW open fractures and east–west steep foliation surfaces. The east–west steep surface also displays the horizontal hinges of the undulating steep foliation. (d) Photograph of a sliding plane (dashed line) localized parallel to the inclined axial surfaces of the small-scale horizontal folds.

rock slope failure activity occurred in the past along such listric planes (Fig. 12a–c). An important feature at these three sites is the role of the preexisting horizontal fjord-inclined fold that guided the formation of the single listric faults. These

faults are parallel to the inclined axial fold surface and act as sliding planes at the base of the instabilities. They correspond to back-cracks where they steepen up towards the surface (Fig. 12a–c: Henderson et al. 2006). At site 12b the listric fault has an

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Fig. 12. (a) – (c) Observations made at sites 11, 11b and 12b in Storfjord (located in Fig. 5; see Henderson et al. 2006). The listric failure planes on this slope are guided by the geometry of horizontal inclined folds. (a) Site 11b is at the position of a previous slide (picture from the helicopter). Traces of foliation as dotted white lines displaying the hinge of an inclined horizontal fold. (b) The low-angle basal shear plane is parallel to the inclined axial surface and steepens up into an extensional back-crack structure (picture of site 11 from the helicopter). (c) Structures observed at site 12b. The basal shear plane has a listric geometry dipping 308 –358 towards the fjord and parallel to the foliation and steepening up into an extensional fracture along the axial plane of an antiform. A block already detached from this structure and was laterally guided by a north–south-trending old fault zone that acted as a transfer fault. (d) Horizontal inclined folds controlling the geometry of the basal sliding planes. At five sites in the Storfjord area (sites 11, 11b, 12b, 20 and 21b located in Fig. 5) the sliding plane at the base of the rock slope instability is developed parallel to the moderately inclined axial plane of horizontal folds. The dip-slip direction of blocks is inferred to be perpendicular to the fold hinges.

angle of 308 to 358 towards the fjord, parallel to the axial plane of an antiform, and steepens up as it reaches the surface to form an extensional fracture truly cutting the hinge of the antiform. The eastern limit of the failed block at site 12b is a north– south epidote-rich fault plane that acted as a transfer fault (i.e. the lateral boundary between the unstable and stable parts of the slope: Fig. 12c). The basal shear planes of rock slope instabilities at

sites 11, 11b, 12b, 20 and 21b in Storfjord (located in Fig. 5: Henderson et al. 2006) developed parallel to axial surfaces of horizontal folds moderately inclined towards the fjord or valley and the inferred slip direction of the instabilities is perpendicular to the fold axes (Fig. 12d). At the active rockslide of Mannen (Romsdalen Valley, located in Fig. 4: Henderson & Saintot 2007; Dahle et al. 2008), the back-crack of the

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collapsing block is localized on a folded zone and opens perpendicularly to the axial surfaces of recumbent horizontal folds (Fig. 13). This site is a good example of the weakening of the rock mass due to folding. The folds promote anisotropy in terms of strength that will be the locus for the development of gravitational failure surfaces. The field examples above demonstrate that the presence of centimetre- to hectometre-scale close – tight ductile folds influence the stability of the slopes. Folded zones can focus the gravitational deformation (as localized folded narrow zones). The contact between folded and unfolded units can be the locus of gravitational structures (due to a contrast of rheology between the units as a result of folding). Development of sliding planes at the base of rockslides can be guided by the geometry of centimetre-scale folds (where favourably orientated with respect to the slope), either parallel to the fold axes or parallel to the axial surfaces. Where decametre- or hectometre-scale fjordinclined fold limbs intersect the slope, they are reactivated as basal sliding planes and the upwards development of the failure planes (back-cracks)

is often deviated from the fold hinges. The latter confirms that the ductile folding decreases the strength of the rocks in the limbs and not in the hinges of folds (see the discussion in the last section of this paper). Influence of the sole detachments of the Caledonian allochthons on the slope stability. Western Norway is characterized by the remnants of the allochthonous sheets of the Caledonian Orogeny (Roberts & Gee 1985; Hossack & Cooper 1986; Roberts 2003). Most of the contacts between the different units (displayed in the map in Fig. 1) are tectonic in origin. The sole detachments of the allochthons developed in the ductile domain of deformation and they are commonly rather flat or shallowdipping. However, in a few places some of these contacts are found to be vertical. Field studies allowed the identification of eight rock slope instabilities that are located in the vicinity of the sole detachments. Such a setting provides an opportunity to evaluate how the presence of sole detachments may contribute to the gravitational destabilization of the slopes.

Fig. 13. The back-crack of the collapsing block opens perpendicularly to the axial surfaces of recumbent horizontal folds (white dashed line) at Mannen (Romsdalen Valley, located in Fig. 4a; see Henderson & Saintot 2007).

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At the sites called Vik, Viddalen, Tussen and Hjellane in Sogn og Fjordane County (located in Fig. 3), Gikling in the Sunndalen Valley (located in Fig. 4b) and site 20 in Storfjord (located in Fig. 5; also illustrated in Fig. 11) the gravitational deformation is limited to the Caledonian allochthonous units above the flat-lying–shallow-dipping sole detachments that intersect the steep slopes. Vik, Viddalen, Tussen and Hjellane slope instabilities (in Sogn og Fjordane County, located in Fig. 3) developed in the phyllitic nappes. As mentioned earlier, steep phyllitic slopes are highly susceptible to rupture only because of the low strength of the rocks. However, and in addition, at the four sites of Vik, Viddalen, Tussen and Hjellane there is the sole detachment underneath the gravitational structures. At Viddalen the instability developed on an anaclinal slope with foliation of the phyllites at about 308. Despite such geometry that may enhance the stability, a collapsing zone of approximately 1 Mm3 developed (Fig. 14: see also in Bo¨hme et al. 2011). The lowermost limit of the instable volume coincides with the 58 south-dipping sole detachment cropping out along the cliff (Fig. 14a, b). The qualitative assessment to explain the development of Viddalen instability against the unfavourable geometry of the foliation is to consider: (1) the intrinsic weakness of the rocks; (2) its pre-existing vertical joint sets (see in Bo¨hme et al. 2011); and (3) the presence of the sole detachment and its hydraulic properties. Because the sole detachment is underlined by water springs, it may act as a conduit for groundwater; the subjacent gneiss being, in this case, quasi-impermeable. Considering the heavy rainfall and snow melting periods, large water pressures, consequent to the rising of the hydraulic head, may build up along the sole detachment, contributing to local destabilization and additional opening of pre-existing joints (Terzaghi 1962; Giraud et al. 1990; Crosta & Agliardi 2003; Cruden 2003; Sartori et al. 2003; Grimstad 2005 among others). Also during wintertime, when the outflow of water at the springs gets blocked by frost, the hydraulic head may lift up yielding large water pressures at the sole detachment. The fact that the water springs never dry from spring to autumn may indicate that the water flows from far inwards in the high mountain plateau (Fig. 14c) and an approximately 400 m-high hydrostatic column may then be considered to cause pressures of some units of MPa at the sole detachment. Such pressures would exceed the low tensile strength of the jointed phyllites (the anisotropic tensile strength of intact phyllites ranging from 1.25 to 15 MPa: Hanssen et al. 1990; Ramamurthy et al. 1993; Saroglou & Tsiambaos 2008) and may subsequently generate (1) further opening of the tectonic joints

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and (2), by decreasing the normal stress acting on the sole detachment, a spreading-like displacement of the overlying unit specifically directed towards the free cliff surface (Fig. 14c). Tussen is a site in Sogn og Fjordane (located in Fig. 3) that displays two rock slope instabilities (Fig. 15a). The southernmost instability was the focus of the field study and consists of two blocks (namely blocks A and B; Fig. 15b: Henderson et al. 2008; Bo¨hme et al. 2011). Both of the instabilities are located in the allochthonous phyllitic sheet thrusting over the autochthonous gneiss along a shallow south-dipping sole detachment (Fig. 15c, d). The cleavage in the phyllites is parallel to the sole detachment (see in Bo¨hme et al. 2011). It is likely that some shears occurred along the foliation coeval with the large shear along the nearby sole structure, rendering the south-dipping foliation more susceptible to being sheared again under gravitational forces. It would imply a southwardsdirected sliding of the blocks, oblique relatively to the NNE– SSW-striking cliff. The opening of large back-bounding crevasses also clearly took place along slope-parallel segments of kilometre-scale NNE –SSW lineaments (Fig. 15a, b). Six different potential rockslides are reported at Hjellane in Sogn og Fjordane (Fig. 15e, f; located in Fig. 3). Five of them lie on the thin gneissic allochthonous cover, overlying a phyllitic allochthon. The sole detachment of the allochthonous gneiss is dipping approximately 258 towards the fjord. Such a position favours the gravitational sliding of the gneiss along the reactivated sole detachment (Fig. 15e, f ). One potential rockslide developed in the weak phyllitic rocks at the knick-point of the slope (see Fig. 8c in Bo¨hme et al. 2011). The site called Gikling in the Sunndalen Valley (located in Fig. 4b) (Henderson & Saintot 2007; Saintot et al. 2008) is a very large active and complex unstable zone that covers an area of approximately 1 km2. It lies on a 258 slope, the uppermost limit of the deformed area being at 1400 m a.s.l. and the inferred toe zone below 900 m (Figs 16 & 17). Our field study at this site emphasized the important role of pre-existing structures in the development of slope instabilities. Not only were the pre-existing structures clearly reactivated by gravitational forces but they also defined the extent of the deformed volume. The foliation is, in general, dipping to the south (Fig. 16). The uppermost limit of the unstable zone corresponds to a spatial arrangement of pre-existing structures (Fig. 17). As such, segments A and B in Figure 16 are reactivated old brecciated faults (Faults A and B in Fig. 17a–d). Segment C (Fig. 16) comprises a north– south crack several metres wide and deep, that displays stepped walls. The steps are made up of the planes of the south-dipping foliation and of

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Fig. 14. (a) Photograph of the Viddalen slope instability (located in Fig. 3) that developed at the edge of the cliff in the phyllitic Caledonian allochthon (see also Bo¨hme et al. 2011). The lowermost limit of the instability coincides with the sole detachment between the allochthon and the autochthonous gneiss. The tectonic contact is underlined by water springs. (b) Topographic map of the surrounding of the slope instability with hydrographical system and tectonic units (UTM co-ordinates, zone 32N). (c) Topographic profile (located in b) with the possible path of groundwater (grey arrows) from subsurface to the sole detachment (dotted black lines represent the foliation). The instability may have partly developed by high water pressures along the flat-lying sole detachment, when outflow of water gets blocked by the frost in wintertime or when high rates of precipitation occur (see text for further explanation).

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Fig. 15. (a) Three-dimensional view (www.norgei3d.no) of the two instabilities at Tussen (Sogn og Fjordane County; located in Fig. 3: see Henderson et al. 2008; Bo¨hme et al. 2011). (b) Three-dimensional view (www.norgei3d.no) of the southern instability. Large back-bounding crevasses follow NNE– SSW regional lineaments. (c) Aerial photograph with tectonic units and (d) geological section across the unstable slope. The instability developed in the phyllitic Caledonian allochthonous nappe. (e) Locations of the six rock slope instabilities at Hjellane (Sogn og Fjordane County; located in Fig. 3: see Henderson et al. 2008) on the geological map (UTM co-ordinates, zone 32N) and (f ) profile. Five of the six potential rockslides developed in the thin Caledonian allochthonous gneissic sheet and the main basal shear plane is probably the reactivated 258 fjord-dipping sole detachment. One of the potential rockslides developed in the underlying allochthonous phyllites at the knick-point of the slope.

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Fig. 16. Detailed structural mapping on the geological map (UTM co-ordinates, zone 32N: Tveten et al. 1998) and cross-section showing the peculiar tectonic setting of the large complex slope instability of Gikling (Sunndalen Valley, located in Fig. 4b: see Saintot et al. 2008). Letters A –D refer to specific segments that limit the gravitational deformation, and A–C to pre-existing structures displayed in photographs in Figure 17.

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Fig. 17. Photographs illustrating the pre-existing structures that limit the complex slope instability at Gikling (Sunndalen Valley, located in Fig. 4b: Saintot et al. 2008). (a) Picture to the NW from the helicopter. Faults A and B, respectively, refer to the segments A and B in Figure 16 that limit the unstable slope. (b) Reactivated pre-existing brecciated NE-dipping fault (fault A in picture a, and segment A in Fig. 16) that marks a sharp border between the unstable and stable slopes. (c) Reactivated pre-existing steep SE-dipping fault with gouge and breccias (fault B in picture (a) and segment B in Fig. 16). (d) Close-up of the fault B with fine-grained breccia and gouge (white frame in c; view to SW). (e) Localization of tensional deformation to form segment D (as referred to in Fig. 16) along a zone of small amplitude gently plunging inclined folds, the newly formed gravitational fractures at this place trend parallel to the azimuth of fold axes (view to the west). (f) Picture to the west from the helicopter of the gravitational graben-like structures on the uppermost unstable sub-block (see the location in a).

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the north –south-trending steep fractures. Segment D (on Fig. 16) is the uppermost limit that we observed. It is a wide zone of newly formed tensional fractures that are localized along a zone of small-amplitude gently plunging inclined folds and whose planes are oriented parallel to the fold axes (Fig. 17e). The lowest limit of the instability is not reachable and therefore not well constrained by direct field observation. However, a semi-horizontal line of springs is observed at about 800 m a.s.l. (Fig. 16). It coincides with the shallow-dipping sole of the gneissic allochthonous unit, in which are confined the gravitational structures (Fig. 16). The sole of the allochthon is composed of metasandstones and schists that are weaker that the overlying gneiss, and that would intuitively not support such a material where cutting across the slope (considered as a free border that will allow the propagation of the gravitational deformation). This level is a good candidate to be the basal limit of the slope instability and, therefore, the volume of the entire slope instability

would be greater than 100 Mm3. In addition, a large vertical fault zone lies at the eastern foot of the unstable zone. It separates the two autochthonous units of the valley and also limits to the east the allochthonous sheet (Fig. 16). Gikling is, therefore, in a very peculiar structural setting that makes the slope vulnerable to gravitational forces. At the sites called Vollan and Ivasnasen, in the Sunndalen Valley (location given in Fig. 4b), the Caledonian contacts between the five squeezed tectonic units are steep to vertical and NE–SW trending (Fig. 18). The potential rockslide identified at Ivasnasen (located in Fig. 4b) is underlain by allochthonous augen gneiss (Fig. 18). The site displays an old slide with a maximum failed volume of 5 Mm3 (Figs 18 & 19a, b). The top of the failed volume was at approximately 500 m a.s.l. and the bottom of the valley is at 200 m a.s.l.. Extensive fractures appear to propagate from the previous failure plane towards the SW and towards the NE, isolating

Fig. 18. Geological map (UTM co-ordinates, zone 32N: Nilsen & Wolff 1989) of Vollan and Ivasnasen surroundings (Sunndalen Valley, located in Fig. 4b: Saintot et al. 2008). Note that sector 1 is a topographic promontory called Ivasnasen, with three free borders that may favour the destabilization of the whole volume.

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Fig. 19. (a) Three-dimensional view (www.norgei3d.no) of the slope instability at Ivasnasen (Sunndalen Valley, located in Fig. 4b) and locations of the different sectors of the slope instability (Saintot et al. 2008). (b) Northeastern part of the back-fracture of the previous failed volume and slope instability that propagates towards the NE (sector 2). (c) Southwestern slope instability (sector 1) and open NE–SW fractures that propagated from the failed back-crack. (d) Aerial photograph (www.norgeibilder.no) of Vollan site in the Sunndalen Valley (located in Fig. 4b) with limits of the slope instability (dashed line: Saintot et al. 2008). The arcuate shape of the cracks that developed within the instability is a typical deformation pattern in weak rocks. (e) The back-scarp of the rock slope instability is a reactivated segment of a regional NE–SW vertical tectonic contact between the phyllites (in which developed the instability) and the quartzites (with a foliation parallel to the contact).

the two new unstable zones (sectors 1 and 2; Figs 18 & 19). The southwestern instability is coincident with a topographic promontory (sector 1; Figs 18 & 19a, c). The three free borders of the topographic promontory may favour the destabilization of the whole body. The limiting back-scarp and the parallel fractures downwards (Fig. 19) strike NE–SW parallel to the steep thrust planes displayed on the geological map (Fig. 18). The estimated volume of sector 1 may be up to 5 Mm3. The large slope instability of Vollan is located in phyllites and micaschists (Fig. 18). Figure 19a shows an aerial photograph with the 1 km2 estimated extent of this large slope instability. Given the arcuate shape of the cracks that developed within the instability, it appears that most of the displacement was accommodated by creeping as a

‘semi-ductile downwards flow’, that is a typical deformation pattern in weak rocks. The opening of the back-scarp (at 1050 m a.s.l.) occurred by reactivation of a segment of the NE– SW-trending steep regional tectonic contact between quartzites and micaschists (Fig. 19b). The lowermost level of the deformed area may also correspond to the tectonic contact between the allochthonous micaschists and the quartzites (Fig. 18) at 400 m a.s.l. These field examples enhance the need to systematically check the slope stability where the allochthonous sole detachments crop out. When dipping towards the valley or fjord, they are reactivated as sliding surfaces. They can be also composed of a de´collement layer of weak rocks (e.g. sandstones, micaschists) that may not bear the overlying allochthon rocks, even if their dip angle and

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direction are not optimum (i.e. flat-lying). They are also often underlined by water springs and act as water conduits. Seasonally, when (1) excessive precipitation occurs and snow melts and when (2) in wintertime the outflow of water is stopped by frost at the springs, the hydraulic head may significantly rise, and hydrostatic pressures of some units of MPa can build up along the plane. Such pressures may be large enough to promote: (i) further opening of the pre-existing fractures in the overlying rocks; and (ii) the free slope surface-directed gliding of the overlying units by decreasing the resistance to shear along the plane.

Reactivation of brittle structures in rockslide development It is understood that the susceptibility of a given slope to rockfalls or rockslides increased with its density of brittle structures. Of note herein is our definition of ‘density of structures’ as a qualitative estimation of the concentration of structures within an area in relation to its size, which includes both the frequency and persistence of structures. As mentioned before, after the ductile Caledonian Orogeny the bedrock of Norway was deformed by a significant amount of tectonic events and a high density of brittle structures (showing, therefore, a high frequency and persistence) are observable at all scales. In the following, we demonstrate through selected field examples how the presence of these brittle structures affects the slope stability. Reactivation of exfoliation. The most recent development of brittle structures in Norway is a widespread exfoliation process caused by the sudden ice-cap removal. However, the unstable volumes along exfoliated slopes are generally too small to be classified with the 72 studied instabilities. The shape and the size of the instabilities are, indeed, constrained by the small spacing of exfoliation joints and, therefore, the instabilities mostly correspond to relatively thin slices of rocks. The most prominent exfoliation we observed during our field campaigns is found in Tafjord and cuts across the metamorphic foliation (for example, between sites 26 and 27 located in Fig. 5). The exfoliation process also occurred along the parallel-to-slope, or nearly, metamorphic foliation and contributed to its weakening through opening (as along the western side of La˚nefjorden in Sogn og Fjordane County; Fig. 3). We also observed that at the scale of western Norway there are few remnants of slopes that display exfoliation joints. This may be due to the weak downwards penetration of the exfoliation and the subsequent superficial rockslides having already failed. At site 5 of Storfjord (location shown in Fig. 5) there is an example of a 5 Mm3

unstable block that slid along an exfoliated plane. The block is also limited by one of the north – south regional faults and by an east –west crack. The latter follows the vertical foliation (Fig. 20). The role of regional faults in slope instability development. The regional faults of western Norway cover a large spectrum in strike and dip (Fig. 1). However, some prominent strikes are identified: north– south; ENE– WSW to NE –SW; and WNW –ESE to NW –SE (cf. in Gabrielsen et al. 2002). These large structures largely controlled the hydrographic pattern and fjord system. As such, on the fjord-sides, we can generally observe an important set of mesoscale brittle structures (and among them, striated faults) trending parallel to the fjord axes and believed to form the damage zone of these regional faults. The ENE –WSW strike is the prominent brittle grain towards, and of, the Møre Trøndelag Fault Complex. The north –south steep faults design the coast westwards of Sogn og Fjordane County. North– south regional faults were also observed in the Storfjord area and the Romsdalen Valley, and played a role in the development of slope instabilities. When cutting the gneissic rocks, the faults are large topographic linear depressions. This differential erosion suggests that they are weaker than the host gneisses. They, indeed, often present a metre-scale core filled by a juxtaposition of cohesive and non-cohesive fault rocks; as, for example, epidote-rich cataclasites, and hematiteand chlorite-rich breccias and gouges. Some field examples described herein illustrate the reactivation or the control of such structures in slope instability development. The role of regional-scale lineament orientations and their angular intersection and subsequent interaction with the fjord geometries is further discussed by Henderson & Saintot (2011). At the site of Mannen (in the Romsdalen Valley, location shown in Fig. 4a) we observed that the western border of the instability is a nearly north – south-trending vertical epidote-rich brecciated fault, largely opened under gravitational forces. It also marks a clear limit between the unstable eastern zone, where numerous secondary parallel pre-existing fractures are opened, and the more stable western zone, where such structural grain is less developed (Fig. 21a). The active collapsing block of Mannen is highly disrupted by the same system of opened fractures (Fig. 21b). The perpendicular fractures opened along the east –west steep foliation (Fig. 21c) and the main back-crack is localized at a folded zone (see Fig. 13). Several sets of moderately- to shallow-dipping joints were also measured and some of them are good candidates to define basal sliding planes (Fig. 21c). The regional system of north–south-trending vertical faults is very well developed in the SW

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Fig. 20. (a) Photograph showing the intense network of north–south regional faults with one of them reactivated as a back-crack of the potential rockslide at site 5 of Storfjord (located in Fig. 5). (b), Photograph showing the unstable block and its limits as reactivated structures. (c) Stereonet of structures measured at the foot of the block: foliation planes as black lines, exfoliation surfaces as grey lines, large regional lineaments as dashed lines. (d) Close view at the foot of the unstable zone, there is a recent rockfall activity at the front of the unstable block, that is, precisely above the basal shear plane (exfoliation and foliation planes as dashed white and dotted black lines, respectively).

part of the studied Storfjord area (Fig. 5). Along the north–south-trending fjords either the faults are largely reopened as back-cracks where they are vertical, as observed at sites 1, 4, 5 (see in Fig. 20a) and 6b (location shown in Fig. 5) (Henderson et al. 2006), or reactivated as basal shear planes where

they are favourably dipping towards the fjord, as observed at sites 2 and 2b (location shown in Fig. 5) (Henderson et al. 2006). Sites 2 and 2b in Storfjord are the locations of two potential rockslides that developed on an approximately 658 dipping east-facing slope with

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Fig. 21. (a) Photograph from the helicopter of the opened roughly north–south-trending epidote-rich faults and parallel or nearly parallel fractures (black dashed lines) on the unstable edge of the plateau at Mannen (Romsdalen Valley, located in Fig. 4a: Henderson & Saintot 2007; Dahle et al. 2008). The opening of the system occurred under gravitational forces. A north– south main fault marks the limit between the unstable and stable areas. The foliation is steep and marked with a white dotted line (b) Photograph from the helicopter of the collapsing block of Mannen with an important set of north–south trending opened fractures. The dotted line is the unstable block limit in this plane view. (c) Stereonet of field data.

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reactivation as a basal sliding plane of the same system of regional faults dipping 408 towards the fjord. These faults are subparallel to, or slightly steeper than, foliation (their development is assumed to have been partly guided by the preexisting metamorphic foliation). Site 2 corresponds to an approximately 80 m-high block. The unstable block is also detached from the slope by the roughly north–south steep regional fractures that are largely reopened and act as back-crack structures. The basal sliding plane is identified along a reactivated 50 cm-thick epidote-rich cataclastic fault zone, with the main sliding surface exactly developed at the interface between the cohesive fault rocks and the host gneiss (Fig. 22). The instability at site 2b

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is located approximately 900 m above sea level on a parallel epidote-rich cataclastic fault that reaches the surface inwards the plateau. This is the whole prominent ridge with an estimated volume of 10 Mm3 that is involved in the gravitational sliding. Examination of the basal shear plane showed that a fine-grained breccia was created by the sliding of the block and superimposed on the pre-existing fault core. Along the east –west-orientated fjord system the north –south regional steep faults behave as transfer fractures that allow the sliding of blocks from their sides (typically at sites 7, 7b and 9, see their locations in Fig. 5; see also Henderson & Saintot 2011). A detailed mapping from west of sites 7 and 7b to

Fig. 22. (a) Basal sliding plane at the foot of an unstable block, site 2 in Storfjord (located in Fig. 5). (b) The basal sliding plane is well developed and reactivated an epidote- rich cataclastic fault. (c) Closer view of a lens of epidote-rich cataclasites in the fault core. The planes of weakness are at the interface between the cohesive cataclasites and the host gneiss.

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site 8 of Storfjord (Fig. 23) showed the importance of these north–south large cataclastic fault zones in assessing the slope stability. The offset of bedrocks along the faults juxtaposed gneiss with foliation dipping towards the fjord and gneiss with foliation dipping into the mountain; in other words, the faults delimit unstable cataclinal and stable anaclinal parts of the slope. Sites 7 and 7b are on the cataclinal part of the slope and are surrounded by large areas where failures occurred along the fjorddipping foliation reactivated as sliding planes. Large WNW –ESE vertical cracks opened in these two remaining instabilities (Fig. 23), they are subparallel to the second prominent set of regional lineaments of the large Storfjord area that partly controlled trends of the fjord (Fig. 1). A fine-grained breccia formed because of the block sliding along a foliation-parallel basal shear plane at site 7b (Henderson et al. 2006; Henderson & Saintot 2011). This susceptible part of the slope is limited westwards and eastwards by two north– south cataclastic large faults, west and east of which the slopes are anaclinal (Fig. 23). This study along such a slope showed again the high relevance of the foliation as a parameter to define the slope susceptibility. It showed also, however, that other structures should be taken into account when carrying out a slope stability analysis. As such, and near site 8, one of these faults juxtaposed a western anaclinal slope and an eastern cataclinal slope (Fig. 23). Whereas it is thus predicted that the western zone would be more stable, unsteady slabs formed because of the occurrence of approximately 558 dipping slope-parallel exfoliation surfaces reactivated under gravity. At site 9 on the 558 dipping southern side of Geiranger Fjord (located in Fig. 5) the foliation dips towards the fjord and ENE– WSW cracks are widely opened in the potential sliding mass (Fig. 24a). Fieldwork demonstrated that these ENE– WSW structures consist of large (several kilometres long and several metres wide) preexisting fault zones that were reactivated by recent extensional and down-dip movement towards the fjord (Fig. 24b –d). No low-angle shearing plane at the base of the unstable block was evidenced in the field owing to the inaccessibility of the area and the severity of the topography. However, it is postulated that the extensional displacement on the steep structures is taken up on shearing along some basal structures, parallel to foliation, and that the down-dip displacement may be coupled with a rotation of the blocks. The sliding mass is guided by north–south pre-existing faults parallel to the motion of the block (Fig. 24e). Site 9 is a good example of how a combination of two sets of regional faults and of a fjord-dipping foliation increases the susceptibility of a steep slope to rockslides (Fig. 24e).

Some field examples in Storfjord also clearly show that these faults, where oblique to the slope, guided the structures that developed under gravitational forces. However, such instabilities formed because of the presence of other pre-existing zones of weakness favourably orientated relatively to the fjord or valley. This is the case at sites 10b, 22, 26 and 27 in Storfjord (location shown in Fig. 5) presented below, where a combination of pre-existing structures favoured rockslide formation. The unstable block at site 10b may be of 2 Mm3 maximum and lies on a 508 dipping SW slope of Geiranger Fjord (Fig. 25). Low-angle basal shear planes are developed subparallel to the foliation and recent movement along the shear structures is evidenced by the development of a fault-rock gouge. The block is fully delimited by: (1) an extensional vertical crack, oblique to the slope, which is a reactivated long-lived regional north–south lineament; and (2) a NE–SW steep large structure that transfers the downwards motion of the block (Fig. 25). The NE –SW steep structures strikes parallel to a system of small-scale inherited epidoteand chlorite-coated vertical faults that we observed at the toe zone of the block (Fig. 25d). One of the best examples of a combination of structural sets was observed at site 22 in Storfjord on a 708 west-dipping slope (Fig. 26; location shown in Fig. 5) (Henderson et al. 2006). Open NE– SW vertical fractures, together with the opening of the NW–SE vertical foliation, form a columnar fabric into the unstable part of the slope. The southern part of the main back-crack follows a NE –SW fracture. Two different types of faults, both characterized by epidote-rich coated surfaces also contribute to the development of an unstable rock mass. Gently west-dipping faults intersect the columns and may form the basal sliding plane required to detach the rock mass at the foot of the columns. The pre-existing epidote-rich vertical north– south faults guided the development of the northern part of the main crack at the back of the unstable area. The main mechanism here is likely to be columnar failure, thereby reducing the likelihood of a large volume of unstable material. The slope instabilities of sites 26 and 27 in Storfjord (location shown in Fig. 5) are found where two gneiss units are juxtaposed along a NW–SE steep contact cutting the 558 –608 dipping parallel slope (Fig. 27). In addition, several steep north–southtrending regional-scale faults obliquely cut the NW– SE-striking slope (Fig. 27a). Prominent fractures trend parallel to these large main structures and are largely opened. The east –west-trending vertical foliation planes guided the development of the second trend of open fractures (Fig. 27c). These two nearly perpendicular vertical trends contributed to detaching large high blocks from the slope. Field

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Fig. 23. Detailed topographic map (UTM co-ordinates, zone 32N) from sites 7 and 7b to site 8 in Storfjord (located in Fig. 5): north–south regional cataclastic fault zones acting as transfer zones and limiting unstable and stable parts of the slope with regard to the attitude of foliation from one side to another side of the faults. At site 8 the foliation dips into the mountain slope but the exfoliation dips towards the fjord, and the latter was the locus for the development of gravitational structures. The data of stereonets correspond to sites 7 and 7b: the observed basal shear planes of the potential rockslides are parallel to the foliation planes; the roughly north–south set of steep fractures corresponds to the trend of transfer faults and the WNW– ESE set to the trend of back-cracks.

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Fig. 24. (a) Structures of the rockslides at site 9 in Storfjord (located in Fig. 5): ENE– WSW extensional fractures steeply dipping towards the fjord and postulated shearing along the undulating shallow– moderate fjord-dipping foliation to accommodate the extensional movement in the steep structures. A down-dip component of displacement is also inferred along the same steep ENE –WSW fractures and may be accommodated by a rotational component of motion of the blocks. (b) Steep pre-existing 3 m-wide ENE –WSW-trending fault zone that consists of a ‘damage zone’ of many smaller fractures. These are partly cemented by quartz. This weak zone focused recent movement, producing three opened fractures. (c) and (d) Secondary structures that evidence both extensional and down-dip (fjord-directed) movement along the steep ENE–WSW fractures. (e) Stereonets of the three pre-existing structures that are reactivated under gravity and led to a completely structured rockslide.

observations at the toe zones of the instabilities documented sliding planes that trend parallel to well-developed 308 –408 fjord-dipping exfoliation joints. Steiggjeberg in Sogn og Fjordane County (location shown in Fig. 3) displays a vertical face with frequent rockfall activity (Fig. 28a). The site is located in the footwall and, more specifically, in the damage zone of a moderately fjord-dipping Devonian regional normal detachment (Fig. 28b). The erosional process deeply incised the slope

down to the thick cataclasites of the fault core and formed steep walls within the overlying granulites. A dense network of faults (Fig. 28c) that characterize the wide damage zone of the detachment dramatically increases the instability of the wall. At Steiggjeberg, the pre-existing tectonic architecture, the subsequent rheological contrast and relief favour the destabilization of the rock under gravitational forces. Our observations at 45 sites lead to the conclusion that whatever their orientation relative to

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Fig. 25. (a) Photograph from the helicopter to the south of site 10b in Storfjord (located in Fig. 5) and of the large north– south regional fault zone. (b) Orthophotograph of the site showing the triangular shape of the site limited by the large north–south and NE– SW pre-existing faults (UTM co-ordinates, zone 32N). The white and black arrows highlight, respectively, the north– south and the NE–SW reactivated faults. (c) Closer view of the site, to the SW. The dashed line is the basal shear plane that developed parallel to foliation. (d) Stereonets of geological structures.

the slope the regional faults may contribute to slope destabilization and rockslide formations. This is particularly true in the gneissic areas (that covers more than the three-quarters of the western Norway area) where the fault cores appear weaker

than the host rocks and focus the gravitational structures (from back-cracks to transfer fractures). Also, and as expected, these regional faults are characterized by wide damage zones that are a dense network of mainly parallel faults and

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Fig. 26. (a) Picture from the helicopter towards the NE showing the 1 Mm3 volume of rock (limited by a dotted line) detached from the slope by a NE–SW vertical crack, site 22, Storfjord (located in Fig. 5). (b) Picture to the SE from the helicopter of the SW– NE crack. The steep NW– SE foliation is seen. (c) Mapping of structures (UTM co-ordinates, zone 32N). (d) Stereonets of field measurements.

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Fig. 27. Geological structures observed on the SW side of Tafjord (sites 26 and 27 of Storfjord region; located in Fig. 5). The slope instabilities are largely developed subsequent to the reactivation of the pre-existing structural pattern with: (1) opening of cracks along north–south steep regional faults and the east– west vertical foliation; and (2) development of basal sliding planes along the moderately fjord-dipping exfoliation joints. (a) Orthophotograph (UTM co-ordinates, zone 32N) with the two unstable sites, the foliation attitude, the north– south steep regional faults (solid white lines) and the geological contact between two gneissic tectonic units (solid black lines). (b) Stereonets of structures with joints as thin black dashed lines, pre-existing faults as thin black filled lines, foliation as thick black lines, exfoliation as thick grey lines and sliding planes as thick black dashed lines. (c) Photograph of the main crack at site 27 that developed along the east–west vertical foliation.

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Fig. 28. (a) Photograph from the helicopter of the unstable face at Steiggjeberg in Sogn og Fjordane (located in Fig. 3). (b) Extract of the geological map with the peculiar tectonic setting of the unstable scarp (star) that lies in the damage zone of a moderately fjord-dipping Devonian regional normal detachment (UTM co-ordinates, zone 32N). (c) Close view of one (supposed normal fault) of the numerous faults that characterizes the damage zone of the Devonian detachment and that contributes to weaken the scarp.

fractures. The observations fit with the consensus that links many gravitational slope deformations worldwide with the presence of regional faults (see the references cited in introductory section of this paper). In that sense, Bo¨hme et al. (2011) show the spatial relationships between the rock slope instabilities in western Norway and the largest fault zones. Also, because large fault zones structurally control the hydrographical network in Norway, the side slopes of the valleys and fjords may contain a non-negligible amount of associated fractures. The latter mainly striking parallel to the slope (to the main fault segment) serve to localize the back-crack of instabilities. In addition, our study reinforces the concept that the high susceptibility to rockslides of the steep slopes in western Norway is directly linked to the high spatial density of faults at various scales and with various orientations. This high fault density is related to the age of the Norwegian bedrocks that cumulated a long history of successive brittle

tectonics from the time of the Caledonian Orogeny. Subsequently, it is not rare to observe two cross-cutting orientations of faults and associated fractures at one site that promote the slope deformation (Table 1).

Combination of favourably orientated structures with respect to the slope in rockslide development Intuitively, and subsequently to the field analyses, it appears that a combination of several geological factors, which taken one by one decrease the slope stability, will dramatically further enhance the destabilization. Site 6b in Storfjord (location shown in Fig. 5) is one example where a combination of pre-existing geological factors allowed the development of a large unstable volume of 15 Mm3 on an approximately 458 SE-dipping slope (Fig. 29). The site is

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Fig. 29. (a) Orthophotograph, (b) structural mapping (UTM co-ordinates, zone 32N) with open fractures in thin lines, wide back-bounding cracks in thick lines, basal sliding plane in festooned line and (c) stereonets of the potential rockslide at site 6b in Storfjord area (located in Fig. 5). An unfortunate combination of six pre-existing geological factors allowed the development of a large unstable volume of 15 Mm3: (1) the instability developed in the weak mica-rich rocks; (2) foliation dips towards the fjord and guided the creation of a basal shear plane, underlined by a newly formed 20 cm thick gouge; (3) the basal shear plane may be located at the interface between the unfolded upper and the folded lower parts of the slope; (4) and (5) NE–SW and NW– SE reactivated pre-existing regional faults (see stereonet) are open fractures and transfer fractures, respectively, into the unstable block and the back-crack on its northern part also developed along a NE–SW vertical fault (thick white line in b); and (6) the southern part of the back-crack (thick white line in b) follows a north–south regional cataclastic fault.

located on the mica-rich units, and these rocks are prone to destabilization because of their intrinsic weakness. The foliation dips approximately 308 SE towards the fjord and a shear plane with newly

formed 20 cm-thick gouge developed subparallel to it at the base of the unstable volume (see also Henderson & Saintot 2011). An important field observation is that below the base of the unstable

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volume the rocks are folded. Because of the contrast in rheology between folded and unfolded rocks, we propose that the basal shear plane formed at the interface between the two units. Open joints in the unstable block trend roughly NE–SW. NW– SE-trending joints act as transfer zones. An analysis in the surrounding area of the unstable part of the slope revealed that both the NW –SE and NE–SW vertical fractures belong to the pre-existing structural brittle grain and are commonly epidote-rich faults. On the northern side, the main back-crack follows a NE –SW structure. Towards the south, the back-crack with a width of 3 m is an extensional fracture that trends north–south, similar to the inherited regional cataclastic faults. Three regional trends of pre-existing vertical faults, the fjorddipping foliation, the different rheology between the folded lower and the unfolded upper parts of the slope (with, at their interface, the localization of the basal sliding plane), and the intrinsic weakness of the mica-rich rocks represent six factors that led to the development of the rockslide.

Discussion, concluding remarks and future work Table 1 summarizes the field observations at the 72 examined sites of western Norway. The table is divided in two parts according the host rocks of the instabilities as we assume that one of the primary

factors to consider is the lithology of the rocks, that is, weak v. hard rocks and, respectively, amphibolites, schists and micaschists, weathered mafic and ultramafics v. gneisses (see the subsection ‘The lithological factor in slope instability development’). Out of the 72 studied instabilities 13 sites are grouped as being hosted by weak rocks and are set in the top part of Table 1. A second criterion of site classification in Table 1 is the number of structures that hold or guide the gravitational deformation (as documented by detailed field examples in previous sections) and a third criterion is the estimated maximal unstable volume, with the criterion values increasing from top to bottom in each of the two lithological groups. Table 1 also provides by sites qualitative assessments on the type and nature of the pre-existing structures involved in the gravitational deformation. Table 1 and the diagram in Figure 30a show that 10 out of the 13 instabilities that developed in the weak rocks do not necessarily comprise reactivation of more than one prominent (i.e. large and persistent at the scale of the instability) pre-existing structure. Only two out of the 13 instabilities hosted by the weak rocks developed with the reactivation of two pre-existing structures. Site 6b in Storfjord is the major exception of these 13 sites, with five preexisting geological structures involved in its development. However, this analysis tends to confirm that such rocks cannot hold on the steep slopes of western Norway simply because they are intrinsically weak (i.e. pre-existing structures would not

Fig. 30. (a) Diagram of the type of gravitational slope deformation and of the estimated maximum volume of the 72 studied instabilities of western Norway as a function of the number of pre-existing structures that guide or hold the gravitational deformation. There is a slight correlation between the numbers of pre-existing structures and the volume of the instability. (Colour code: white and black for sites in weak and hard rocks, respectively; shape code: circle, triangle and star for rockfall, rock topple and rockslide, respectively.) (b) Hypothetical scheme (perspective view) of pre-existing structures that control in a way the volume of the unstable zone (grey part): the old large fault zone focuses a back-bounding crack, a transfer fracture is guided by a axial fold plane and the basal shear plane forms along foliation (with h being a sufficient height to generate weight/force that can destabilize a mica-rich layer and allow the sliding).

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be necessary to the development of a gravitational instability). That is not the case if we analyse the 59 unstable sites developed in the harder gneissic rocks of Norway. We observed that at such 54 localities pre-existing prominent structures are required to develop the instabilities (in agreement with worldwide structural studies on rockslides; see the references in the introductory section to this paper) and, to a greater extent, that at 37 localities more than one inherited prominent structures played a significant role in the formation of the instability (Table 1 & Fig. 30a). This does not preclude that the other necessary gravitational structures (assuming that three orientations of gravitational structures lead to fully detach a rockslide at the back, at the side and at the bottom) are not reactivated structures. They may also be reactivated but would be smaller and more local pre-existing structures (like typical sets of joints) than the prominent structures that we presently count (like regional brecciated/cataclastic faults or decametre- and hectometre-scale folds). The gravitational deformation in the hard rock like the gneisses of western Norway is only possible because of the anisotropy of strength subsequent to the tectonics; that is, the metamorphic foliation, the presence of folds and brittle structures at all scale. The strength of such hard rock slopes will dramatically decrease if structural features are favourably orientated with regard to the slope. For instance, in the metamorphic rocks of western Norway, one obvious structural factor that controls development of slope instabilities is the foliation and, more specifically in the dioritic gneiss, its mafic weak layers. As observed at 11 localities, the foliation is vertical, roughly slope-parallel and contributes to columnar shape-like failures. The foliation can also be fjord- or valley-dipping and reactivated to form basal sliding planes as observed at 30 localities. However, many field examples showed that foliation dipping towards the fjord or valley is far from being the only structural factor that led to the full development of slope instabilities. The associated large failure surfaces (back-crack and transfer fractures) responsible for the detachment of a block from the slope should be obviously present and these are rarely newly formed structures. When such failure planes are newly formed they are, indeed, guided by the ductile grain of the rock (as a particular fold geometry), but very commonly gravitational forces simply reactivate preexisting brittle structures. It follows that at most of the sites there is a combination of several favourably orientated structures with respect to the steep slopes that will dramatically decrease the slope stability and promote gravitational failures. Also, Henderson & Saintot (2011) present a new qualitative approach to the spatial variation in susceptibility

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at the scale of Storfjord based on structural criteria and demonstrate that the sites with the highest susceptibility record the greatest number of identifiable gravitational structures. An important loss of rock-mass strength is linked to the high density of pre-existing ductile and brittle structures of various orientations that characterized the bedrock of western Norway. The direct consequence is that there are few slopes exempt of gravitational deformation in western Norway. The rock mass of western Norway suffered at a large scale an important tectonic weakening. Any further weathering (as meteoric) will enhance the destabilization of the slope by the weakening of rocks and preferentially of its geological structures (Bachmann et al. 2004). This last point will also provide the assumption that the volume of the slope instability may be in a certain limit, controlled by the presence and the reactivation of the weakest parts of the slope, that is, the large geological structures (Fig. 30b). Indeed, we observed at 24 sites that borders of the gravitationally deformed slope coincide with large inherited structures. This may explain the slight correlation between the numbers of pre-existing prominent structures and the volume of the slope instability (Fig. 30). Where the gravitational deformation occurred with a destabilization involving two or more pre-existing prominent structures the volume of the instability tends to exceed 1 Mm3 (with the exception of one rockfall of 0.001 Mm3; Fig. 30a). The dimension of the unstable zone may be constrained by the hundreds-of-metre-scale spacing and persistence of large structures (as well exemplified by the distribution of the slope instabilities in the vicinity of sites 7, 7b and 8 in Storfjord; see Fig. 23). Henderson & Saintot (2011) established this relationship in more detail in the Storfjord area where the largest unstable volumes are at sites with the largest number of well-developed gravitational structures. These gravitational structures reactivated preexisting structures and also clearly correspond to the borders of the instabilities. Therefore, the preexisting large structures not only locate the gravitational deformation, but may limit and impose the unstable volume. The two best examples to illustrate ˚ knes (described in Jaboyedoff such setting are the A et al. 2011) and Gikling rockslides (Saintot et al. 2008) (see Figs 16 & 17). The analysis of the spatial distribution of the number of pre-existing structures observed by instabilities coupled with the unstable volume over western Norway tends to confirm this point. To the north of western Norway and in Møre og Romsdal County, where large ductile and brittle structures are more abundant, there is a clear increase in the volume of instabilities together with an increase in the number of structures involved (Fig. 31).

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Fig. 31. Map of the 72 slope instabilities of western Norway with an indication of the type of gravitational slope deformation, the estimated volume and the number of pre-existing structures that guide or hold the gravitational deformation (background is Fig. 2b).

In addition, the spatial distribution of potential rock slope failures over western Norway shows that there exists a specific clustering in Storfjord (Fig. 31). The assumption to explain such clustering

is, in addition to the steepness of the slopes and the gradient of the ongoing high rate uplift (see Fig. 2 and discussion in Henderson & Saintot 2011), the occurrence of close –tight metre-scale polyphase

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folds in the gneiss (which is not so pronounced towards the Møre Trøndelag Fault Complex and towards Sogn og Fjordane County). We propose that there is a strong correlation between the occurrence of ductile folds and slope stability decrease, and that fold interferences will provide further loss in slope stability. This may be because ductile folding is expected to: (1) decrease the intrinsic strength of minerals (and thus rocks) by stretching and flattening, specifically along the limbs; and (2) increase along a given slope the probability to have favourably orientated grains prone to be reactivated by gravitational forces (or to guide newly formed gravitational structures). Interference of several phases of folding would increase these factors. The typical gravitational slope deformation of western Norway mostly corresponds to rockslide or complex slide (also involving toppling and flexural toppling) with, out of the 72 studied sites, 56 instabilities that developed along a basal sliding planes (i.e. all classified as rockslides in Figs 30a & 31). We previously mentioned that 30 rockslides have a basal sliding plane that formed along the metamorphic foliation on cataclinal slopes. According to Figures 30a and 31, the type of slope deformation is exclusively slide with the increase of both: (1) the number of pre-existing structures involved in the instability formation; and (2) the volume of the instability. Five rock topples and 11

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rockfalls are observed along slopes where none or only one geological structure is clearly reactivated in the process of destabilization. Although most of the studied slope instabilities display well-developed gravitational structures (see above and Henderson & Saintot 2011), thus testifying to quite significant down-slope cumulative displacements, many of the basal sliding planes of the rockslides are rather shallow dipping. Such geometry implies high values of normal stresses that prevent any gliding of the blocks (Table 2). In order to explore this point, we calculated for 20 sites the potential values of cohesion, C, as a function of various friction angles, w, of the sliding basal planes (Fig. 32). Using Selby’s equation (1993, p. 328):

C ¼ r  g  H

sin(b  a)  (sin a  cos a  tan w) 2sin b

where r is the density taken equal to 2700 kg m23, g is the acceleration of gravity (9.8 m s22), H is the the height of the unstable volume, b is the slope gradient and a is the angle of the sliding plane (Table 2). At most of the sites the basal sliding plane is not accessible due to the steepness of the slope. Hence, the analysis is carried out on a relatively limited number of 20 sites for which the geometry of the basal sliding plane is constrained.

Table 2. Parameters used for the calculation of the possible cohesion and friction angle of 20 sliding planes (see Fig. 32) Sites Site 1, Storfjord Site 2, Storfjord Site 2b, Storfjord Site 5, Storfjord Site 6b, Storfjord Site 7, Storfjord Site 7b, Storfjord Site 9, Storfjord Site 10, Storfjord Site 10b, Storfjord Site 12, Storfjord Site 14b, Storfjord Site 22, Storfjord ˚ knes, Storfjord A Oppstadhornet Ha˚rstad Rustøyane Viddalen Tussen Hjellane (down-slope part)

H (m)

b (8)

a (8)

t (MPa)

sn (MPa)

200 100 300 300 100 200 60 200 250 150 100 250 130 200 100 200 80 150 200 100

60 65 65 55 46 60 45 55 60 50 60 65 70 45 32 45 40 45 60 40

40 50 45 25 28 25 35 33 30 38 30 30 25 30 30 35 25 10 20 30

3.4 2.0 5.6 3.4 1.2 2.2 0.9 2.9 3.3 2.4 1.3 3.3. 1.5 2.6 1.3 3.0 0.9 0.7 1.8 1.3

4.1 1.7 5.6 7.2 2.3 4.8 1.3 4.4 5.7 3.1 2.3 5.7 3.1 4.6 2.3 4.3 1.9 3.9 5.0 2.3

a dip angle of the sliding plane; b, slope gradient; H, height of the unstable block; in addition, first-order magnitudes of the shear and normal stresses, t and sn, respectively, acting on the plane due to the weight of the lithostatic column of height H.

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Fig. 32. Calculation of the possible cohesion and friction angle of 20 sliding planes (from Selby 1993; see Table 2 for the values of parameters). The shaded area represents the physically meaningless part of the graph with, towards the black side, unrealistic friction angle values decreasing to 08.

The analysis shows very low cohesion values (from 0 to approximately 1000 kPa) when varying the friction angle values reasonably taken above 58 (Fig. 32). Of note herein is the precise determination of the friction angle close to 308 at the ˚ knes rockslide (the cohesion being negliactive A gible as the mass is moving). At the other sites that are considered inactive or dormant, cohesion values are clearly one –two orders lower than shear stress magnitudes (some units of MPa; Table 2). Hence, the analysis predicts that these instabilities would fail but actually they do not. A first explanation is that the planes still display asperities, typically bridges, giving higher cohesion values than the one predicted by Selby’s equation. This roughness of the planes will be smoothed out by chemical and/or mechanical weathering. Such process is part of the long-term fatigue of the medium and contributes to the required maturation of the basal sliding planes. An additional reason is probably the high values of the normal stresses (Table 2). However, the calculated normal stresses

(Table 2) are first-order values, for example without taking into account fluctuations due to the water pressure. The normal stress may dramatically drop off with large seasonal water inputs, thus triggering motion (Terzaghi 1962; Giraud et al. 1990; Crosta & Agliardi 2003; Cruden 2003; Sartori et al. 2003; Grimstad 2005 among others). For example, the shear and normal stress magnitudes due to the weight of the lithostatic column on the extreme flat-lying plane at the base of the rock slope instability at Viddalen (see the earlier subsection ‘Influence of the sole detachments of the Caledonian allochthons on the slope stability’; Fig. 14) would be approximately 0.7 and 3.9 MPa, respectively (Table 2). In this case, an increase in the isotropic water pressure to a value of only 1.6 MPa would be enough to provoke the reversing of the shear and normal stress relationships triggering the gliding along the plane. Water pressures of some few units of MPa are predictable along most of the basal planes of the potential investigated rockslides when considering the important relief

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at the sites and, thus, the height of the hydrostatic columns. Both roughness of the plane and normal stress variation interplay for the stability of the rockslide. Finally, this geological study provides a basis to construct a map of susceptibility to large gravitational deformation along selected slope gradients taking into account lithology (weak v. strong rocks), anisotropy due to structures (i.e. considering their orientation and their types) and the density of structures. The following list summarizes some crucial parameters that we observed during the fieldwork and may be used to index the loss of stability at a regional scale along high slopes with gradients above 358 in western Norway: (1) weak rocks; (2) foliation towards the fjords or the valley or steep foliation striking roughly slope-parallel; (3) folds and interference folds; (4) Caledonian thrusts cutting the slope; and (5) regional brecciated/ cataclastic faults close to the slope. If all of these features are present, gravitational deformation is very likely to occur. The susceptibility map based on geological/structural criteria, combined with an analysis of past events, will help to complete hazard mapping of western Norway. We are grateful to M. Bo¨hme, E. Anda, H. Dahle, S. Sætre, Tim Redfield and all our companions during the fieldwork in western Norway; and L. Blikra and R. Hermanns for discussions on slope stability. We thank two anonymous reviewers for their help in improving the manuscript. The work described in this contribution is partially supported by the Research Council of Norway through the International Centre for Geohazards (ICG). Their support is gratefully acknowledged. This is ICG contribution No. 279.

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Regional spatial variations in rockslide distribution from structural geology ranking: an example from Storfjorden, western Norway IAIN H. C. HENDERSON* & ALINE SAINTOT Geological Survey of Norway, Leiv Eirikssons Veg 39, 7491 Trondheim, Norway *Corresponding author (e-mail: [email protected]) Abstract: Structural geology has recently become a key topic in landslide research. However, the link between regional structures, their cumulative contribution to rockslide development and their significance in a spatial framework is uncertain. We examine the influence of structures on rockslide susceptibility in the Storfjorden area, a 900 km2 fjord complex in western Norway that ˚ knes and Heggursaksla. We have newly identified 52 includes the monitored rockslide sites of A potential rockslide sites from aerial photographs. The structural features critical for the development of large rockslides (fjord-dipping foliation, basal shear plane and breccia, extensional fracture and transfer fault) have a spatial bias in orientation. Rockslides are more likely to develop in specific fjord orientations that have favourably oriented structures. Therefore, the development of rockslides has a marked spatial distribution that we describe qualitatively with an inventory of structural features. Sites with the full plethora of features display the most movement, the largest volumes and are already under the closest scrutiny with regard to monitoring. These sites are also spatially biased, the largest clustering occurring in west Sunnylyvsfjorden. The utilization of structural criteria can show trends in spatial distribution of rockslide potential and on a regional scale can be an important tool in susceptibility analysis.

The study of rockslides has increased rapidly in recent years owing to the increasing needs of society and as a result of the expanding use of remotesensing techniques (e.g. Van Westen et al. 1997; Dai & Lee 2001; Jaboyedoff et al. 2005). Despite this, the development of multi-scalar, regional- to outcrop-scale field studies (e.g. Chigira 1992; Hungr et al. 1999; Agliardi et al. 2001) are still the exception. Recent studies examine the spatial distribution of rockslides where tectonic, structural and topographical data are combined with spatial analysis (DeGraff 1978; Nichol et al. 2002; Ambrosi & Crosta 2006, 2011; Di Crescenzo & Santo 2007). Other studies shed new light on the underlying geometrical, structural and tectonic causes of rockslides (e.g. Cruden 1976; Varnes 1978; Braathen et al. 2004; Redfield & Osmundsen 2009; Agliardi et al. 2009b; Osmundsen et al. 2009). Our understanding of the underlying geological causes of, and controls on, the development of rockslides and their regional spatial distribution is often based on intuitive or subjective geological observation. However it is only seldom that a more objective and empirical control on rockslide causes and distribution is forthcoming (Sauchyn et al. 1998; Kawabata et al. 2004). Studies focusing on the combination of individual structural controls on rockslides (e.g. Sauchyn et al. 1998; Amin 1999; Bo¨hme et al. 2011), particularly in relation to the interplay between lineament orientation and valley orientation is sometimes overlooked. For example, the control of fall-line sloping structures parallel to planar slopes is a

direct and intuitive control on rockslide development and it is one which will be clearly shown to be an important factor in our field study. However, studies of the interaction of a varying large-scale fjord (valley) architecture relative to different structural trends are rather few and are mostly confined to the field of geomorphology (e.g. Schmidt & Montgomery 1995). Very few have focused on structural aspects of susceptibility (e.g. Irigaray et al. 2003; Romana et al. 2003). Some studies focus on the controlling factors for the spatial distribution of rockslides (Osmundsen et al. 2009; Redfield & Osmundsen 2009) without invoking susceptibility considerations, whereas others have focused on the susceptibility from geographical information systems (GIS), slope aspect and numerical modelling (Van Westen et al. 1997; Dai & Lee 2001; Dorren et al. 2004; Guthrie & Evans 2004; Mansor et al. 2004) without detailed structural analysis. Our aim is therefore not to contribute to further development of geotechnical models on a site scale (e.g. Markland 1972; Hocking 1976; Hoek & Bray 1981) or on a regional scale (e.g. Ward et al. 1981; Fall & Azzam 2001; Fall et al. 2006) but rather to integrate more conventional structural geology techniques with a spatial study of rockslide susceptibility. Such studies of slope tectonics have been relatively few until recently (Gu¨nther 2003; Braathen et al. 2004; Roering et al. 2005; Agliardi et al. 2009a; Saintot et al. 2011). The Western Gneiss Region (WGR) of Norway (Fig. 1a) is historically a place with many

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 79–95. DOI: 10.1144/SP351.4 0305-8719/11/$15.00 # The Geological Society of London 2011.

80 I. H. C. HENDERSON & A. SAINTOT Fig. 1. (a) Regional location map showing the study area (black box) relative to the MTFC and the relatively high density of rockslides Precambrian basement is in pink, Caledonian thrust sheets are in greens and yellows. (b) Profile X–Y through A showing uplift data (Olesen et al. 2000) relative to the clustering of both historic rockslides and potential rockslides. (c) Storfjorden field area showing bedrock geology and simplified foliation form lines (dashed red) in the Precambrian basement. Blue lines are open folds in the Precambrian basement. Green dots are identified potential rockslides.

STRUCTURAL GEOLOGY RANKING

documented rockslides. Enhanced clustering of rockslides in this area may be owing to the presence of a strong gradient in regional uplift (Fig. 1b). Other studies suggest that areas with strong uplift gradients are more prone to rockslide clustering (Jaboyedoff et al. 2003). Our study area, in the Storfjorden Fjord Complex (Fig. 1c), is approximately 900 km2, and comprises Sunnylyvsfjorden, Geirangerfjorden, Norddalsfjorden, Tafjorden and the southern part of Storfjorden, where rockslide occurrence is exceptionally enhanced. We have identified 52 new rockslide sites from initial aerial photograph analysis, airborne LiDAR-DEM (Light Detection and Ranging digital elevation model) and digital orthophotographs. Our study incorporates an inventory of structure types present on these sites (the reader is referred to Saintot et al. 2011 for a detailed description of the sites), regional lineament analysis from the remote sensing, and orientation analysis of the different structures present at the sites relative to the many different fjord and slope orientations. We determined which orientations and types of structures were likely to increase the spatial susceptibility depending on variable fjord orientation. From a qualitative numerical ranking of the different structural types we present a novel method of demonstrating the regional spatial susceptibility of rockslide occurrence.

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gneisses and the overlying Caledonian thrust sheets (Fig. 1c). Over 80% of the surface area consists of granodioritic –dioritic gneiss of the Precambrian basement. Small amounts of quartz mica gneiss and coarse-grained granitic gneiss are present. In the eastern part of the field area a Caledonian thrust sheet consisting of augen gneiss and minor mica schist is present. The majority of potential rockslide sites, up to 70%, are located within the basement granodioritic –dioritic gneiss and appear to display no localization along lithological or tectonic contacts, with the exception of sites 20–23 and Heggursaksla in Tafjord, which are located on the base of the Caledonian thrust sheet.

Methods Stereographical aerial photograph analysis initially identified the sites, followed by helicopter reconnaissance and detailed field structural analysis. Volume calculations and regional lineament mapping were undertaken using regional airborne LiDAR-DEM and associated high-resolution digital orthophotographs. Regional structural data were collected, as well as site-specific structural data.

Results Regional setting Tectonic Our study area is located in an area of enhanced rockslide occurrence associated with enhanced present-day uplift gradients (Fig. 1a, b) that are, in part, glacially-controlled but can also be related to neotectonic activity (e.g. Blikra 1990; Dehls et al. 2000; Blikra et al. 2002; Braathen et al. 2004; Olesen et al. 2004; Redfield et al. 2004) on the Møre– Trøndelag Fault Complex (MTFC). Redfield & Osmundsen (2009) postulated a long-lived topographical high related to lithospheric flexure and normal faulting, creating a steep footwall margin (‘Great Escarpment’). Our study area lies directly on the western edge of this footwall high. This is controversially purported to be still active today. In addition, the only definitive description of a neotectonically active fault in southern Norway appears to directly segment the study area (Anda et al. 2002; Blikra et al. 2002), and can be an influencing factor in the development and spatial clustering of these rockslides.

Lithological The field area transects the contact between the Western Gneiss Region crystalline basement

Site-specific structural controls on rockslide development Our regional mapping has identified 52 rockslide sites. Site-specific structural analysis has generated over 3000 measurements of different types: foliation, exfoliation, joints, fractures, sliding planes, breccias and pre-existing faults, which often form mappable regional lineaments (Fig. 2). On all of the rockslides where the block is fully detached, a well-defined basal sliding plane (BSP) lies subparallel to a fjord-dipping foliation and sites that demonstrate fully developed extensional fractures are localized on fjord-dipping foliation with 208 –358 dipping sliding planes (Fig. 2a). The BSP is most often parallel/subparallel to foliation or is localized in a mica-rich sliver or along a pre-existing breccia structure that is very often subparallel to foliation. In rare cases the BSP is parallel to exfoliation, which may be unrelated to foliation orientation. Block detachment varies from several centimetres up to 10 m. Where the extensional fractures are not fully opened a topographical depression is observed at the rear of the block. Examples of evidence for down-slope gravitational motion are gravitational duplexes at the base of the block (Fig. 2b), partial disintegration of the hanging-wall block by extensional fracture development above the BSP

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Fig. 2. Examples of the structures observed in the field common to most of the examined sites. (a) Site 6B, which shows a detached block with a several-metre-wide back-crack and several secondary subvertical fractures in the block, fjord-dipping foliation and a sliding plane at the base of the block subparallel to foliation. (b) Brittle extensional duplex at the base of site 2B showing down-slope movement. (c) Extensional fractures in the hanging wall to the basal shear zone at site 2B. (d) Non-cohesive fault rock breccia on the sliding plane at the base of site 6B. The inset photographs shows a SEM (scanning electron microscope) image of the breccia texture.

(Fig. 2c) and the development of a non-cohesive breccia (Henderson et al. 2009) where sufficient movement along the BSP has occurred (Fig. 2d). The main regional lineament trend is NNE– SSW or NE –SW (Fig. 3), with minor trends of WNW–ESE and east– west. These consist of steeply dipping, cemented breccia fault zones (Fig. 2e). Translational rockslides are by far the most predominant geometrical type of rockslide and often have large potential volumes, occurring where the pre-existing structural geometries are favourable: the presence of fjord-dipping foliation, a BSP, a fault breccia on the BSP, the presence of exfoliation (to weaken the foliation planes), a steep extensional fracture at the back of the block and a through-going transfer fault to detach the block. A conceptual model of the development of rockslides within the Storfjorden area is shown in Figure 2g and is similar to that developed by others (Hoek & Bray 1981; Chigira 1992; Braathen

et al. 2004). Where few or none of these structural criteria are present, there is generally no evidence for activity in the hanging-wall block, such as preferential hanging-wall disintegration by fracturing (Fig. 2c), kinematic evidence (Fig. 2b) and breccia development (Fig. 2d). Extension fractures are localized on, and reactivate many, pre-existing geological lineaments. Small-fold development within the foliation also appears to have a large effect on the development and movement of rockslides. More details of the relationships between pre-existing ductile structures and the geometrical development of potential rockslides are discussed by Saintot et al. (2011).

Regional structural controls on rockslide development Regional mapping demonstrates that the foliation forms a series of complex recumbent isoclinal

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Fig. 2. (Continued) (e) A 1 m-wide pre-existing cataclastic zone forming the eastern margin of site 7B acting as a transfer structure. (f) Basal shear zone at site 2B intersecting topography showing a 15 cm displacement away from underlying talus. This suggests recent movement. (g) Schematic model of the assemblage of different structural types necessary to cause failure, based on the 52 sites studied here.

folds in the host granodioritic –dioritic gneisses, which is then refolded on open folds with ENE – WSW fold axes where the dip of the foliation is between 208 and 408 (Fig. 1c). The fjord system in Storfjorden consists of several east –west- and several north –south-trending fjords. We contend that the combination of the different structural

types, their regional variation in orientation, the variation in regional lineament orientation and how these different structural aspects interact with a varying fjord orientation can produce a spatial variation in rockslide susceptibility. For example, we identify eight areas where the regional lineament orientation is different. Lineament mapping (Fig. 3,

84 I. H. C. HENDERSON & A. SAINTOT Fig. 3. Digital aerial photographs acquired simultaneously with airborne LiDAR-DEM data over the whole study area. Both the digital orthophotographs and the airborne LiDAR-DEM data have been used to create a lineament map. We divided the area into eight different domains and demonstrated that different lineament sets are present in different ˚ knes site). areas of the study area (the inset shows an example of the combined airborne LiDAR-DEM and orthophotography from the A

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Fig. 4. The orientation of different key structures relative to the fjord geometry (‘quarter’ stereonets are structure dip). The background map is the detailed digital orthophotography. Structures have a distinct preferred orientation, suggesting that specific geographical areas will have an enhanced potential for rockslides if several structures have favourable orientations and that there will therefore be a spatial variation in rockslide susceptibility.

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Fig. 5. Matrix constructed from the structural mapping of the different sites and the ranking of these based on the number of structural criteria present that we contend are important for the development of a rockslide.

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areas A –H) from the digital orthophotographs and airborne LiDAR-DEM shows a wide variation in lineament orientation in different fjords. (The inset in Fig. 3 shows an example from one site with the combined digital orthophotographs, airborne LiDAR-DEM and simplified mapping.) Area A (Fig. 3a) displays predominantly WNW– ENE lineaments that may, for example, act as transfer structures, whereas a limited number of north– south lineaments may act as extensional backfractures. For a rockslide in area B (Fig. 3b) both east –west (predominant) and NNE– SSW (minor) structures may act as transfer structures and extensional back-fractures, respectively. Area C (Fig. 3c) displays a major NE– SW trend (and a minor north–set set) that may act as transfer structures, whereas a limited number of NW–SE structures may act as extensional back-fractures. In area D (Fig. 3d) a predominant set of WNW –ESE structures may act as extensional back-fractures.

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There are several minor lineament trends but not in any abundance to function as north –south transfer structures. Very few potential rockslides are observed here (except for sites 11 and 12) relative to the other areas. In area E (Fig. 3e) where there are also very few lineaments developed (with many oblique to the slope). This area is also limited in rockslide development (cf. Fig. 1c). Area F (Fig. 3f) has two well-developed lineament trends, ENE –WSW and north–south. The NW –SE orientation of the fjord means that these structures are not ideally oriented for potential rockslides, except in the NW part where several instabilities are observed (see Fig. 1c). Area G (Fig. 3g) has a major NE–SW and a minor NW–SE trend. These orientations are not ideal for rockslide development as they are both highly oblique to the east –west fjord orientation and, indeed, this is a geographical domain with limited rockslide potential. Area H (Fig. 3h) has few north–south lineaments and also has poor

Fig. 6. The spatial distribution of different localities plotted by volume (as calculated using the airborne LiDAR-DEM data). The localities with the largest volumes are concentrated on the western side of Sunnylyvsfjorden.

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rockslide potential. In area H there is extremely limited evidence for back fractures. Insufficient lineaments are present to localize fracture and those that are present have a high obliquity to the slope. The different structural types have a strong orientation bias (Fig. 4), which may mean that their influence in rockslide development is biased towards particular orientations of fjords. For example, foliation dips, on average, 308 but shows a strong bias to SE-dipping (Fig. 4a). Therefore, foliation is likely to play a major role in developing potential rockslides (Jaboyedoff et al. 2011) preferentially on SE-facing fjords. This is indeed the case ˚ knes, Heggursakin reality where the largest sites (A sla, 2B, 4, 5, and 6B) are all located on SE-facing slopes. Extensional fractures have a wide variation in orientation (Fig. 4d), but have a predominant NNE–SSW trend, suggesting that north –south

fjords (in this case predominantly Sunnylyvsfjorden) are more susceptible to rockslides. On a regional scale, structures potentially acting as transfers have a predominantly north –south trend (Fig. 4e). Therefore, fjord systems that are north – south trending may develop rockslides more easily because structures able to function as transfer structures are already present. Basal shear planes and associated breccias have two main trends. These are NW- and SE-dipping and NNE-dipping (Fig. 4f), suggesting that SE-facing parts of fjords (the southern part of Sunnylyvsfjorden) and the outer northern part of Tafjord, the southern parts of Tafjord and Norddalsfjorden, and south Geirangerfjorden would be more susceptible to rockslides. Joint orientation shows an extremely variable orientation and does not represent a strong bias for any fjord orientation on a regional basis and is isotropic on a regional scale (Fig. 4b). However, small-fold

Fig. 7. The spatial distribution of different localities plotted by the amount of structural criteria present (based on the structural mapping). This, like the volume data presented in Figure 6, demonstrates that the sites with the greatest number of structural criteria are located in the same areas.

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orientations show a very distinct SE-plunging orientation, suggesting that SE-facing slopes would be more susceptible (Fig. 4c). Exfoliation (Fig. 4g) has two main trends; either NNE- or SSW-dipping or SE-dipping, many exfoliation planes are parallel to foliation and are therefore controlled by foliation orientation. All three of these orientations would affect susceptibility if the exfoliation enhanced weakening of foliation planes, therefore allowing easier development of BSPs.

Structural ranking of rockslide sites We constructed a simple cumulative structural inventory that categorized the 52 rockslides in terms of structures favourable for sliding (Fig. 5). We also included geomorphological evidence for recent activity and GPS (global positioning system) measurements. These last two factors

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were weighted relative to the structural criteria because they constitute direct evidence for movement (these have a value of 2; Fig. 5). The results represent a relative susceptibility to failure of the various sites on a regional scale. First, a very wide spectrum of cumulative structural features was observed per rockslide site, based on a scale from 0 to 9. Secondly, approximately 80 –85% of sites show three criteria or less. It is notable that none of these sites display a fault breccia on the sliding plane, suggesting that the presence of a fault rock is necessary to facilitate movement. Seventy per cent of the sites with four structural criteria or more in the structural inventory have a gouge. In addition, it is clear that both geomorphological evidence for recent activity and GPS movement occur only where a breccia is present on the sliding plane. Of the 52 sites only 21 (c. 40%) have a fully developed sliding plane. Of this subset of localities with a

Fig. 8. The spatial distribution of different localities plotted by evidence for recent activity (based on geomorphological observations in the field, such as those shown in Fig. 3f) and those displaying a gouge in the BSP. This shows remarkable spatial agreement with both Figure 6 and Figure 7 (volume and amount of structural criteria, respectively).

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sliding plane, only eight (15% of the total sites) have a fault rock developed. Figure 5 also shows that 80% of the localities with four or more structures have small-scale structures that are favourable for slide development, such as fjord-plunging fold axes. This point is discussed in more detail in Saintot et al. (2011).

Spatial variation in susceptibility There is a distinct spatial distribution of the different rockslide sites relative to their volumes (Fig. 6), with the largest potential slides concentrated in west Sunnylyvsfjorden or in Tafjord. The regional variation in structural criteria presented in Figure 5 is also notable (Fig. 7). Sites with the most structural criteria (eight or nine) are found exclusively on the western side of Sunnylyvsfjorden or in Tafjord. Site 21 is the only exception and is probably the result of its localization on the

Caledonian thrust zone. Figure 8 shows that most of the sites that contain a gouge have also had recent activity. Comparison of these features with the structural criteria presented in Figure 7 demonstrates that there is also a relationship between structural criteria and the presence of a gouge. Figure 9a shows the relationship between fjorddipping foliation orientation and potential rockslide location: the shaded areas show where foliation dips between 208 and 358 down towards the fjord. There is a remarkable correlation between the areas with favourable foliation for sliding and those areas where the biggest potential slides are observed (notably including both of the key localities presently ˚ knes and Tafjord). Only two being monitored: A minor rockslides volumes occur outside of these foliation demarcated areas. Figure 9b shows a susceptibility map constructed from the site information on volume alone and this shows remarkable correlation with areas of fjord-dipping foliation.

Fig. 9. (a) Susceptibility map based on fjord-dipping foliation relative to mapped areas with potential failures. There is good agreement between the inferred susceptible areas and the area where rockslides are actually mapped. These delineated higher susceptibility areas agree remarkably well with the largest mapped potential rockslides, most of which lie within these areas. Only two small volume rockslides are located outside of these areas.

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West Sunnylyvsfjorden and Tafjord have the highest average volumes of potential rockslides and also areas with the most consistently fjord-dipping foliation. It is also notable that these are the areas with the two highest structural criteria inventory. Figure 10 demonstrates that the vast majority (over 75%) of the potential rockslides are less than 1 Mm3. There is also a rapid drop-off in potential rockslide volume over approximately 5 Mm3. ˚ knes and Heggursaksla (Fig. 1c) are the only two A sites that show active movement and these happen to be amongst the largest sites in terms of volume. Movement here is documented from several millimetres up to 8 cm per year on average. Therefore the fact that these sites are high-priority monitored objects is noticeable from both our structural inventory ranking and from the spatial distribution data that we present.

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Discussion We have newly identified 52 potential rockslides clustering in an area of approximately 900 km2 in western Norway. Braathen et al. (2004) provided a comprehensive review of the structural geometries of the most prominent rockslide sites in Norway; we identify the same structures here and demonstrate that these are cumulatively important for the development of rockslides. Only the most active and/or the largest potential rockslides have a full inventory of these structures (Henderson et al. 2006; Saintot et al. 2011). We demonstrate that a spatial variation of these structural elements leads to a spatial variation in the susceptibility to rockslide development. Ganerød et al. (2008) have identified the critical geological structures on the ˚ knes slide that are identical to the highly active A

Fig. 9. (Continued) (b) Susceptibility map based on the average slide volume for individual geographical areas (fjords). The average structural criteria inventory is shown as numbers. The areas of west Sunnylyvsfjorden and both south and north Tafjord have the largest average volumes. West Sunnylyvsfjorden and north Tafjord also have the greatest number of structural criteria.

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Fig. 10. (a) Frequency histogram of the occurrence of potential rockslide sites with respect to their volume. This demonstrates that most sites have volumes of less than 1 Mm3 and only nine sites have volumes above 8 Mm3. However, ˚ knes and Heggursaksla). (b) Log-normal two of the largest sites are also those that are documented as being active (A plot of structural criteria with rockslide volume. This demonstrates that it is only the largest volumes that display the greatest possible number of structural features.

structures we observe regionally and play an obvious role in the development of the gravitational structures over the whole fjord system. Ambrosi & Crosta (2006) presented a similar study concerning the spatial distribution of rockslides or ‘sackung’ within the Alps related to tectonic features. They demonstrate the importance of similar lithological and structural constrains on rockslide development, and their spatial distribution. Bovis & Evans (1996) also described an identical suite of structures to those that we determined as critical to develop large potential rockslide volumes. Such regional analytical approaches to rockslide susceptibility based on geological information have also been carried out. Hungr et al. (1999) provided an extensive study (up to 3500 sites) of rockslides and rockfalls on a regional scale in Canada, and outlined a risk analysis based on magnitude–cumulate frequency considerations. However, we believe that our approach, where the relative susceptibility of different rockslides on a regional scale is based on a spatial variability of structural elements, is a valuable concept that could be applied elsewhere. The regional clustering of rockslides may be related to the proximity of the long-lived and still active MTFC to the NW. This is also the location of the highest gradient of uplift recorded in western Norway (Hicks et al. 2000; Bungum et al. 2005). We propose that rockslide development in this area is, in part, related to recognized, but hitherto unassessed, neotectonic factors (Blikra et al. 2002; Olesen et al. 2004). We document an

interplay of structural controls, their spatial variability relative to fjord geometry, and what role they may play in potential rockslide development and distribution. The make-up of the structural inventory (fjord-dipping foliation, sliding plane and breccia development, extensional detachment, and transfer structures) and how these structures interact with a variable fjord geometry is critical to the geographical distribution of rockslides. The spatial pattern of rockslides is also determined by the preexisting lineaments, which we have mapped as having different orientations in the different fjords. The underlying geology and, particularly, the lineament pattern often plays a critical role in rockslide distribution (Bachmann et al. 2004). We are therefore able to determine areas of spatially enhanced rockslide susceptibility based on the structural inventory. We observe that fjorddipping foliation is generally the main precursor for the development of fjord-dipping sliding planes and the development of subsequent gouges that weaken the base of the rockslide. Suitably oriented foliation is therefore an important factor in determining spatial variation in rockslide susceptibility on a regional scale but cannot solely account for the distribution of rockslides. Ranking of the 52 different sites based on their structural inventory to assess the structural influence ˚ knes and Hegon rockslides development places A gursaksla, which have the full plethora of structural criteria, as the most susceptible. This confirms that this novel approach is a viable method of identifying

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the key structures, structural processes and the most important sites within a rockslide susceptible area on a regional scale. We also observe a distinct spatial clustering of rockslides with the largest volumes and with the most favourable structures for rockslide development. Our data suggest that the west side of Sunnylyvsfjorden and Tafjorden are the two areas most prone to rockslide development (where both ˚ knes & Heggursaksla are located). A Notably, the clustering of rockslides that we observe has little to do with the pattern of regional bedrock geology. The area with the clustering of most rockslides, displaying the highest average volumes and the most structural criteria in west Sunnylyvsfjorden, lies in the middle of the Precambrian Basement granodioritic gneiss, far from the contacts with the overlying Caledonian cover. The only exception to this is the northern side of Tafjord, where the Heggursaksla slide is located close to the autochthon–allochthon interface. Application of this methodology to similar areas containing a high concentration of potential rockslides, where no monitoring is yet present, is a powerful and inexpensive tool in prioritizing sites for monitoring. We believe that the analysis of regional structural data in this way in active rockslide systems can: (1) help to determine the spatial distribution of potential rockslide activity; (2) help to understand the contributing factors promoting failure; and (3) provide a better geological foundation for susceptibility analysis.

Conclusions † The 900 km2 Storfjorden area is coincident with a high regional uplift gradient, adjacent to the still active Møre –Trøndelag Fault Complex (MTFC). Fifty-two new potential rockslide sites have been discovered during this study, their clustering showing a correlation with uplift and historical rockslides pattern. † Regional lineaments and mapped structures (fjord-dipping foliation, sliding planes, basal breccias, extension fractures and transfer structures) that delineate potential rockslides have preferred orientations, meaning that rockslides are more likely to develop in specific fjord orientations that have favourably oriented structures. † Cumulative ranking of all sites by structural features demonstrates that those with the highest structural ranking are those that display movement, form amongst the largest volumes and are already being monitored. † The largest volumes and those with the most structural criteria are spatially biased, the largest clustering occurring in west Sunnylyvsfjorden.

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† Comparison of the structural ranking method shows that areas with suitably oriented fjorddipping foliation (between 208 and 408) are coincident with sites displaying the bestdeveloped fault gouges and also sites that have the largest volumes. † The spatial distribution of potential rockslides appears to have only a limited relationship to the regional-scale bedrock pattern, as demonstrated by the clustering of sites within the large, crystalline, hard granodioritic mass on the west of Sunnylyvsfjorden and the few located on the minor lithological units and/or close to the major tectonic contacts. L. H. Blikra is thanked for providing interesting discussions in the field. Reviews by G. Crosta and an anonymous reviewer greatly improved the manuscript. This project was funded by the Norwegian Government through the intermunicipality project ‘Aaknes-Tafjord’. K. Jogerud and T. Bergeng from Stranda Kommune are thanked for logistical help in the field. Partial funding was also received from the International Centre for Geohazards. This work is ICG publication No. 278.

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STRUCTURAL GEOLOGY RANKING Geophysical Union, 85, (Fall Meeting Supplement), Abstract H41B-0296. Mansor, S., Shariah, M. A., Billa, L., Setiawan, I. & Jabar, F. 2004. Spatial technology for natural risk management. Disaster Prevention and Management, 13, 364–373. Markland, J. T. 1972. A useful technique for estimating the stability of rock slopes when the rigid wedge sliding type of failure is expected. Imperial College Rock Mechanics Research Report 19. Nichol, S. L., Hungr, O. & Evans, S. G. 2002. Large-scale brittle and ductile toppling of rock slopes. Canadian Geotechnical Journal, 16, 773– 788. Olesen, O., Blikra, L. H. et al. 2004. Neotectonic deformation in Norway and its implications: a review. Norwegian Journal of Geology, 84, 3– 34. Olesen, O., Dehls, J. et al. 2000. Neotectonics in Norway, Final Report. Geological Survey of Norway Report 2000.002, 135. Osmundsen, P. T., Henderson, I. H. C., Lauknes, T. R., Larsen, Y., Redfield, T. F. & Dehls, J. 2009. Active normal fault control on landscape and rock-slope failure in northern Norway. Geology, 37, 135– 138. Redfield, T. F. & Osmundsen, P. T. 2009. The Tjellefonna fault system of Western Norway: Linking late-Caledonian extension, post-Caledonian normal faulting, and Tertiary rock column uplift with the landslide-generated tsunami event of 1756. Tectonophysics, 474, 106–123. Redfield, T. F., Torsvik, T. H., Andriessen, P. A. M. & Gabrielsen, R. H. 2004. Mesozoic and Cenozoic tectonics of the Møre Trøndelag Fault Complex, central Norway: constraints from new apatite fission track data. Physics and Chemistry of the Earth, 29, 673–682. Roering, J. J., Kirchner, J. W. & Dietrich, W. E. 2005. Characterizing structural and lithologic controls on

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deep-seated landsliding: implications for topographic relief and landscape evolution in the Oregon Coast Range, USA. Geological Society of America Bulletin, 117, 654–668. Romana, M., Sero´n, J. B. & Montalar, E. 2003. SMR Geomechanics classification: Application, experience and validation. In: ISRM 2003 – Technology Roadmap for Rock Mechanics. South African Institute of Mining and Metallurgy, Johannesburg, 1– 4. Saintot, A., Henderson, I. H. C. & Derron, M. H. 2011. Inheritance of ductile and brittle structures in the development of large rock slope instabilities: examples from western Norway. In: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 27– 78. Sauchyn, D. J., Cruden, D. M. & Hu, X. Q. 1998. Structural control of the morphometry of open rock basins, Kananaskis region, Canadian Rocky Mountains. Geomorphology, 22, 313–324. Schmidt, K. M. & Montgomery, D. R. 1995. Limits to relief. Science, 270, 617– 620. Varnes, D. J. 1978. In: Schuster, R. L. & Krizek, R. J. (eds) Landslides, Analysis and Control. Transportation Research Board Special Report 176. National Research Council, Washington, DC, 11–33. Ward, T. J., Li, R. M. & Simons, D. B. 1981. Use of a mathematical model for estimating potential landslide sites in steep forested drainage basins. In: Timothy, R. H. & Pearce, A. J. (eds) Erosion and Sediment Transport in Pacific Rim Steeplands Symposium. International Association of Hydrological Sciences Publications, 132, 21–41. Van Westen, C. J., Rengers, N. & Terlien, M. T. J. 1997. Prediction of the occurrence of slope instability phenomena through GIS-based hazard zonation. Geologische Rundschau, 86, 4004–4414.

Rock slope instabilities in Sogn and Fjordane County, Norway: a detailed structural and geomorphological analysis ¨ HME1,2*, ALINE SAINTOT1,2, IAIN H. C. HENDERSON1,2, MARTINA BO HELGE HENRIKSEN3 & REGINALD L. HERMANNS1,2 1

Geological Survey of Norway, 7491 Trondheim, Norway

2

International Centre for Geohazards (ICG), P.O. Box 3930, 0806 Oslo, Norway

3

Høgskulen i Sogn og Fjordane, Avdeling for Ingeniør og naturfag, P.O. Box 133, 6851 Sogndal, Norway *Corresponding author (e-mail: [email protected]) Abstract: More than 250 rock slope failures have occurred in Sogn and Fjordane County in historical times. So far, 28 sites are known where open cracks indicate that rock slope failures may occur in the future. Detailed structural and geomorphological analyses of these sites have been conducted, and form the basis for an evaluation and comparison of the unstable rock slopes. Four of these sites are described in detail herein. The main characteristics for rock slope instabilities in Sogn and Fjordane are: (1) a preferred location within relatively weak rock units, such as phyllites and weathered mafic gneisses; and (2) the development of most instabilities at convex slope breaks, which are evident as knick-points in the slope profile. Sogn and Fjordane is compared with other Norwegian regions, particularly Møre and Romsdal County, with respect to the spatial distribution of past and current rock slope instabilities. Sogn and Fjordane shows the greatest number of historical slope failures, whereas in Møre and Romsdal a larger amount of potential instabilities is observed. We propose that the larger amount of unstable rock slopes in Møre and Romsdal may be controlled by a locally high gradient of ongoing postglacial uplift and a higher rate of neotectonic activity.

During the last century Norway has suffered several natural disasters with large losses of life caused by rockslides and related tsunamis (Furseth 2006). Owing to the present situation of global climatic change, which leads to a warmer and wetter climate in Norway (Flatøy et al. 2008), an increase in the frequency of rockslides is likely. Therefore a better understanding of present-day rock slope instabilities will optimize monitoring and further research on the susceptibility of future rock slope failures in that region. Historical data and geological studies show that a high concentration of post-glacial gravitational slope failures as well as current instabilities can be found in two Norwegian counties: Møre and Romsdal, and Sogn and Fjordane, both situated in western Norway (Blikra et al. 2006). Our summary shows that Sogn and Fjordane is the county that is historically most affected by rock slope failures, and has the largest loss of life resulting from rock slope failures and related tsunamis (Fig. 1). Seven large historical events in Loen Lake are documented, all of them triggering displacement waves. Two of these lake tsunamis killed a total of 135 people in the years 1905 and 1936 (Bjerrum & Jørstad 1968; Grimstad & Nesdal 1990). In spite of this, a systematic registration

and detailed study of all ancient and current rock slope instabilities has not yet been carried out in Sogn and Fjordane. However, to predict rock slope failures and, therefore, prevent such extensive humanitarian disasters in the future, it is important to investigate the potential for rock slope failures in this county and to systematically study current rock slope instabilities. We present an overview of unstable rock slope areas detected so far in the county of Sogn and Fjordane, and summarize historical events (Fig. 2). To highlight that local geology is one of the major conditioning factors for slope instabilities, the individual characteristics of four selected sites are presented. The different geological features are discussed in order to define the controlling factors for rock slope instabilities in Sogn and Fjordane. Finally, these controlling factors are compared with those of rock slope instabilities in other Norwegian regions and discussed within the framework of international studies. We define rock slope failures as gravitational mass movements of rock masses down a mountain slope and rock slope instabilities as the sources of these failures. Both are classified according to Cruden & Varnes (1996), based on the type of observed or expected slope movement.

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 97–111. DOI: 10.1144/SP351.5 0305-8719/11/$15.00 # The Geological Society of London 2011.

¨ HME ET AL. M. BO

98

1000

Number of casualties

Sogn & Fjordane Møre & Romsdal

100

Hordaland Oppland Rogaland Troms Nordland Sør-Trøndelag Vest Agder Buskerud Telemark Aust Agder Finnmark Vestfold Nord-Trøndelag

10

Østfold

1 1

10

100

1000

Number of rock-slope failures Fig. 1. Number of rock slope failures and related casualties for each affected Norwegian county, (Geological Survey of Norway (NGU) – Landslide database). Sogn and Fjordane County is historically most affected by rock slope failures and has the greatest loss of life caused by rock slope failures and related tsunamis.

Fig. 2. Overview of the study area including historic rock slope failures in Sogn and Fjordane (NGU – Landslide database), identified rock slope instabilities from the NGU’s database and localities described within this paper. For the numbers see Table 1. The inset shows the location of the study area within Norway and indicates the two discussed counties.

ROCK SLOPE INSTABILITIES IN NORWAY

Methodology All of the 28 localities were surveyed from a helicopter for evidence of active movement and for critical factors such as the presence of open fractures at the back of potentially moving blocks (Table 1). In a second step we carried out a structural analysis on the sites following the method suggested by Henderson & Saintot (2011). This included the investigation of: (1) dip and azimuth of the foliation with respect to the slope; (2) the presence of a weakened plane at the bottom of the block that could form a basal sliding plane; and (3) the presence of lateral boundaries of the block, such as transfer structures that could possibly act as structures detaching the block from the mountainside. Volumes that are given in this paper are approximations of the expected maximum volume. They have been calculated with an assumption about the shape of the rockslides based on bordering structures visible in the field.

Overview of rock slope instabilities in Sogn and Fjordane County The foundation of this study is a database with approximately 30 localities that have been identified from previous studies mainly based on the interpretation of aerial photographs or from field observations (Blikra et al. 2002; Bøe et al. 2002) (Fig. 2 and Table 1). These localities could represent a prominent risk for inhabitants and infrastructure, and have therefore been the subject for field analysis (Henderson et al. 2008). Most of the steep cliffs that were indicated as possible rock slope instabilities are indeed potential rockfall areas with small unstable volumes, but do not show all of the structures for large rock slope instability development (Table 1). Henderson & Saintot (2011) demonstrate that rockslides with the largest volumes, that is at least 3 Mm3, only develop when the full set of critical structures are present, that is, a valley-dipping foliation and a weakened plane at the base of the potentially unstable block, as well as existing lateral boundaries of the unstable block. Localities where several of these structures are developed form a high potential for future rock slope failure and, therefore, were the subject for subsequent fieldwork. This is the case for the four localities presented in detail in this study, namely Tussen, Viddalen, Strandanipa and Hyllestad (Fig. 2).

Tussen This potential rockslide area is located at the NE end of the inner fjord system of Sognefjorden (Fig. 2). It is situated 1100 m above the valley

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floor on a NNE–SSW-trending slope that has an average gradient of 358, but exceeds 608 in the uppermost part where the instability is located at the edge of the plateau (Fig. 3a). In this region a Caledonian thrust boundary forms the contact between the phyllitic nappe, in which the instability is located, and the underlying Precambrian gneiss. However, a direct relationship between the presence of this sole detachment and the development of the rock slope instability is unlikely because the detachment is located relatively deep (.500 m) below the unstable area, as well as shallowly dipping (c. 108) and trending oblique to the valley. The bedrock of the unstable area consists of mica schists and phyllites, as well as some regions with mylonitic quartzrich gneisses and quartzites. The foliation dips approximately 308 to SSW and is, therefore, oblique to the valley (Fig. 3b). The whole area is characterized by two sets of regional structures. NNE –SSW-striking lineaments with a persistence of several hundreds of metres contribute to the major structural influence of the unstable area. They are most probably reactivated structures that have now developed large graben-like structures behind the unstable area (Fig. 3b, c). A set of less-developed SE– NW-striking joints act as transfer structures (Fig. 3b). The currently most apparent active block comprises only a part at the front of the whole instability and the volume of this block is estimated to be approximately 7 Mm3 (Fig. 3a, c). It is bounded by a 300 m-long graben-like structure to the NW, while the other three sides are free (Fig. 3c, d). Northeastwards and southwestwards are scarps of older slope failures with widespread talus at their base. The instability is thus a remaining part of the former slope. Several smaller fractures behind the main unstable block have been observed until approximately 250 m inside the plateau (Fig. 3a). Those fractures are up to 1.5 m open and up to several tens of metres deep, but die out laterally. Using these fractures to delimit a maximum volume and to estimate a worst-case scenario gives a volume of 10 Mm3. This unstable rock slope area is therefore one of the largest instabilities in Norway. Measured opening vectors, determined by matching definite pairs of edges on the opposite walls of a fracture, indicate movement directions perpendicular to the NNE–SSW-striking grabenlike structures (Fig. 3b). A major basal sliding plane could not be observed in the field, but thick scree deposits may cover its trace. However, a limited number of field data indicate a third set of joints, dipping approximately 308 to the ESE and hence towards the valley (Fig. 3b). These structures are clearly visible on both side-cliffs and may contribute to the formation of a basal sliding plane. A sliding plane at the base of the unstable block with

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Table 1. Overview of the 28 investigated sites No.

Studied sites

Lithology

Type of expected movement*

Open fractures at the back of potentially moving blocks

Tussen

Phyllite

Rockslide

Yes

2

Viddalen

Phyllite

Yes

3

Hyllestad

Conglomerate

4

Strandanipa

Phyllite

5

Almenipa

6

Byttejuvet

Quartz monzonite (metamorphic) Gneiss (mafic)

Rockslide, rock topple Rockfall, rock topple Rock topple, rockfall Rockslide

7

Aurland-Fla˚m

8 9

Yes Yes Yes

Rock topple

Yes

Phyllite

Rockslide

Yes

Fosse Gjøringbøfjellet

Phyllite Gneiss (felsic)

No Yes

10

Gra˚berget

Gneiss (felsic)

Rockfall Rockfall, rock topple Rock topple

11

Hallandsberget

Gneiss (felsic)

Rockslide

Yes

12

Hjellane

Phyllite / Gneiss (felsic)

Rock topple, rockslide

Yes

Yes

Geomorph. model (Fig. 8)

Yes

b

Yes

c

Yes

a

Yes

a

Yes

b

One side defined Diffuse lateral boundaries

Yes

a

Yes

c

Not defined One side defined One side defined Not defined

Yes Yes

a a

Yes

a

Yes

c

Diffuse lateral boundaries

Yes

c

Weakened plane at the bottom of the block

Lateral boundaries of the unstable block

Oblique to valley, moderate dip Towards mountain, moderate dip No foliation

Not observed

Towards mountain, moderate dip Towards valley, steep dip

Not observed

All sides defined All sides defined All sides defined All sides defined All sides defined

Towards mountain, steep dip Gentle fold with axis azimuth perpendicular to slope strike and shallow dip towards valley (southern part) Unknown Towards mountain, moderate dip Towards mountain, moderate dip Varying, lower part towards valley Towards fjord, moderate dip

Not observed

Yes Not observed

Yes

Not observed

Not observed Not observed Not observed Yes Not observed

¨ HME ET AL. M. BO

1

Knickpoint

Dip and azimuth of the foliation with respect to the slope

Jølstravatnet

Gneiss (felsic)

Rockfall

No

Not observed

Not defined

No

–

Not observed

No

–

Not observed

One side defined Not defined

No

–

Yes

Not defined

Yes

b

Not observed

All sides defined Not defined

No

–

Yes

Towards valley, moderate dip Towards mountain, moderate dip Oblique to valley, vertical Towards valley, moderate dip Oblique to valley, moderate dip Unknown

14

Katlenova

Rockslide

Yes

15

Kloppuri

Quartz diorite (metamorphic) Gneiss (mafic)

Rockfall

No

16

La˚nefjorden

Gneiss (felsic)

Rockslide

Yes

17

Nepangerhaugane

Gneiss (felsic)

Yes

18

Nærøydalen

Gneiss (mafic)

19

Oppigardshyrna

20

Osmundsneset

Quartz monzonite (metamorphic) Gneiss (felsic)

Rockfall, rockslide Rockfall, rock topple Rock topple, rockslide

No

–

Yes

Towards mountain

Yes

One side defined

Yes

c

Rockslide

Yes

Not observed

b

Rockslide

Yes

Yes

b

Rockslide

Yes

Yes

c

Rockfall Rockslide

No Yes

Unknown Oblique to valley, steep dip

Not observed Not observed

Not defined Not defined

Yes Yes

a c

Terakamben Tirskardskreda

Quartz monzonite (metamorphic) Gneiss (mafic) Quartz monzonite (metamorphic) Gneiss (felsic) Gneiss (felsic)

All sides defined All sides defined One side defined

Yes

Phyllite

Oblique to valley, moderate dip Towards valley, moderate dip Towards valley, shallow dip

21

Ovrisdalen

22

Rustøyane

23 24

Steiggjeberget Stopelen

25 26

Rock topple Rockfall

Yes No

Not observed Not observed

Not defined Not defined

Yes No

a –

27

Vidme

Phyllite

Rockslide

Yes

Not observed

c

Vik (Midtfjellet and Hesten)

Phyllite

Rockfall, rockslide

No

All sides defined Not defined

Yes

28

Unknown Vertical, slope parallel Towards valley, moderate dip Subhorizontal

Yes

c

Not observed

Not observed Yes

Yes

ROCK SLOPE INSTABILITIES IN NORWAY

13

*According to Cruden & Varnes (1996).

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¨ HME ET AL. M. BO

Fig. 3. Tussen rockslide area. (a) Sketch cross-section of the unstable area. Structures that dip 308 towards the valley display weakened planes that could develop into a basal sliding plane. (b) Structural field data (equal-area projection, lower hemisphere; in black: foliation poles and mean plane; in grey: fracture poles and mean planes, grey arrows: opening vectors, black, dashed: general slope trend). Grey crosses represent the joint set that may form a basal sliding plane. (c) Three-dimensional view with aerial photograph overlay (www.norgei3d.no) of the study area showing the spatial arrangement of structures. View to the SW. The NW border of the unstable block is a back-bounding graben-like structure and the other three sides are free. Dashed line A –B marks the location of the cross-section. (d) Unstable block, view to the NE.

a dip of 308 coincides with the furthest open fractures going inwards the plateau (Fig. 3a). Recent rockfall activity is detected along the cliff at both ends of the back-bounding graben-like structure, as well as at the front of the block where the proposed sliding plane may crop out. Both ends of the graben-like structure exhibit several smaller blocks or columns with average volumes of 90 m3 that are almost completely detached from the cliff.

Viddalen The unstable rock slope of Viddalen is located in the Aurland region at the SE end of the inner fjord

system of Sognefjorden (Fig. 2). It is situated 260 m above a water reservoir on a north–southtrending slope with an average gradient of 408 (Fig. 4a). The bedrocks are mica-rich schists and phyllites with a foliation that dips moderately SE towards the mountains (Fig. 4b). A flat-lying thrust plane, which forms the contact with subjacent Precambrian gneiss, is cropping out approximately 200 m below the unstable area (Fig. 4a). This site is especially important for detailed investigation because of the possible formation of a post-failure displacement wave in the reservoir and subsequent overtopping of the dam, similarly to the Vaiont slide in 1963 (Hendron & Patton 1987). The

ROCK SLOPE INSTABILITIES IN NORWAY

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Fig. 4. Viddalen unstable area. (a) Sketch cross-section of the unstable area. (b) Structural field data (keys for stereonet as for Fig. 3). (c) Three-dimensional view of a shaded digital elevation model (DEM) of the unstable area with a resolution of 1 m, based on airborne LiDAR (Light Detection and Ranging) data. View to the SE. The dashed line indicates the extent of the unstable block on the surface and the dotted line A–B marks the location of the cross-section. (d) Overview of the area displaying a spring level that coincides with the sole detachment between the phyllitic and the gneissic unit. The dashed line marks the limits of the unstable block. View to the east.

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¨ HME ET AL. M. BO

consequences for the local community in such a scenario would be catastrophic. A 150 m-long NNE –SSW-striking backbounding crevasse and a large NNW–SSE-striking transfer fracture limit the unstable block to the east and to the north, respectively; whereas only a ground depression indicates the southern border (Fig. 4c, d). In the field, the back-bounding crevasse shows a maximum down-throw of 13 m and an opening of up to 15 m in the northern part of the instability. The volume of the unstable block is uncertain because the depth of the unstable part is not known. On the surface it covers an area of 13 600 m2, but it is separated into several smaller blocks by three main fracture sets (Fig. 4b). One set, which includes the back-bounding crevasse, strikes NNE –SSW and is thus slightly oblique to the general slope trend. The second set strikes NW–SE, while the last set strikes west–east (Fig. 4b). Several fractures are open with a width of up to 1 m and a visible depth of maximum 10 m, but there are also some fractures that are only expressed as ground depressions on the surface. The unstable area is highly fractured at its front, as highlighted by the occurrence of many smaller detached blocks (Fig. 4d). Present-day rockfall activity is observed at the front and the toe of the unstable area. A spring is situated along the outcropping thrust plane and, consequently, indicates a layer with high permeability that drains the surface (Fig. 4a, d). This sole detachment may form the lower limit of the unstable area and may also have contributed to destabilize the slope. However, it can probably not act as a basal sliding plane because it is horizontal. Two theories exist about how the instability might

have developed. The first theory is based on the toppling of big rock columns that are separated by steep discontinuities. Flat lying discontinuities, which might be fractures parallel to the main thrust plane, constrain the column height. Another possible failure mode is a planar failure along a complex basal sliding plane that develops by connecting various pre-existing planes with unfavourable orientations by failure of rock bridges. Field observations suggest that the gravitational deformation is propagating towards the south. The back-bounding crevasse shows the largest opening, as well as the largest down-throw, in the northern part of the instability, whereas only a depression is visible as a continuation of it at the southern end of the unstable block (Fig. 4c). Furthermore, a higher rockfall activity at the front of the northern part indicates that the deformation rate seems to be higher in this area (Fig. 4d). However, this needs to be confirmed by monitoring data.

Hyllestad There are two unstable rock slope areas located on the northern coast of the peninsula of Hyllestad in the SW coastal area of Sogn and Fjordane (Fig. 2). Both lie only 1 km from each other and are situated directly above the sea on the edge of a subvertical, 440 m-high cliff that consists of massive Devonian conglomerates (Figs 5 & 6). There were two larger registered rockfalls in 1992 and 1998 with volumes of approximately 30 000 m3, one from the eastern unstable area that is described below, and another one further east along the same fjord at Katlenova, location 14 (Fig. 2 & Table 1). Each

Fig. 5. Hyllestad unstable area, western part. The overhanging unstable block is detached from the cliff by a subvertical, 120 m-long fracture that has an opening of up to 1 m. (a) Sketch cross-section of the unstable area. (b) View to the east. Dashed line A–B marks the location of the cross-section. (c) View to the west.

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Fig. 6. Hyllestad unstable area, eastern part. (a) Sketch cross-section of the unstable area. (b) Overview. Solid white line marks the border of the whole unstable area. Dashed lines indicate open fractures and depressions that separate the area into several blocks. Short black lines indicate the location of the two bolts. Dotted line A–B marks the location of the cross-section. View to the SE. (c) High rockfall activity is present at the front of the cliff. View to the south.

created tsunamis that reached a maximum run-up height of 6 m on the opposite side of the fjord and caused severe damage (Harbitz et al. 2001). The westernmost unstable area forms a massive block of approximately 0.17 Mm3 that is detached from the cliff by a 120 m-long and about 100 m-deep back-bounding fracture (Fig. 5). This subvertical extensional fracture has an opening of up to 1 m. The block is overhanging and shows almost no internal fracturing. Failure would occur when all rock bridges of this major crack have broken, resulting in a free fall. As the volume of this block is much larger than the previously observed events in this area, run-up heights of the triggered tsunami would significantly exceed previous ones. The eastern unstable area is divided into several blocks by a set of cliff-parallel fractures and several fractures perpendicular to the cliff (Fig. 6a, b). Fractures are subvertical with an opening of up to 1 m and visible depths of up to approximately 15 m or are only expressed as depressions in the topography. The lower limit of the unstable area cannot be accurately defined, but fractures on the cliff and a lowermost level of rockfall sources along the cliff suggests a depth of 60–100 m (Fig. 6c). All the blocks together have an approximate maximum volume of 90 000 m3, whereas the largest single block presents a maximum volume of 10 000 m3. The single blocks become smaller towards the rear of the instability. Intense rockfall activity is observed at the front of the cliff (Fig. 6c) and there are some smaller completely detached columns with volumes of up to 200 m3. Two bolts at one of the major cliff-parallel fractures indicate a movement of 2 cm within the last 4 years (Fig. 6b) (Henderson et al. 2008). It is likely that the failure of one of the frontal blocks triggers the

failures of the rear blocks by pressure release and induces a rapid succession of several rockfall events.

Strandanipa This unstable rock slope area is located in the central coastal area of Sogn and Fjordane (Fig. 2). It is situated on top of a 100 m-high cliff 620 m above the fjord on a WNW –ESE trending-slope that has an average gradient of 358 (Fig. 7a). The bedrock consists of mica schist and phyllite, and the foliation is in general dipping NE towards the mountains, but folds are locally observed. Widespread blocky rockfall deposits are present at the base of the cliff, some reaching the fjord. The unstable part of Strandanipa is characterized by an irregular pattern of steep fractures that separate the unstable volume into many small blocks (Fig. 7b). Several fractures are open with widths of up to 0.5 m and visible depths of 5 m, while others are only expressed as topographic depressions on the surface (Fig. 7b). Fractures that are visible on the top surface are not clearly noticeable on the cliff. A well-defined boundary at the bottom of the unstable block that could act as a basal sliding plane could not be observed in the field. Instead the bedrock shows irregular layering and fracturing increasing with depth (Fig. 7c). Below an approximately 10 m-thick upper massive layer the bedrock is highly fractured. Hence, the strength of the highly fractured bedrock is lower and, therefore, probably quite unstable. This suggests, similar to observation of sub-recent rockfall events, that this unstable area might only be affected by successive smaller rockfall or toppling events, and not by large slope failures that could reach the fjord and cause a tsunami.

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Fig. 7. Strandanipa unstable area. (a) Sketch cross-section of the unstable area. (b) Overview. The solid line marks the border of the unstable area, and the dashed lines indicate open fractures and depressions that separate the block into several parts. Dotted line A– B marks the location of the cross-section. View to the south. (c) Exposed cliff on which the unstable part is located. The bedrock is highly fractured and is therefore probably quite unstable below the uppermost approximately 10 m-thick massive layer. View to the NE.

Summary and discussion of controlling factors of rock slope instabilities in Sogn and Fjordane County Rock slope instabilities in Sogn and Fjordane County occur basically in six different rock types, which are: Devonian conglomerate, phyllites, mafic and felsic gneisses, metamorphic quartz monzonite as well as metamorphic quartz diorite. All these rock types have undergone a certain degree of metamorphism and most of them show a distinct foliation. Foliation is one of the most important discontinuity sets regarding the development of rock slope instabilities in western Norway (Saintot et al. 2011). The majority of the possible unstable rock slopes in Sogn and Fjordane (about 46%, Table 1) are situated within relatively weak rocks, mainly intrinsically weak rocks like phyllites, but also in rocks that were prone to an intense weathering like mafic gneisses and thus became weak over time. Among them, we can find the largest unstable volumes like Tussen or the Aurland-Fla˚m area

(previously described by Braathen et al. 2004; Blikra et al. 2006). Also, a large number of rock slope instabilities, about 36%, are within felsic gneisses. Only four localities, accounting for 14% of all instabilities observed so far, are situated within metamorphic quartz monzonite. The localities at Hyllestad and Katlenova form exceptions, as they developed within massive Devonian conglomerate and metamorphic quartz diorite, respectively. The presence of the four types of critical structures, including open fractures at the back of the potentially moving block, a valley-dipping foliation, a weakened plane at the base of the potentially unstable block, as well as existing lateral boundaries of the unstable block, was systematically checked for each of the investigated sites in Sogn and Fjordane County (Table 1). As a result it can be shown that in none of the localities all the critical structures are present. Most sites (79%) show a distinct open fracture at the back of the potentially moving blocks, and those that do not comprise potential rockfall areas with minor involved volumes (,5000 m3). Lateral boundaries, such as

ROCK SLOPE INSTABILITIES IN NORWAY

transfer structures, are only completely defined for 32% of the localities, whereas no lateral borders at all could be found for 39% of the localities. The latter comprise mainly rockfall areas. Structural studies of many high-priority sites, like Tussen or Viddalen, show that the foliation is dipping towards the mountain or oblique to the valley so that its activation as a basal sliding plane is not favoured. From the rock slope instabilities where the orientation of the foliation is known, 13 sites show a foliation that is favourable to the observed initial failing mechanisms. Of these sites eight have a foliation that dips towards the valley and thus might be reactivated as a basal sliding plane. Back-bounding fractures or transfer structures may develop along a vertical or steep foliation, which has been observed at five sites. The limit to the base of the instability and thus the vertical thickness remains undetermined for most sites (75%). It is assumed that this is mainly due to the limited accessibility of the very steep lower parts of the studied rock slope instabilities. A possible basal sliding plane is observed only at six sites, accounting for 25% of all observed instabilities. These are actually sites where the lower parts of the slope are accessible or where the structures are distinct enough to be visible from the helicopter. As movement rates are probably small, it is likely that erosion at the front masks deformations due to the slope instability and that the basal sliding planes are therefore not cropping out. Also, no distinct basal sliding plane might be visible on the face if a complex sliding surface develops by destroying rock bridges and stepping down along foliation or other discontinuities. In any case, for defining the shape of the lower limits of the unstable area, further geophysical investigations or drilling would be required. Lineaments on aerial photographs indicate a connection of some rock slope instabilities with pre-existing faults or regional structures. Agliardi et al. (2001) showed that the structural control of pre-existing brittle structures is one of the most important factor in the development of large rock slope instabilities. Pre-existing fractures may form the lateral limits of an unstable area and backbounding crevasses or transfer structures preferentially develop along them (Mahr 1977; Bachmann et al. 2004). For example at the Tussen rockslide area, the back-bounding graben-like structures developed along pre-existing regional lineaments. In Saintot et al. (2011) the influence of reactivated brittle structures on the slope stabilities in western Norway are described in more detail. A slope profile was drawn for each of the rock slope instability sites using elevation data with a resolution of 25 m. An analysis of the location of the instabilities on the corresponding profile shows

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that 79% of the rock slope instabilities in Sogn and Fjordane County have developed along convex slope breaks, which appear as knick-points in the slope profile or at the top of a slope, that is, at the edges of plateaus (Table 1). Figure 8 illustrates all 22 slope profiles where the slope instability developed at a knick-point aligned at this knickpoint. It is so argued that these instabilities are part of the general processes of denudational slope development that try to even knick-points towards an equilibrium state of the slope (Ahnert 1987; Selby 1993, p. 369). Likewise, Holm et al. (2004) observed a concentration of rockfall along convex slope breaks that are interpreted to be of glacial origin.

Discussion and comparison with other areas in Norway As mentioned in the previous section, the slope instabilities in Sogn and Fjordane County are situated at knick-points in the slope profiles, as well as at the edges of high plateaus (Fig. 9a–c). This is similar to several unstable sites along the Romsdalen or Sunndalen valleys (Braathen et al. 2004; Blikra et al. 2006; Saintot et al. 2008). Counterexamples can be found in other highly susceptible ˚ knes rockslide areas of Norway, for example at the A (Braathen et al. 2004; Blikra et al. 2005; Ganerød et al. 2008; Jaboyedoff et al. 2011), where the slope instabilities are developed directly on steep slopes (Fig. 9d). Braathen et al. (2004) defined three geometrical models for rock slope instabilities in Norway based on pre-avalanche deformation patterns, which are: (1) rockfall areas: (2) rockslide areas: and (3) complex fields. Rockfall areas are located on subvertical slopes, and steep cliffparallel or -oblique extension fractures limit the unstable blocks. Rockslide areas are found on slopes with moderate dips (,468) and are characterized by slope-parallel basal sliding planes that develop along zones of weakness. Complex fields show an intricate deformation pattern with the development of a wide range of different structures. However, all complex fields have a low-angle basal sliding plane in common. These geometrical groups also show the typical location of the instability on the slope as we have discussed above. Rockfall areas and complex fields defined by Braathen et al. (2004) are located at the edge of a plateau and thus agree with our three geomorphological models (a), (b) and (c) in Figure 9, respectively, whereas rockslide areas are located directly on the slope and consequently agree with model (d) in Figure 9. Field investigations have been carried out in several areas of Norway and therefore allow a

¨ HME ET AL. M. BO Elevation

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Location of knick-point and rock-slope instability

Distance

500 m

Fig. 8. All 22 slope profiles of rock slope instabilities in Sogn and Fjordane that developed at a knick-point. Profiles are based on a 25 m DEM. The location of current rock slope instabilities coincides with the location of knick-points. However, the intensity of the knick-points is varying. While most sites are situated at very clear knick-points, others show only small breaks in the slope. This results in a relatively broad distribution of slopes below and above the knick-point. Nevertheless, all slope profiles are located below the straight dashed reference line close to the knick-point.

back-bounding fracture

knick-point in slope profile

back-bounding fracture

knick-point in slope profile

basal sliding plane basal plane

(c)

(a) back-bounding fracture

knick-point in slope profile

back-bounding fracture

basal sliding plane

basal sliding plane

(b)

(d)

Fig. 9. General sketch of types of rock slope instabilities in Norway based on pre-failure geomorphology. Rock slope instabilities in Sogn and Fjordane are located at unstable edges of plateau-like surfaces (a and b), as well as at knick-points of slopes (c) and are not situated directly on a steep slope with constant slope angle (d). The expected types of movement are fall or topple for (a) and slide for (b) –(d).

ROCK SLOPE INSTABILITIES IN NORWAY

comparison of the spatial distribution of current rock slope instabilities in different regions. In the Storfjorden region, Møre and Romsdal County, for example, more than 40 localities display structures, which are critical for the development of large rock slope instabilities (Henderson & Saintot 2011). In several localities in the Storfjorden area that have been identified from aerial photographs all critical factors are present, such as: (1) distinct open fractures at the back of possible moving blocks; (2) foliation dipping towards the valley; (3) a potential basal sliding plane; and (4) lateral free boundaries of the block, such as well-developed transfer structures (Henderson et al. 2006; Henderson & Saintot 2011). In contrast, none of the localities in Sogn and Fjordane displays all the critical structures that are named above. Many of the localities that have been identified from aerial photographs eventually only showed some rockfall activity but no signs of large unstable rock slopes that may develop into rockslides. Furthermore, the majority of the recognized unstable rock slopes in Sogn and Fjordane County are situated within highly weathered and, hence, weak phyllites as well as mafic gneisses, whereas the rock slope instabilities in the Storfjorden region are mainly located within Precambrian gneiss. Henderson & Saintot (2011) argue that a fjord- or valley-dipping foliation orientation is one of the main influencing factors for the development of large rock slope instabilities (.3 Mm3) in the Storfjord region of Møre and Romsdal County. This cannot be demonstrated in the same order of magnitude for Sogn and Fjordane County, as discussed in the previous section. It is shown that there is a discrepancy in the amount of historical and current rock slope instabilities in Møre and Romsdal and in Sogn and Fjordane. Whereas the historical record indicates a much larger extent of rock slope failures in Sogn and Fjordane County, the number of potential current instabilities is greater in Møre and Romsdal. This could be due to the incompleteness of the historical and present database. However, the higher density of large current rock slope instabilities in Møre and Romsdal County may result from a locally high gradient of present-day post-glacial rebound combined with a higher tectonic activity in post-glacial times that has strongly disrupted the slopes (Olesen et al. 2000; Henderson & Saintot 2011). For example, the post-glacial Berill fault, up to now the only described neotectonically active fault in southern Norway, lies in this county (Anda et al. 2002). A large historic rockslide occurred in its vicinity and a complex rock slope instability currently develops in its hanging wall (Anda et al. 2002; Blikra et al. 2006). Finally, many branches of the large Møre– Trondelag Fault

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Complex (MTFC) dissect the northern part of the Møre and Romsdal County and are still active. Redfield & Osmundsen (2009) argued that an earthquake that originated at one of the fault segments of the MTFC triggered the Tjelle rockslide in the eighteenth century. It has been discussed in many studies that active faults strongly influence the development and spatial distribution of gravitational slope instabilities (e.g. Hasegawa 1992; Hermanns et al. 2001; Sauro & Zampieri 2001). Ambrosi & Crosta (2006) showed for a study in the Central Italian Alps that major tectonic features control the strength of the surrounding rock mass and thus support the development of failure zones. However, regional uplift can be the cause for an increased relief and, therefore, may negatively affect the stability of slopes (Martino et al. 2004; Galadini 2006). In Sogn and Fjordane County, the gradient of present-day post-glacial uplift is lower and no active faults have been reported up to now (Olesen et al. 2000; Henderson & Saintot 2011). The differing amounts of current rock slope instabilities in the two counties may be controlled by differences in the effect of permafrost or climatic conditions. However, there is no significant difference in the permafrost-affected areas in both counties. Furthermore, climatic variations within each county are much larger than potential differences between the two counties and thus cannot explain the observed distribution of current rock slope instabilities.

Conclusions From these first results, the most important factor for the development of instabilities in Sogn and Fjordane County appears to be the rock type, as the majority of the current rock slope instabilities are situated within relatively weak rock units, such as phyllites and weathered mafic gneisses. Furthermore, a very important controlling factor for the location of rock slope instabilities in Sogn and Fjordane County is a particular slope profile with knick-points. Instabilities develop preferentially at convex slope breaks and are thus part of the general processes of denudational slope development that try to even out knick-points of a slope profile towards an equilibrium state of the slope. The larger number of potential current rock slope instabilities in Møre and Romsdal may be explained by a locally high gradient of present-day postglacial uplift and the greater tectonic activity that occurred in post-glacial times in this county. Therefore, the location of current rock slope instabilities in Sogn and Fjordane seems to have controlling factors other than those found in the Storfjorden region of Møre and Romsdal County.

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We are grateful to Sogn and Fjordane County for support, as well as to L.H. Blikra and M.-H. Derron for their helpful discussions during our work on this manuscript. Also, thanks go to the editor of this Special Publication, M. Jaboyedoff, for a constructive review and to the two anonymous persons for reviewing this paper. The work described in this paper is partially supported by the Research Council of Norway through the International Centre for Geohazards (ICG). Their support is gratefully acknowledged. This is ICG contribution No. 266.

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Landslides: Investigation and Mitigation. Special Report 247, Transportation Research Board, National Research Council, Washington, DC, 36– 75. Flatøy, F., Barstad, I. & Sorteberg, A. 2008. The GeoExtreme project, Module B: Description of the current climate and investigation of future climate scenarios. In: 33rd International Geological Congress, (33IGC), Oslo, Norway, 6– 14 August, 2008. Furseth, A. 2006. Skredulykker i Norge. Tun Forlag, Oslo [in Norwegian]. Galadini, F. 2006. Quaternary tectonics and large-scale gravitational deformations with evidence of rock-slide displacements in the Central Apennines (central Italy). Geomorphology, 82, 201– 228. Ganerød, G. V., Grøneng, G. et al. 2008. Geological ˚ knes rockslide, western Norway. Enginmodel of the A eering Geology, 102, 1– 18. Grimstad, E. & Nesdal, S. 1990. The Loen rockslides – A historical review. In: Barton, N. & Stephansson, O. (eds) Rock Joints, Proceedings of the International Symposium on Rock Joints, Loen, Norway. A.A. Balkema, Rotterdam, 3– 8. Harbitz, C. B., Domaas, U. & Varlid, E. 2001. Rock slide generated tsunamis – probability and hazard ˚ fjorden, Western Norway. In: Sharma, zoning in A V. M. & Saxena, K. R. (eds) Fjellsprengningsteknikk/Bergmekanikk/Geoteknikk, Oslo 2001. A.A. Balkema, Rotterdam, 27.1– 27.13. Hasegawa, S. 1992. Large-scale rock mass slides along the fault scarp of the Median Tectonic Line in northeastern Shikoku, southwest Japan. In: Bell, D. H. (ed.) Landslides: Proceedings of the 6th International Symposium on Landslides, Christchurch, New Zealand, Volume 1. A.A. Balkema, Rotterdam, 119–125. Henderson, I. H. C. & Saintot, A. 2011. Regional spatial variations in rockslide distribution from structural geology ranking: an example from Storfjorden, western Norway. In: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 79– 95. Henderson, I. H. C., Saintot, A., Bo¨hme, M. & Henriksen, H. 2008. Kartlegging av mulig ustabile fjellpartier, Sogn og Fjordane. Norges geologiske undersøkelse Report 2008.026. Henderson, I. H. C., Saintot, A. & Derron, M.-H. 2006. Structural mapping of potential rockslide sites in the Storfjorden area, western Norway: the influence of bedrock geology on hazard analysis. Norges geologiske undersøkelse Report 2006.052. Hendron, A. J. Jr. & Patton, F. D. 1987. The Vaiont slide – a geotechnical analysis based on new geologic observations of the failure surface. In: Leonards, G. A. (ed.) Dam Failures. Engineering Geology, 24, 475–491. Hermanns, R., Niedermann, S., Villanueva Garcia, A., Gomez, J. S. & Strecker, M. R. 2001. Neotectonics and catastrophic failure of mountain fronts in the southern intra-Andean Puna Plateau, Argentina. Geology, 29, 619 –623. Holm, K., Bovis, M. & Jakob, M. 2004. The landslide response of alpine basins to post-Little Ice Age glacial thinning and retreat in southwestern British Columbia. Geomorphology, 57, 201–216.

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Controlling factors for deep-seated gravitational slope deformation (DSGSD) in the Aosta Valley (NW Alps, Italy) G. MARTINOTTI1, D. GIORDAN2, M. GIARDINO1* & S. RATTO3 1

Dipartimento di Scienze della Terra, Universita` degli Studi di Torino, Torino, Italy

2

Consiglio Nazionale delle Ricerche, Istituto di Ricerca per la Protezione Idrogeologica, Torino, Italy

3

Regione Autonoma Valle d’Aosta, Dipartimento difesa del suolo e risorse idriche, Aosta, Italy *Corresponding author (e-mail: [email protected]) Abstract: Deep-seated gravitational slope deformation (DSGSD) is a common and widespread type of large slope instability in the Alps. In the Aosta Valley region in NW Italy, DSGSDs occupy at least 13.5% of the regional territory. In this study, regional distribution analyses have been coupled with local detailed geological and geomorphological surveys of individual phenomena to detect the controlling factors, deformation processes and evolution stages of DSGSD. Data and maps from field and remote-sensing investigations have been supported by drill data and geomechanical and hydrogeochemical analyses from project studies for hydroelectric plants and tunnels. Several phenomena related to DSGSD have been studied thoroughly: gravity-induced stresses, tectonic– metamorphic setting, morphostructural relations, glacial and periglacial morphodynamics, recent tectonic evolution, hydrogeological conditions and karst phenomena have been generically indicated as controlling factors. In the studied area three of the controlling factors were crucial in differentiating the form and evolution of DSGSDs: deep dissolution, surface tectonics, and tectonostructural setting. They are presented as possible end members of a classification scheme for DSGSDs.

Large portions of mountain relief along the Alpine chain are affected by deep-seated gravitational slope deformation (DSGSD) (Mortara & Sorzana 1987; Ambrosi & Crosta 2006). These huge gravitational phenomena are characterized by their large areal extension (multiple km2), and by distinctive complex geometrical, geomorphological and geomechanical settings (Crosta 1996). Over the past four decades DSGSD have been studied throughout the world (Jahn 1964; Ter-Stepanian 1966; Beck 1968; Nemcock 1972; Mahr 1977; Savage & Varnes 1987; Varnes et al. 1989; Chighira 1992; Sorriso Valvo 1995; Crosta 1996), but their original recognition and first detailed studies (Zischinsky 1966, 1969) were conducted in the European Alps where they are widespread. Historically, there have been two different kinds of approaches to the studies of DSGSD: one based on geomorphological and another based on structural data. By integrating these two approaches it is possible to analyse DSGSD geometry and form, interpret the controlling factors, and model their evolution. Recent papers on DSGSD by French researchers accomplished this comprehensive objective by considering both geomorphological (relief forms, weathering of slope material) and structural– tectonic (large fracture patterns, inherited faults) factors as influential in large slope

instabilities (Bachmann et al. 2004, 2006); they made a significant contribution to both two- and three-dimensional (2D and 3D) modelling of DSGSD (Chemenda et al. 2005; Bois et al. 2008; Bachmann et al. 2009). As stated above, activity of DSGSD implies different elements performing interactively as a complex system. To analyse and interpret a DSGSD, it is necessary to conduct a multidisciplinary research project (e.g. Agliardi et al. 2001) by collecting, organizing and processing a considerable amount of data on different features: † geomorphological data on landslide-related elements and other landforms; † stratigraphical and sedimentological data on surficial deposits; † lithological and geostructural data on bedrock units; † hydrological data on surface waters; † hydrogeological and geochemical data on groundwater; † geotechnical and other monitoring data on deformed materials and related deformational features; † pedological and vegetation data pertaining to land use; † historical data on past instability events.

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 113– 131. DOI: 10.1144/SP351.6 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Different data sources mean different acquisition methods and, possibly, different data formulations. In remote-sensing studies, geometrical data are required for thematic mapping, and numerical data are required for geotechnical in situ analysis. From field surveys, geometrical data are obtained from geological and geomorphological sketch maps. From the microscopic analysis of rock-forming minerals, geometrical, but not mappable, data are derived. Such a complex multidisciplinary research project has to comply with the interpretation needs of different variables and overcome the difficulties of their different formulations. For this reason, all of these aspects have been investigated in the DSGSDs presented in this work, representing a selection of some of the most interesting Aosta Valley large-slope instability phenomena. The Aosta Valley has a great number of types and dimension DSGSD that makes it a great study field in the Western Alps. The Regione Autonoma Valle d’Aosta (‘Aosta Valley Autonomous Region’) is the smallest Italian region, with an area of 3262 km2; it is one of the main Alpine valley systems surrounded by the greatest peaks of the Pennine and Graie Alps. It covers the whole mountain zone of the Dora Baltea hydrographical basin. The elevations range from 400 m a.s.l. (metres above sea level) of the valley mouth into the Po Plain up to 4810 m a.s.l. of Mont Blanc, the highest peak of the European Alps. From a geological point of view, the Aosta Valley belongs to the Western Alps, the axial zone of the Alpine chain where it is recognizable as an imbricated stack of continental crust and oceanic units (Dal Piaz 1992). They are the result of the convergence phenomena between the European and Insubric (‘Adria’) palaeocontinent (Polino et al. 1990). The Alpine tectonometamorphic units of the Aosta Valley (Fig. 1) also show signs of postcollisional tectonic activity and neotectonic fault systems (e.g. Aosta-Ranzola fault: Balle`vre et al. 1986; Bistacchi et al. 2001). Similar to other Alpine regions, all of these geostructural and tectonic characters influence not only the Alpine relief evolution and uplift (Hunziker et al. 1992) but also the mountain slope dynamics (Giardino & Polino 1997). Of all of the morphodynamic agents that have influenced the regional geomorphology, the glaciers are the most important, especially by means of the Pleistocene glacial advance phases. The valley configuration still preserves, at various observation scales, the direct traces of glacial modelling (Carraro & Giardino 2004) both as erosional and as depositional landforms. Glaciation also influences indirectly the actual slope dynamics because of the pressure released by decaying glaciers (Panizza 1974).

The erosive activity of watercourses affected all of the Aosta Valley territory through the progressive deepening of glaciated side valleys, and the erosion of glacial landforms and deposits. Important fluvial deposits and landforms are present on the valley bottoms, as are imposing mixed alluvial and debrisflow fans, products of recurrent meteorological events (e.g. the last major one occurring on 13–16 October 2000: Ratto et al. 2003). Gravity has played a fundamental role in the geomorphic evolution of the Aosta Valley and has operated in tandem with a variety of other exogenic processes in shaping the landscape. Gravitational phenomena on mountain slopes can be differentiated by size and type: from simple shallow landslides (mainly soil slips) to large slope instability phenomena (mainly DSGSDs) (Fig. 2) (Giardino et al. 2004). The latter are widespread in the Aosta Valley, as indicated by the Inventory of Landslides in Italy (IFFI Project: Amanti et al. 2001; Colombo et al. 2005): they cover at least 13.5% (441 km2) of the valley (Ratto et al. 2007). Among the many DSGSD phenomena in the Aosta Valley, four representative case studies (Hoˆne, Villeneuve, Breuil– Cervinia and Quart) have been chosen to highlight possible end members of the different factors that control the structure, form and dynamic evolution of large slope instabilities. Figure 1 shows the locations of studied DSGSD in the Aosta Valley with respect to the main geological units in the region. The selection has been made based on geological and geomorphological characteristics considered useful to determine the timedependent and time-independent factors that controlled their formation and continuing movement. For each DSGSD we collected data and produced geotemathic maps by means of field and remote-sensing investigations. Interpretations have been supported by drill data, and by results from geomechanical and hydrogeochemical analysis carried out during the project of hydroelectric plants and tunnels in the Aosta Valley. The result of our studies have been used in this paper to develop a generic classification scheme for DSGSDs in the Aosta Valley: by condensing a large amount of information from the four selected case studies and by comparing to the results of the regional inventory of landslides (IFFI Project), we propose to use the controlling factors as an approach to DSGSD studies in general, regardless of their geographical location.

Hoˆne DSGSD: the role of pre-existing shear zones in the stability of large slopes The Hoˆne DSGSD is located on the right slope of Champorcher Valley (lower Aosta Valley; Fig. 3),

CONTROLLING FACTORS FOR DSGSD Fig. 1. Geological sketch map of the Aosta Valley (modified from Bonetto & Gianotti 1998) and the locations of the studied DSGSDs (1, Hone; 2, Villeneuve; 3, Cervinia; 4, Quart).

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Fig. 2. DSGSD distribution in the Aosta Valley showing the location of the studied phenomena (1, Hone; 2, Villeneuve; 3, Cervinia; 4, Quart): (a) shows the percentages of landslide types in the Aosta Valley in regard to the total number of landslides; and (b) shows the percentage of landslide type area in regard to total landslide area (modified from Ratto et al. 2007). Only the 8% of phenomena mapped in the IFFI Project (Giardino et al. 2004) are classified as DSGSD, but they occupy more than 74% of the total landslide area (Ratto et al. 2007). The IFFI Project (Amanti et al. 2001) recognized 5218 landslide phenomena (Ratto et al. 2007), corresponding to 580 km2 (17.8% of the Aosta Valley territory).

at elevations of between 500 and 2200 m a.s.l. The area is characterized by erosional features of both glacial and fluvial origin. The middle and lower slopes are crossed by the deeply incised Ayasse Torrent. In the higher part of the area, landforms are mainly the result of shallow-slope instabilities related to gravitational and cryogenic processes. Micaschists and gneisses in the lower part of the valley have been assigned to the Sesia Lanzo complex (Compagnoni et al. 1977; Venturini 1995). Bedrock is fairly competent from a geomechanical point of view, the instability of the Hoˆne slope is conditioned by a moderately dipping, planar shear zone that is many kilometres long.

The Hoˆne DSGSD was documented first in 1996 during studies conducted for a hydroelectric plant. Detailed geological and geomorphological surveys and local geostructural analyses carried out by the authors outlined a large (23 km2) area of gravitational sliding characterized by fragmented rock overlying more stable bedrock (Giordan 2006; Martinotti & Giordan 2008). Structural geological surveys and drilling holes verified the characteristics and role of the planar shear zone in the control of the DSGSD dynamics and evolution. A geological cross-section of the south slope of the Champorcher Valley is shown in Figure 4. The major controlling shear plane dips 258NW: it separates a relatively unjointed gneiss and micaschists

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Fig. 3. Vertical view of the Hoˆne DSGSD (dashed pink line). In the upper-right part of the image the Ayasse Torrent confluences into the Dora Baltea River, close to Hoˆne village. The red line represents the Hoˆne hydroelectric power plant. Line A– A0 is the cross-section profile of Figure 4.

lower unit from a heavily fractured gneiss and micaschists upper unit, locally evolving to rock block slide deposits. The thickness of the shear zone ranges from a few metres to tens of metres. Microscale geostructural analyses conducted on oriented samples from rocks cropping out along the shear zone (Fig. 5) suggest that displacement along the shear zone evolved from a deeper crustal position, with ductile deformation, to a shallow one, involving brittle reactivations. Interpretation of petrographical data from thin sections showed the presence of both cataclasites and ultracataclasites. Cohesive breccias with silicified, epitodized and sericitized bedrock clasts represent the product of a hydrothermal phase in lower greenschist condition. They are associated with foliated cataclasites, overprinting older mylonite kinematic indicators. Discordant

pseudotachilites, associated with non-foliated cataclasites, represent a further shallower brittle deformational phase. Clay-rich fault gouges are the latest product of superficial deformation. The microscale observations have been used together with field evidences of the outcropping deformed bedrock to realize a 3D model of the shear plane (Giordan 2006) whose large extent and lateral continuity has been proved in the Hoˆne slope. Interactions between the complex distribution of geomorphological processes and landforms (first: valley glacial modelling; later: deep slope incision by running water) and the relatively homogeneous structural setting (constant dip of shear plane) indicate that the DSGSD involved local, but coherent, sliding in different areas of the Champorcher Valley (Fig. 5). The main shear plane emerges along adjacent tributary valleys separated by

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Fig. 4. Schematic cross-section of the Hoˆne DSGSD showing the position of the reactivated shear plane.

slightly deformed ridges. The hanging wall is intensely fractured with open cracks near the surface and local water seepage on the slope. When conditions were favourable for initial slope instability, the superficial rock mass moved

down-slope along the shear plane. Later, on the higher part of the slope, the DSGSD evolved as a collapse: very coarse gravitational deposits formed, with a large number of huge towering blocks lying over a structural slope subparallel to

Fig. 5. Three-dimensional reconstruction of the relationships between the Hoˆne DSGSD shear plane and the topography; arrows indicate the outcrops of the shear plane mapped during the study; red shaded areas show the projection of the shear plane. A and B are outcrops where the samples for microstructural analysis were taken.

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the shear plane. Some drill holes penetrated these collapsed deposits, reaching loose and disjointed rocks at depth of tens of metres, below which are more stable rock masses (Martinotti & Giordan 2008).

Villeneuve DSGSD: the role of deep dissolution in slope deformation The DSGSD at Villeneuve affects a north–southelongated ridge on the south slope of the middle Aosta Valley, between the confluence of the

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Cogne and Valsavaranche tributaries valleys. The DSGSD was identified in the late 1980s after geological surveys were conducted for the underground power station of a planned hydroelectric plant below the Mt Poignon ridge (Fig. 6). As shown in the geological map in Figure 7, the ridge is composed of micaschists and gneisses of the crystalline basement (Grand St Bernard unit), and calcshists, prasinites, marbles and carnioles of the Mesozoic Piemonte nappe system. The rocks were complexly folded during two isoclinal phases and an open folding event. The folds are crossed by younger shear planes (Fig. 8).

Fig. 6. Hillshade view of the Villeneuve ridge elongated north to the main Aosta Valley bottom, and separating the Cogne Valley (left ¼ east) and Valsavaranche (right ¼ west). The dotted red shape represents the borders of a collapse zone affecting the slope. The dashed yellow line is the projected hydroelectric penstock down to the Dora Baltea River; the solid red line is the alternative penstock, relocated and realized after DSGSD studies.

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Fig. 7. Geological sketch map of the Villeneuve area showing the cross-sections of Figure 8.

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Fig. 8. Sequence of geological cross-sections of the Villeneuve area with the limits of the collapsed area.

The Villeneuve area has a variety of glacial landforms and deposits at different elevations. Tributary valleys have steep rock walls and have been deeply incised by post-glacial fluvial erosion. The slope affected by the DSGSD heads at Mt Poignon (1500 m a.s.l.) and extends down the valley side to 800 m a.s.l. Detailed geostructural and geomorphological field analysis and mapping identified evidence of an anomalous gravitational instability localized in the central part of the slope. The origin of this feature was established through deep drilling (500 m), which enabled the integration of surficial and underground data. The surficial data indicated the presence along the slope of a zone of disjointed and loose rocks, bordered by diffused rockslides and rockfalls. A depression (area c. 200 000 m2) is located in the central part of the ridge, into which glacial landforms and deposits have subsided (Fig. 9a). The zone is characterized by both active and inactive surficial deformation: open fractures, trenches, antislope scarps and closed depressions. The depression, covered by a dense conifer forest, is peculiar because of the absence of streams or springs. The surrounding slopes are much steeper and entirely stable because the rocks are not jointed (Fig. 9a). Surface data suggest the presence of a sinkhole, a phenomenon normally caused by the dissolution of underlying soluble rocks, such as evaporites. Anhydrite and gypsum, although not present at the surface of the depression, may be present at depth,

near the St Bernard–Piemonte nappe contact (Martinotti 1989). Because the underground power plant was to be built in a possible zone of dissolution, drilling was undertaken to investigate the phenomenon. Of particular importance were the results from the deepest (500 m) drill hole (Martinotti 1989). It penetrated, in sequence: landslide deposits, undisturbed– moderately jointed calcschists with many open fractures, folded marbles and calcschists with extremely bad geomechanical characteristics [rock quality designation (RQD) ¼ 0: Fig. 9c], anhydrite breccias, and joint-free Great St Bernard micaschists (RQD nearly 100). As shown in the cross-section in Figure 9b, deep solution disarticulated the rock mass above the soluble rocks, propagating deformation to the surface. Rock collapse has occurred below the surface depression. This interpretation has been confirmed by geological data provided by two highway tunnels in the lower part of the slope, and by nearby surveys and structural analysis. In summary, the Villeneuve DSGSD is mainly caused by dissolution of deep anhydrite. This phenomenon, first documented in the Western Alps at the end the nineteenth century, was first described as an active process by a Swiss geologist (Schardt 1905, 1906) while studying the Sempione railway tunnel. In the Villeneuve case, most of the dissolution probably occurred at the Last Glacial Maximum when the main valley glacier introduced abundant cold water under pressure to the valley bottom. The poor ion content of glacial water is an

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Fig. 9. (a) Geological and geomorphological map of the central-lower part of the Villeneuve DSGSD; (b) schematic representation of the collapsed area and location of the 500 m-long drill hole; and (c) drill cores of very poor RQD, highly fractured marbles at the contact with anhydrite.

enhancing factor of water ‘aggressiveness’ toward soluble rocks (Alberto et al. 2008). As a first stage, gypsum and anhydrite rocks experienced massive dissolution and consequent volume reduction. Subsequently, by means of sulphate-ion-rich fluids, induced dissolution extended to marbles and calcschists (Alberto et al. 2007). This process is still active, although at a lower rate, as shown by the fresh appearance of surficial deformation and by the high sulphate content of the deep waters sampled in the highway tunnel mentioned earlier.

Breuil – Cervinia: litho-structural conditions on DSGSD in high mountain relief The Breuil –Cervinia DSGSD is located on the west side of the upper part of the Valtournanche, near the local ski resort. It is a large slope instability (8 km2) that has been subject to recent intensive urbanization and is crossed by a pipeline connecting Goillet Lake and the Perreres hydroelectric station (Fig. 10). In comparison to the other case studies presented in this paper, the Breuil –Cervinia DSGSD has the

best-preserved and freshest evidence of gravityinduced slope deformation. The deformation extends from high elevations, near Motte de Plete`Cime Bianche (2900 m a.s.l.), which were deglaciated at the end of the Little Ice Age, to Breuil– Cervinia (1900 m a.s.l.), deglaciated during Late Glacial time (Fig. 10). The deformed area therefore has been affected by different geomorphic processes, with respect to those described in the preceding case studies: here fluvial activity is of lesser importance than glacial activity. The high elevation, hummocky topography and abundance of rock outcrops allow reliable geological mapping of the surface geology. The evolution of the Breuil– Cervinia DSGSD is complex. Deformation is concentrated and more active in the lower and marginal portions of the slope than in higher areas. A 10-year record of deformation of a penstock for conveying water down to the Perreres hydroelectric plant on the SW portion of the DSGSD (Fig. 10 section B) shows mean movement rates of 2–3 cm year21, with displacement vectors parallel to the slope. The bedrock comprises units of the Piemonte Nappe System (Vannay & Alleman 1990): the

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Fig. 10. Panoramic view and geological model of the Cervinia DSGSD. Mont Fort unit is shown in black. Dashed line A is the hydroelectric tunnel and solid line B is the penstock; the circle is the DSGSD–tunnel intersection, where the most damaged sections of the tunnel are located.

Zermatt – Saas Unit (serpentinites, serpentineschists and metabasites, with minor metasediments) and the Combin Zone Unit (sensu stricto) (Fig. 1). The latter consists of two superposed sub-units (Marthaler 1984; Sartori 1987): the Tsate´ Unit (calcshists and prasinites) and the Mont Fort Unit (quartzites, dolomitic marbles, calcareous marbles, pseudocarnioles and, possibly, evaporites). The main schistosity in the high Valtournanche dips by about 108 to the west, which favours the activation of deep-seated slope deformation. The geomorphological and geological analysis indicates that only the Combin Zone units have been involved in the gravitational phenomena; the underlying Zermatt –Saas Unit has not been affected. The lateral variations in thickness, the presence nearby of pseudocarnioles (Alberto et al. 2005, 2008) and the poor strength of the Mont Fort Unit, which is bounded by two large shear zones, show that the deep-seated gravitational phenomena result from the progressive deformation of the Mont Fort Unit (Fig. 10). Either the pseudocarnioles or, possibly, minor evaporite masses along the main basal plane of the DSGSD probably acted as weak and soluble layers, playing an active role in the activation of DSGSD. The overlying rocks of the Tsate` Unit were consequently deformed and displaced by the reactivation of the pre-existing joint systems. Confirmation of this hypothesis came from a geological inspection of the damaged tunnel connecting the Goillet Dam to the Perreres hydroelectric pipe. The tunnel passes through the Zermatt– Saas unit, which is only slightly fractured, and into the Mont Fort Unit, which is highly fractured. Tunnel remedial work was concentrated at the tectonic contact between the two units, where

water seepage and ground deformation were encountered and the tunnel lining was damaged. During tunnel remedial work, clay-rich cataclasite and fault gouge were found along this contact, but not in outcrops outside the tunnel. These observations confirm the role that DSGSD played in deformation of the tunnel. Other indications of large-scale slope instability are visible at the foot of the Breuil– Cervinia DSGSD, where highly fractured and altered rock masses alternate with thick coarse debris units. Here walls of buildings and other structures have been progressively fractured and tilted in a manner consistent with DSGSD. This DSGSD is characterized, in addition to the typical slow and progressive movements, by several rapid episodes of movement localized at the base of the DSGSD. An example is the Motta di Plete` rockslide, with a surface area of about 0.3 km2.

The Quart DSGSD: the dynamic role of fault zones in the stability of large valley slopes The Quart DSGSD extends over several kilometres of the slope on the north side of the broad middle section of the Aosta Valley. The section of the valley owes its morphology, in part, to the Aosta – Ranzola fault zone, one of the more important neotectonic lineaments in the NW Alps that has been active from the Oligocene to the present (Giardino 1995; Bistacchi & Massironi 1999). The geomorphology of the Quart DSGSD (Fig. 11) is a product of several agents: glaciers, responsible for erosional and depositional landforms (roches moutonne´es, spillways, erratics,

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Fig. 11. Evidence of the deep-seated deformations related to the Quart DSGSD are visible at the base of the slope in a hydroelectric tunnel and plant. Three-dimensional view of the Quart DSGSD and principal geomorphological and deformational elements: 1, glacial deposits, distributed at different elevation as a response to the deepening of the Aosta Valley by the Balteo Glacier; 2, main DSGSD down-slope scarp, the upper one being a reactivated tectonic discontinuity of the Aosta –Ranzola system; 3, main DSGSD counterslope scarps; 4, outcrops of bedrock thrusts over glacial deposits; 5, DSGSD displacement vectors; 6, the Vollein area; 7, hydroelectric penstock; 8, hydroelectric power plant. (a) Subvertical attitude of the tectonic discontinuity along the main DSGSD upper scarp. (b) Schematic representation of the kinematic relations between the master fault of the Aosta–Ranzola system (in the lower part of the image) and the Quart DSGSD.

tills), gravity, tectonic processes, and deep dissolution. These agents are responsible for nearsurface deformational features, including trenches, counterslope scarps and closed basins, many of which displace glacial landforms. Pre-Quaternary bedrock comprises rocks of the Piemonte Domain (calcshists with intercalations of mica marbles and metabasites) and the Austroalpine System (orthogneisses, fine-grained gneisses and subordinate amphibolite lenses) (Canepa et al. 1990). The rocks have undergone multiphase isoclinal ductile deformation (phases 1 and 2), overprinted by post-foliation folding (phase 3) with east– west horizontal fold axes. Ductile low-angle shear planes cut the calcschist units; some of these shears have been reactivated and locally involve Quaternary deposits. The reactivated shear planes are possibly phase 3 structures. Rock outcrops along the slope are cut by several high-angle fractures, commonly showing relative

displacements. The two principal groups of discontinuities – both subvertical, striking N58E – N258E and N808E–N1108E – show striations, furrows, rock and calcite steps, and calcite, quartz and other fibrous mineral growths. Kinematic analysis of these high-angle planes suggests the presence of fault surfaces with carbonate or silty in-fillings, showing later striae from mechanical effects. In the eastern area there is field evidence that the last displacement was left-lateral strike–slip. The highangle faults acted as headscarps for gravitational sliding along low-angle shear planes. The evolution of the Quart DSGSD has been controlled by several factors, whose importance and mutual interaction have changed through time. The principal factors are: † neotectonic activity of the Aosta –Ranzola shear zone (Giardino 1995); † deep dissolution of evaporites, confirmed by the high sulphate content of the deep waters sampled

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in the hydroelectric tunnel (Paolina et al. 1986; Alberto et al. 2008); † glacial and fluvial–glacial erosion, which overdeepened the valley and produced steep valley walls, leading to later failure during and after deglaciation; † gravity, which induced both deep-seated and shallow-slope movements. Faults associated with the left-lateral Aosta– Ranzola shear zone control the geometrical pattern and the kinematics of the Quart DSGSD. Although primarily a strike– slip structure, the Aosta – Ranzola master fault has a significant normal component responsible for recent uplift of the south slope of the Aosta Valley. The Quart DSGSD is actively deforming within this conjugate extensional system (Fig. 11b). Ancient valley floors at different elevation testify to the progressive deepening of the Aosta Valley resulting from the Balteo Glacier and the Dora Baltea watercourse during the Quaternary. The study of Quaternary landforms and deposits show that the DSGD was active at least from the Late Pleistocene. The Vollein necropolis, dated to between 3000 and 1550 years BC and located on the lower part of the DSGSD, was itself involved in the gravitational deformation (Giardino et al. 2000). But, from a geological point of view, this gravitational activity represents only a single episode of the long history of DSGSD of the Quart slope (Fig. 12). Deformation of modern buildings and other recent engineered structures, in addition to levelling and GPS measurements, confirm that the DSGDG is still active. Deformation of one important underground structure allowed an estimate to be made of present-day activity. Damage to the concrete lining of the Quart hydroelectric station at the base of the slope (600 m a.s.l.) has been monitored since the start of the operation in 1950. Remedial work has been supported by studies and investigations that have provided a detailed description of the underground phenomena and a quantification of the movements (Paolina et al. 1986; Fornero et al. 1988; Chiesa et al. 1991). Deformation has also been observed in the underground hydroelectric power plant located below the present-day valley floor and in the large feeder tunnel located at the base of the slope. Movement has also been recorded in a drill hole equipped with an inclinometer and extensometer that is located above the power station. The offsets in the concrete have extensional and lateral components, which is in contrast to the main movements parallel to the slope along tectonic features and also parallel to the deep-seated collapse or progressive settling caused by deep dissolution phenomena.

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Precise topographical measurements were made between 1994 and 1998 at four sites located on the west side of the Vollein necropolis, between 650 and 940 m a.s.l. Although the periods of observation are brief, meaningful and coherent movements of 0.5 –0.8 cm have been found on the slope. Figure 12 summarizes the complex evolution of the Quart DSGSD. The present-day slope is the result of a complex interaction between glacial and fluvial processes and gravitational movements, operating over a long period of time. The DSGSD is characterized by an upper extensional zone and a lower compressional one. As the valley was deepened by erosion, displacements on the slope propagated to progressively lower levels, creating complex interference patterns with the local geological structures; former extensional sectors were later subject to compressional forces, and vice versa. At present, extensive portions of the lower slope are subject to gravitational activation along pre-existing tectonic or gravitational discontinuities. Continuing dissolution of gypsum and anhydrite beneath the slope may contribute to this instability.

Controlling factors and a classification of DSGSD In the Alpine geological literature, DSGSD is commonly defined on the basis of geomorphological features and surficial deformation (Mortara & Sorzana 1987; Varnes et al. 1989). Important DSGSD criteria include: † movement over a large area, commonly an entire slope (Agliardi et al. 2001), in some cases on both sides of a drainage divide; † small displacements in comparison to size of the mass, but occurring over thousands of years (Cruden & Hu 1993; Ballantyne 2002); † toe bulging (Crosta 1996); † the presence of large masses of fractured rock; † the lack of a drainage network or, in one is present, it is anomalous with strong structural control; † the presence of trenches, closed and linear depressions, scarps and counterscarps, and double ridges (Crosta 1996; Radbruch-Hall et al. 1997); † small discrete landslides within the area of the DSGSD (Agliardi et al. 2001; Tibaldi et al. 2004). In the past 10 –15 years great progress has been made in defining these geomorphological features and their evolution. These features provide a good starting point for DSGSD classification. Our studies in the Aosta Valley indicate that, in addition to classic geomorphological investigations,

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Fig. 12. Idealized cross-sections of the Quart slope showing the long-term DSGSD evolution related to the river and glacial erosional activity. Glacial deposits of different age involved in the low-angle shear planes at different elevations confirm different stages of activity stages of the DSGSD. The ‘zip’ lines show the changing limit between the extensional and compressional zones of the unstable slope (diagrams modified from Giardino 1995).

a detailed geological and structural study of the associated basement rocks involved in the gravitational deformation is an essential element in understanding the phenomenon. Moreover, classifications

based solely on landforms are inadequate because different sectors of the same DSGSD can have different kinematic behaviour, spanning more than one geomorphically based classification field.

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Another limitation of the geomorphological approach is that DSGSDs are long-lasting phenomena and thus can be affected by a variety of changing phenomena; for example, large-scale climate change, variable tectonic activity, earthquakes and changes in slope morphology. The size of DSGSDs and their persistence in time force the researcher to thoroughly study bedrock, geological structure, slope geomorphology, surficial deposits, hydrogeology and changing land use. By using this approach, one can identify the main factors that control DSGSD evolution. Some attempts at classification have been based on predisposing and activating factors. This exercise has proven to be complicated and unfruitful; it thus has been set aside in preference for a different approach based on ‘controlling factors’, which condition and differentiate the evolution of DSGSD. † Lithotectonic aspects: for example, the involvement of foliated rocks such as phyllites or calcschists that can behave plastically (Zischinsky 1969); and the presence of shear planes, even those unrelated to recent tectonic evolution, that can be reactivated by gravity under suitable geomorphological conditions (Crosta & Zanchi 2000; Ambrosi & Crosta 2006). † Neotectonic history: recent tectonic activity can serve as an effective driver of large slope instability (Carraro et al. 1979; Hippolyte et al. 2006; Agliardi et al. 2009). † Relations between lithotectonic and geomorphological features and slope evolution (Hu¨rlimann et al. 2006). † Deep-seated dissolution, which is widespread and underappreciated in the Alps because outcrops of evaporites are rare (Schardt 1906). In all cases of DSGSD, where deep drilling data are available, evaporitic units have been intercepted at depth even though they do not crop out at the surface. Examples include Crodo and the Simplon Tunnel (Ossola Valley), Villeneuve (Aosta Valley region) and the Vizze Valley (Bolzano Province). Figure 13 shows a possible classification of the DSGSD in the Aosta Valley and their distinctive features based on the above-mentioned factors. Case (a) is a DSGSD that evolves based mainly on movement along pre-existing, mainly slopeparallel, shear planes or other planar tectonic elements. Case (b) involves plastic flow, as described by Zischinsky (1969), where the deformation of the slope is controlled by a fairly homogeneous rock mass under the influence of gravity. Specific lithotypes, such as phyllites, and the absence of a single shear plane are the distinctive features of this class of DSGSD. Case (c) is a

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DSGSD involving deep dissolution and related surface collapse. Case (d) is similar to case (a), but is partly driven by neotectonic activity. Case (e) involves a combination of two or more of the above mechanisms. It is common in the Alpine environment owing to the complex interplay of structural and metamorphic evolution, neotectonic activity and geomorphological processes. It should be emphasized that DSGSD are longlasting phenomena, in which many controlling factors come into play at different times and with different effects on slope evolution. A single factor can, at different times, activate movement or, alternatively, predispose the slope to movement. Recognition that relations between different factors can change over time has led to the possibility of discriminating time-independent factors from time-dependent factors. The main timeindependent factors are: † lithology; † Alpine tectonometamorphic structure, including inactive deformational elements of ductile and brittle style; † gravity; The main time-dependent factors are as follows. † Relations between the lithotectonic framework, geomorphology and the slope evolution. Slope evolution can condition the potential energy of the slope. Valley forms can change, and it is possible that specific lithotectonic factors can switch from passive to active causing the slope to transition from stable to unstable (Tibaldi et al. 2004). For example, erosional deepening of the Ayasse stream has probably played an important active role in activating the Hoˆne’s tectonic shear zone. † On a temporal scale of thousands of years, as is the case with our DSGSD, recent tectonics can be considered a time-dependent variable. There are instances where changes in the local stress field have been caused by the creation of new structures or the reactivation of preexisting ones (Koons 1986; Bistacchi & Massironi 2000). † Glacial and fluvial processes can change with climate, affecting both surface drainage and the groundwater circulation, inducing or accelerating dissolution of evaporites or limestones at depth. The Quart DSGSD is a good example of how complex the evolution of a DSGSD can be. The Aosta – Ranzola shear zone plays an important role in slope deformation at Quart because the basement rocks along the shear zone are highly fractured. Secondary permeability induced by shearing allows surface and near-surface waters to reach evaporites

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Fig. 13. Possible classification scheme for the DSGSD in the Aosta Valley based on their controlling factors: (a) DSGSD characterized by gravity-induced reactivation of pre-existing shear planes; (b) DSGSD with plastic flow, mainly in schists and calcschists bedrocks; (c) DSGSD with collapse and gravity-induced movements related to deep-seated dissolution phenomena; (d) DSGSD with gravity-induced movements mainly caused by the activity of neotectonic structures; and (e) DSGSDs controlled by mixed factors.

layers at depth. These waters have dissolved gypsum and anhydrite, activating upwards collapse that has extended to the surface. Deformation has also propagated downwards because of progressive erosional down-cutting of the Aosta Valley. The

evolution of the Quart DSGSD is therefore dependent on a series of factors: † the regional tectonic structure (Aosta –Ranzola shear zone) and its recent evolution;

CONTROLLING FACTORS FOR DSGSD

† the relation of the slope to exogenous factors such glacial erosion, deglaciation and fluvial deepening of the valley. All of these factors, combined in space and time, and with differing levels of importance, produced the present the Quart slope. DSGSDs commonly have a complex history that makes it difficult to evaluate separately each controlling factor. However, the data we have collected allow us to formulate an interpretative model for each of these factors. The model must take into account that relations among controlling factors may vary through time. The implication is that the representation of DSGSD can not be confined to a simple time, but is better represented as an evolutionary trend, during which the relative importance of the various controlling factors can vary.

Conclusion The paper presents results of studies on DSGSDs in the Aosta Valley, based on detailed geological and geomorphological surveys of four large slope instabilities and on comparison of data at a regional scale. Over 13.5% of the area of the Aosta Valley is affected by many types and sizes of DSGSDs, making the valley an ideal environment for studying this phenomenon in the Western Alps. Four case studies have been selected from the regional database on the basis of their representativeness in terms of size, geological and geomorphological features, state of activity, and related hazards and risks. Data on geomorphology and structural geology are presented and compared for the four selected DSGSDs (Hoˆne, Villeneuve, Breuil – Cervinia and Quart). Special attention was paid to tectonic, hydrogeological, geoengineeering and archeological features where relevant to the interpretation of individual DSGSDs. Data and maps from field and remote-sensing investigations are supported by drill hole data and geomechanical and hydrogeochemical analyses obtained from studies and remedial work carried out the for hydroelectric plants and tunnels. By analysing and comparing the entire dataset of each case study, we identified and interpreted controlling factors on DSGSD activity and evolution. We propose a fivefold classification scheme based on the controlling factors of the observed DSGSD: ‘Tectonic’, ‘Plastic flow’, ‘Dissolution’, ‘Neotectonic’ and ‘Mixed’. The five types of DSGSD effectively integrate the geomorphological research approach with the geological and structural ones. The widespread occurrence of similar features in other areas of the Alps suggests that the ‘controlling factor’ concept is useful for classifying DSGSD outside the Aosta Valley. This concept may help

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in interpreting long-term, complex, large slope instability phenomena by enabling analysis of regional, time-independent factors controlled by structural geology and local time-dependent factors controlled by geomorphological and seismotectonic processes. Moreover, the approach using the controlling factors could also enhance applied studies by facilitating recognition of hazardous processes, interpretation of their causes and the design of targeted remedial measures.

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Palaeostress analysis of a giant Holocene rockslide near Boaco and Santa Lucia (Nicaragua, Central America) ´ 3, ROMAN NOVOTNY ˘ 1,2*, MARKE´TA KERNSTOCKOVA ´ 4, IVO BARON 4 5 4 ´ ˘ ´ DAVID BURIANEK , PETR HRADECKY , PAVEL HAVLICEK & ROSTISLAV MELICHAR3 1

Czech Geological Survey, Leitnerova 22, 658 69 Brno, Czech Republic

2

Present address: Geologische Bundesanstalt, Neulingsgasse 38, 1030 Wien, Austria 3

Department of Geological Sciences, Faculty of Science, Masaryk University, Kotlarska 2, 612 00 Brno, Czech Republic 4

Czech Geological Survey, Leitnerova 22, 658 69 Brno, Czech Republic 5

Czech Geological Survey, Kla´rov 3, 658 69 Praha, Czech Republic *Corresponding author (e-mail: [email protected])

Abstract: A giant rockslide occurred on the southern side of an Upper Tertiary shield volcano in central Nicaragua in the Holocene. The failure caused tectonic-like deformation of rock masses and changed the local stress regime. The lower, compressional part of the rockslide produced a stress field with the axis of maximum stress (s1) parallel to the displacement vector of the main body. The upper part of the rockslide was gravity-driven with s1 vertical and s3 horizontal, and oriented SE– NW. The mass tended to move SE. In the crown, the stress field had a subvertical s1 steeply dipping towards the west. Data at the base of the Santa Lucia Depression, where eastand west-dipping reverse and thrust faults developed, showed that the compressional stress, s1, was nearly horizontal and east– west oriented, the horizontal s2 was north–south oriented, and the s3 was subvertical. These compressional conditions resulted from the collapse of the crown after the main slope failure phase. Simultaneously, along with the gravity relaxation of the main displaced mass, the slopes and mountain slopes along the main scarp depression underwent deep-seated sliding, sagging and flowing.

A giant rockslide occurred on the southern side of an ancient volcano near Boaco (central Nicaragua) in the Holocene. The rockslide was recognized during geological mapping in the area of Boaco and Santa Lucia (Hradecky´ et al. 2007). We used the term ‘giant’ to accentuate its extraordinary dimensions, which are in excess of 10 km3. Such slope failures are called ‘large’ if their volume is more than 1 km3 (Castelli et al. 2009) or classified as ‘extremely large’ (.5 km3) from a landslidemagnitude point of view (Fell 1994). Gravitational failures of stratovolcano edifices are quite common and they tend to be dramatically larger than non-volcanic terrestrial landslides (Reid et al. 2000, Siebert 2002) because the inherent instability of volcanoes is attributable to a large number of factors, for example, earthquakes, elevated fluid pressures, presence of unconsolidated pyroclastic rocks, hydrothermal alteration and dike or magma emplacement, acting individually or complexly (Siebert 2002). Thus, large volcanic landslides represent a major hazard in Central America. The rockslide near Boaco (nomenclature after Cruden & Varnes 1996) caused tectonic-like

deformation of rocks, produced or reactivated faults and changed the local stress field. Gomberg et al. (1995) noted the similarity between faults produced by large slides and those produced by regional tectonic processes. Some authors mentioned regional tectonic stress field as the principal controlling factor for the evolution and spatial distribution of landslides. Tectonic stresses, in particular, play an indirect role in landslide control through their direct impact on patterns of joints (Gupta 2005) and through their major role in shaping the mountain topography (e.g. Ollier 1981; Embleton 1987). The spatial orientation of valleys and drainage networks as a geomorphic factor in controlling landslide distribution is also closely related to tectonic joints (Scheidegger 1980; Abrahams & Flint 1983; Pohn 1983). However, the direct effect of regional tectonic stresses on near-surface stresses, changed by local topography, was given theoretically by Savage & Swolfs (1986). Contrary to the above-mentioned approaches, this paper describes how a giant rockslide can modify the local palaeostress field.

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 133– 145. DOI: 10.1144/SP351.7 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Geological setting of the giant slope failure The study area is located in mountainous terrain of central Nicaragua, northern Central America (Figs 1 & 2), which is a southern part of the Chortis block of the Caribbean plate. The Chortis block is situated in the northern part of Central America (Honduras, Nicaragua, El Salvador, Guatemala); it consists of Grenville –Palaeozoic and Mezosoic basement (phyllite, gneiss and orthogneiss metamorphoses in greenschist –amphibolite facies) partially modified by Tertiary metamorphism and intrusions (Rogers et al. 2007). The southern part of the Chortis block is almost completely covered by Miocene –Pliocene pyroclastic and effusive strata. Large accumulations of mainly Neogene volcanic rocks occur in the Nicaraguan Highlands region around Santa Lucia and Boaco (Weyl 1980). The volcanic material was produced at several volcanic centres. We anticipate that at least seven large volcanoes formed in the study area during the period of main volcanic activity; some of them have since been destroyed by erosion or buried by younger ones, while others are still quite

well preserved, such as the shield volcano complex of Santa Lucia its 4 km-wide caldera in the centre. The approximately 15 km-wide predominantly intermediate –basic volcanic structure of Santa Lucia is characterized by a large volume of pyroclastic rocks and andesitic lava flows that erupted probably during the Neogene. Recently the following five lithological units of volcanic rocks have been recognized in the Santa Lucia volcanic complex. † Acid ignimbrites, accompanied by surges, fine falls and debris flows (in lower position) are grey– brown. Some ignimbrite layers, several metres thick, have been altered (argilitization) and locally strongly affected by brittle deformation. † ‘Lower andesites’ are deeply altered sets of individual lava flows and clinkers mainly of basaltic andesite composition, alternating with layers of agglomerates. They are porous, which permits strong alteration and weathering processes. Chemical weathering (e.g. zeolitization, carbonatization, argilitization) alters primary unstable minerals and glass into a mixture of limonite,

Fig. 1. Digital terrain model (DTM) of Nicaragua, with the location of the study area near Boaco indicated. Co-ordinates for town of Boaco are: 128280 1500 N and 858390 3000 W (source of data: SRTM – USGS/NASA).

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Fig. 2. DTM of the entire rockslide and adjacent area with definition of distinct regions: I.A, zone of compression in the main body; I.B, block part of the slope failure, recently partly in tension; II.A, trench upslope of the main body, partly filled by subsequent deep-seated rockslides; II.B, relict of the original caldera of Santa Lucia; III.A and III.B, crown of the slope failure, now separated into blocks from subsequent very deep-seated rockslides; IV.A, eastern palaeo-dam, IV.B, SE palaeo-dam. The large arrow indicates the displacement vector of the giant rockslide; smaller arrows indicate subsequent deep-seated rockslide displacement vectors, which activated from unloading of the main rockslide crown. The numbers indicate the site location of palaeostress analysis.

chlorites and clay minerals, and makes the rock weak and sensitive to deformation. † The third unit represents mainly deposits of agglomerate pyroclastic flows, recorded by Garayar (1972) as agglomerates of the Coyol group. Total thickness reaches up to 250 m. † Upper andesitic basalts and olivine basalts predominate on plateaus of northern surroundings of the area. Total thickness reaches up to 70 m. Lateritic weathering is well and deeply developed. Younger basaltic flows and dykes are common throughout the area. † Coarse agglomerates dominate in the uppermost position along the rims of the caldera of Santa Lucia. They form imposing coulisses around the caldera. The base of the individual flows displays fine matrix, and reverse grading is common. This epiclastic unit is the youngest one in the area.

Methods The main aim of the geological project was a geohazard susceptibility assessment of the area of Santa

Lucia and the Boaco towns, an area of about 500 km2, which is a much wider area than that of the rockslide itself. When we came across the giant rockslide, we tried to survey it in the limited time and spectrum of methods available as complexly as possible. The main tools of the geomorphic analysis were: † satellite image interpretation: Landsat 7 and Aster; † stereoscopic interpretation of aerial photographs at a scale 1:25 000 and 1:40 000; † GIS-based analyses of a digital terrain model (DTM) of SRTM (Shuttle Radar Topography Mission) with resolution of 90 and 30 m; † field reconnaissance, geomorphic and geological, and engineering –geological mapping; † sedimentological logging, structural–geological measurement, and so on. Geological mapping was provided at scale of 1:50 000 and it was complemented with petrological analyses in the laboratory. For the purpose of dating, several trenches and probes were dug in order to take organic samples related to geodynamic

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events. The organic samples were later dated using a conventional 14C method at the Kyiv Radiocarbon Laboratory (Dr M. Kovalyukh, Ukraine) or by accelerator mass spectrometry (AMS) 14C method in the laboratories of Poznan University (Dr T. Goslar, Poland). The main goal of the palaeostress analysis was to analyse the change in local stress field caused by the rockslide. Faults with slickensides and striations, and with a known sense of displacement, were measured in natural outcrops, road-cuts and other artificial excavations along the rockslide body. For the analysis the data of total 146 faults were grouped to represent different typical parts of the rockslide. The palaeostress analysis was based on the multiple inverse method (Kernstockova 2005, Yamaji et al. 2006), which allows the process fault-slip data to be expressed as 9D vectors (Melichar & Kernstockova, 2008) and numerically separated into different deformational phases. Data were processed using computer program MARK (Kernstockova 2005). The software enables polyphase data analysis to be performed using a new numerical method presented by Kernstockova & Melichar (2008) and Melichar & Kernstockova (2008). The output results are represented by separated groups of slips on fault surfaces and palaeostress inversion may be visualized in contoured density plots of stress inversions for four-faults groups. Maxima of density are the best-estimated directions of the principal stresses s1, s2 and s3, respectively (Kernstockova & Melichar 2008).

Results Geomorphic and geological observations The rockslide body has an elliptic and slightly asymmetric shape with a length of 11 km, maximum width of 8 km, summit altitude of the body of about 940 m a.s.l. (metres above sea level) and summit altitude of the crown of about 1070 m a.s.l. The minimum altitude near the toe is about 190 m a.s.l. The slope failure was first recognized through the analysis of DTMs. We estimate the rockslide’s thickness at 400– 500 m. Thus, the approximate volume of the displaced mass would be at about 28–35 km3. Regarding the main trench dimension, we expect total maximum displacement of the main body to be about 1 km towards SSE, which is relatively short with respect to the mass dimensions (about 10% of the rockslide length). The basal shear zone was probably consequent along less competent tuff rocks related to a much older shield volcano now located below the volcano of Santa Lucia. The tuffs and related ignimbrites crop out along the base of the slope failure in

the NW and west part of the rockslide, as well as in erosive cuts near Boaco. However, most of the rockslide body is composed of andesites, agglomerates and epiclastic rocks. Southeast gently dipping stratification of the underlying older shield volcano was probably the dominant factor that controlled displacement of the rockslide to the SSE. The rockslide consists of different parts (Fig. 2): that is, the zone of compression, the main body, the trench and the crown. The compression zone (I.A) has hummocky relief developed on strongly deformed and weathered andesites (e.g. cuestas, pressure ridges and other elevations, hog-backs, deep erosional, cuts and scarps). At some places, relict basaltic dykes crop out locally. In contrast, the main body (I.B in Fig. 2) consists of relicts of old structural relief on agglomerates and andesites, which are variably inclined and bounded by scarps and rock cliffs. Typical scarps (surfaces of normal faults) run perpendicular to the longer axis of the main body and are inclined towards the trench area. These scarps were created from subsequent gravitational collapse of the main body towards the trench (Figs 3 & 4). The main trench developed at the location of mass detachment SE of the town of Santa Lucia (II.A in Fig. 2). Recently, the trench has been partly filled with deposits from subsequent gravitational collapse of its margins in the form of secondary deep-seated rockslides thicker than 300 m. The rest of the Santa Lucia Depression (the caldera) has a semi-circular shape in ground-plan view (II.B in Fig. 2) and is also affected by gravitational collapse of the depression margins. Relicts of the original volcanic palaeosurface around the margin of the Santa Lucia Depression (III.A and B in Fig. 2) represent the crown of the giant slope failure. The area was strongly affected by unloading, and blocks ranging in size from several hundreds of metres up to a kilometre have been sagging or sliding rotationally towards the depression (Fig. 5). The slopes of the depression have been the subject of other superimposed rather shallow sliding, falling and flowing, posing a great risk to the town of Santa Lucia.

History of the rockslide The history of the rockslide and the subsequent geodynamic events were studied in sedimentary deposits related to valley-blockage, faults, landslides and fluvial activity in the area. No sedimentary evidence of any of large-scale fragmentation of rock related to the failure (no debris fan, etc.) was recognized. The rockslide blocked two valleys, and dam-lakes developed along its eastern and SE margins (IV.A and IV.B in Fig. 2). Only remnants of the lake deposits have been preserved there up to the present because both dam fills have

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Fig. 3. Oblique view and interpreted cross-section along the line X– Y, depicted in Figure 2, with the expected subsidence caldera, the interpreted internal structure of the giant rockslide (dark grey) and the subsequent gravitational collapse of the main trench near Santa Lucia. Two dams are also depicted; one of them was dated.

Fig. 4. Oblique 3D view on an aerial photograph of the rockslide’s NW part and southern part of the Santa Lucia Depression approximately near to the place of the mass detachment. The mass (with the Cerro Grande Mt) has experienced two modes of movement; that is, first, towards the SE during the main mass movement (large arrow and dark dash line) and towards the NW, subsequently into the opened space of the trench (part of the Santa Lucia Depression). This latter motion is proved by secondary deep-seated rockslides (contoured) and normal faults (dot and dash line).

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Fig. 5. Oblique 3D view on an aerial photograph of the northern part of the Santa Lucia Depression. This depression has a complex evolution. Originally, it was an ancient volcanic caldera that was, in its SE part, enlarged as a trench during the main phase of the giant rockslide. Slopes of the depression were later affected by very deep-seated rotational rockslides up to 200– 500 m thick. These slides probably caused thrust and reverse faults (point SL 41) and hummocky relief that dominated in the centre of the depression. For scale: the length of the rotational rockslide left of the Cerro Chiscolapa Mt is about 3 km and its expected thickness is about 500 m.

been intensively eroded. The eastern and larger one comprises a flat area within the original valley, deeply modified by back-erosion in the west (erosion reached the ignimbritic bedrock there). The fill of the preserved part comprises blackbrownish– reddish clays with intercalations of fluvial fine gravels and sands. Proof of the second dam has been preserved in the form of two terracelike bodies 1 km south of Boaco, separated by gullies and back-erosion. They completely cover an area measuring 550 m long by 300 m wide. Their flat surface is related to the reddish and black organic-rich and stratified clayey lacustrine deposit. The clays are intercalated with coarse and partly weathered fluvial gravel and sand close to tributaries of the former lake. The blackish fine organic clay at the base of the lacustrine deposit was determined as 3410 + 35 14C years BP (Poz22073) using the AMS method. This is the approximate age of the main activity phase of the entire giant rockslide. The rockslide has undergone other failure phases since then. As evidenced in several outcrops within its body, it has been affected by several younger faults. Two activity phases of strike-slip faults

changed depositional conditions of colluvial and fluvial sediments on the ground surface within the zone of compression about 2 km SW of Boaco. Two of these phases were determined to be 240 + 30 (Poz-22072) and 115 + 30 14C years BP old (Poz-22071) by dating of the organic matter in the soil sediment at a base of each cycle. The interpretation of the faults is quite difficult: they could be of tectonic origin or, indeed, the result of the giant rockslide reactivation. The main body underwent subsequent gravitational failure in the back-scarp area in the form of secondary, very deep-seated, rockslides that moved towards the trench. One such activity phase was identified within the soil buried under a thrust of andesite debris near the trench base about 1 km SW of Santa Lucia (on the left-hand part of Fig. 4). The buried charcoal is 160 + 30 14C years BP old, as determined by the AMS dating (Poz22074). Several other reactivation phases, however, should also be expected there. The tension conditions caused, in the combination of heavy rainfall and earthquakes, several superimposed shallow landslides and earthflows near the backscarp. These rather shallow landslides could be

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relatively very active and dangerous. One of them was dated to 230 + 50 14C years BP (Ki-14374). However, most of the landslides have not yet been dated.

Tectonic features During field mapping we recognized that the area of Santa Lucia and Boaco is strongly affected by faults and joints. The faults have developed owing to both the regional tectonic stress and the giant rockslide activity. Specific studies on the current regional stress field in the Central Nicaraguan Highlands are not available in the literature. However, as recognized from focal mechanisms of recent earthquakes in the Nicaraguan trough (western Nicaragua), the prevailing regional tectonic stress field about 40 km to the west of the study site has maximum stress, s1, oriented horizontal and NE–SW to north–south, and the minimum stress, s3, oriented horizontal NW–SE to east –west (Caceres et al. 2005). The area of the rockslide has the characteristic occurrence of normal and strike-slip faults related to the regional tectonic processes. In general, NE–SW- and NW– SE-striking faults prevail in the southern and western parts of the study area. A system of distinct circular or radial normal faults is clearly visible on the DTM in the close vicinity of the Santa Lucia volcano (see Fig. 2). We expect that they were formed by caldera subsidence that was centred at the town of Santa Lucia. Folds and faults with striations, significant groove casts, strain shadows and spread-grains developed within the rockslide body. Some reverse and thrust faults occur in the compression zone. Small drag folds with duplexes and tension calcite veins also accompanied the faults occurring there. In addition, normal and strike-slip faults are present in the zone of compression. It was quite difficult to distinguish rockslide-related faults from the tectonic faults as both types occur in the hard rock and are, therefore, very difficult to differentiate by origin. However, younger faults, which developed near the ground surface, are often filled with striated clay or dissipated in the form of dense joint set towards the ground surface. It was also recognized that some older tectonic faults were reactivated during slope failure as documented by: (1) a set of two or more different striation systems on single fault plane at some outcrops; or (2) a cleavage of opaline fill in common older joints. In the crown of the entire rockslide, mainly normal faults were mapped in contrast to the base of the Santa Lucia Depression, where mostly compression structures were mapped. Thrust and reverse faults, attendant drag folds, and, in some cases, small open buckle folds developed in the base of the depression

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resulting from the gravitational collapse of the crown towards the depression. The quality of the ‘slickenside record’ differed depending on the faulted material and on the depth in which the fault originated. Weathered andesites and basalts had the worst slickenside records, as did the faults that originated very close to the ground surface. These faults rather dispersed into joints. However, the faults in tuffs and ignimbrites, as well as faults that developed deeper below the ground surface, had the best slickenside records.

Palaeostress analysis For the palaeostress analysis 146 faults with kinematics markers were measured along all of the rockslide and surrounding area (Fig. 2). Data from close sites were grouped to represent each different part of the rockslide, and to give a good input. Data from outside the rockslide body served as reference data. The compressional part of the main rockslide, the area along the compressional margin of the main rockslide, the upper (block) part of the main rockslide, the crown of the main rockslide and the centre of the Santa Lucia Depression were the regions of interest because a different palaeostress field was expected. Area outside the giant rockslide. Two reference areas were analysed in the area outside of the giant rockslide to determine the regional tectonic stress field before the slope failure as a reference for our study. Only one tectonic phase was identified in the area north (site 1) and west (site 2) of the slope failure within the Oligocene volcanic terrain (Fig. 2). In both cases, the principal stress, s1, was vertical or subvertical, whilst s2 and s3 were subhorizontal and with similar values within this gravity-driven stress field (Fig. 6a). Compressional part of the main rockslide. The compressional part of the rockslide (area I.A on Fig. 2) is represented by a very complicated structural domain including all fault types (e.g. normal, reverse and strike-slip) and rough topography. In the western part of the zone of compression (sites 3 and 4), two separate tectonic episodes can be identified (Fig. 6b). The ‘A’ phase is more clearly developed, with s1 vertical; the s2 and s3 stresses are almost equal and oriented along the horizontal plane. However, the ‘B’ phase has s1 horizontal and NNW–SSE oriented, s2 is horizontal and directed to the ENE –WSW, and s3 is vertical. The ‘A’ phase is clearly gravity-driven, whilst the ‘B’ phase is compressional and accompanied by the evolution of reverse faults. Axis s1 in the ‘B’ phase is parallel to the displacement vector of the main rockslide.

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Fig. 6. Density maxima of the principal stress directions s1, s2, s3 plotted in an equal-area projection obtained from palaeostress analysis of data from (a) the area outside the giant rockslide, (b) the compressional part of the main rockslide and (c) the area along the compressional margin of the main rockslide.

Two palaeostress phases can also be identified in the area south and SW of Boaco City (sites 5–7 in Fig. 2). The ‘A’ phase is clearly gravity-driven, with s1 vertical; s2 and s3 horizontal and more diversified in contrast to sites 3 and 4, that is, s2 is NW–SE oriented and the s3 axis is perpendicular

to it (Fig. 6b). The ‘B’ phase is also very similar to sites 3 and 4. However, the axis of the maximum stress, s1, is inclined 58 –208 towards NW and the s2 axis is inclined about 708 –858 towards SE (Fig. 6b). As at sites 3 and 4, the axis of s1 is parallel to the displacement vector of the rockslide.

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Area along the compressional margin of the main rockslide. The area along the southern margin (sites 8 –10 in Fig. 2) performed different results (Fig. 6c). The s1 axis is subhorizontal and ENE– WSW oriented, which is almost perpendicular to data from other parts of the zone of compression (as presented earlier). The s2 axis is horizontal and directed to the NNW–SSE, whereas s3 is subvertical and steeply inclined towards SE. We expect that this stress field is rotated compared to other areas of compression because of the original local topography (the rockslide filled an ancient valley in this area), and from the large diversity and architecture of competent and weak rocks. Two stress phases were recognized in the area to the SW outside of the rockslide body (Fig. 2, sites 11– 14). The clearest phase (Phase ‘A’, Fig. 6c) had a vertical s1 axis, and s2 and s3 were oriented nearly horizontal with almost equal values. The second phase is represented by a subvertical s1 (538/728), steeply dipping towards the slide mass, a subhorizontal s2 (SW –NE oriented) and a s3 (SE –NW oriented). The s3 probably represents the gravity-driven and slope-failure-related phase (Fig. 6c). Principal stress s1 was vertical, and s2 and s3 were subhorizontal (Fig. 6c), in the area east of

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the slope failure (Fig. 2, sites 15– 18) and clearly represent a gravity-driven phase. Upper block of the main rockslide. Data from this part of the slope failure are quite limited owing to abundant subsequent and relatively shallow landsliding, and the lack of relevant outcrops of fresh rock. However, two tectonic phases were distinguished in the eastern part of the slope failure (Fig. 2, sites 19 and 20). The most evident phase was characterized by a vertical s1, and a north– south-oriented subhorizontal s2 and s3 (Fig. 7a). The second, and less evident tectonic phase, had s1 subvertical and steeply inclined towards east. The s2 was slightly inclined towards the west, and s3 was horizontal and NNE– SSW oriented. The slope failure mass was displaced during this relatively poorly defined, and probably older, ‘B’ palaeostress phase. Only 10 fault surfaces were measured in the western upper part of the rockslide body (Fig. 2, site 21 and 22) owing to the lack of high-quality outcrops. However, two palaeostress phases were interpreted (Fig. 7a). In the both phases s1 was nearly vertical, and s2 and s3 were horizontal. In the first phase s2 was oriented east –west and s3 was oriented north –south. The second phase had s2

Fig. 7. Density maxima of principal stress directions s1, s2, s3 plotted in an equal-area projection obtained from palaeostress analysis of data from (a) the upper block of the main rockslide and (b) the crown of the giant rockslide.

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oriented NNW–SSE and s3 oriented ENE –WSW. Both phases are interpreted to be gravity-driven, most probably related to mass-wasting processes. Crown of the giant rockslide. Two palaeostress phases were recognized in the area of the crown of the giant rockslide (areas IIIA and III.B in Fig. 2, sites 23–25). The first phase has a subvertical s1 steeply dipping towards the west, subhorizontal s2 (east–west oriented), and horizontal north –southoriented s3 (Fig. 7b). These data suggest gravitydriven normal faulting with east –west-striking faults. The poorly defined older phase had s1 horizontal and north–south oriented, s2 subhorizontal and east –west oriented, and the s3 near vertical and steeply inclined towards the west. The ‘A’ phase was probably related to the unloading of the crown area, whilst the ‘B’ phase is very difficult to interpret. We suspect that the ‘B’ phase was related to subsequent deep-seated rockslides along the trench that moved northwards or southwards and influenced site 24, which was already situated in its zone of compression. We obtained surprising results from the base of the Santa Lucia Depression to the north of Santa Lucia town (Fig. 2) at site 26. This area is dominated by hummocky relief of north–south-oriented ridges. Stress s1 is nearly horizontal and east– west oriented, s2 is horizontal, and s3 is subvertically oriented (Fig. 7b). Reverse and thrust faults, both east- and west-dipping, developed under such conditions. These compressional conditions probably resulted from the collapse of the crown (block III.B on the Fig. 2) owing to rapid unloading after the main slope-failure phase.

Discussion The study presents a new point of view on the relationship between extremely large, deep-seated rockslides and the local stress field. As presented in the introduction, several papers have addressed regional and local palaeostress fields as one of principal direct or indirect factors controlling the spatial distribution of landslides (e.g. Scheidegger 1980; Ollier 1981; Abrahams & Flint 1983; Pohn 1983; Embleton 1987; Gupta 2005). However, our results show how a giant rockslide modified the local palaeostress field. The giant rockslide caused local rock deformation very similar to tectonic deformation. As we have no sedimentary evidence of large-scale fragmentation of rock during the failure, we expect that the rockslide movement was rather moderate –slow (velocity class 3–4 of Cruden & Varnes 1996). An extensional regime prevailed in the upper part of the rockslide body and near the head, accompanied by normal and strike-slip

faults. Quasi-intact blocks with original geological structures were preserved in the central part of the rockslide body. Compression with thrusts, reverse and strike-slip faults, and fault-related deposits, occur in the toe portion of the rockslide. The rockslide obscured evidence of older palaeostress regimes recorded in the volcanic rocks of the area. We used the original tectonic stress field outside the giant rockslide as a basis for comparison to data from our palaeostress analysis. In this stress field s1 was vertical –subvertical, whilst s2 and s3 were subhorizontal and with similar directions. This stress field was clearly gravity-driven in areas near the towns Santa Lucia and Boaco, and it supports the hypothesis of a large subsidence caldera near the centre of the Santa Lucia Depression (Fig. 8a). The compressional part of the giant rockslide was represented by two tectonic phases that could be identified; that is, a clearly gravity-driven phase and a compressional phase, with the axis of maximum stress (s1) parallel to the displacement vector of the main rockslide body (Fig. 8b). The gravity phase could be related to either the forming of a large subsidence caldera or to gravity relaxation after the occurrence of the main rockslide displacement (Fig. 8c). The upper part of the rockslide yielded quite limited data owing to extensive subsequent shallow –deep-seated landslides and, thus, a lack of relevant outcrops of fresh rock. However, two tectonic phases were distinguished in the upper part of the rockslide. Both of the phases were gravity-driven with s1 vertical. Both phases are probably related to mass-wasting processes. In the rockslide crown, two palaeostress modes were recognized. The first phase had a subvertical s1 steeply dipping towards the west, which suggests gravity-driven normal faulting with east –west-striking faults related to the unloading of the crown area. Results obtained from the base of the Santa Lucia Depression, where east- or west-dipping reverse and thrust faults developed, show that s1 was nearly horizontal and east –west oriented, the horizontal intermediate stress was north– south oriented and the minimum stress was subvertical. These local compression conditions probably resulted from the subsequent collapse of the crown and margins of the caldera. The displaced mass of the rockslide only slightly influenced the surrounding areas. The area along the southern margin of the rockslide except for the southernmost parts, in general, showed just small changes in the stress field, such as the slight tilting of originally vertical s1 outwards from the rockslide mass. These observations of the rockslide displacement vector and its related stress field completely differ from observations of the tectonic stress field control of landslide displacement in the Himalayas (Gupta 2005). There it was found that the

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Fig. 8. Overview of the palaeostress field registered by our analysis along the giant rockslide near Boaco: (a) original palaeostress outside of the rockslide body: (b) palaeostress field related to the main rockslide phase: (c) pre- or post-failure palaeostress field in the rockslide body. Black ellipse, maximum stress (s1) axis; grey ellipse, intermediate stress (s2) axis; white ellipse, minimum stress (s3) axis plotted in an equal-area projection (white circles). For an explanation of the regions see the Figure 2 caption.

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displacement vector of all landslides was either along the minimum stress axis (s3) or along the intermediate stress axis s2, but not in any case along the maximum stress axis (s1). The vector of displacement was either orthogonal to s1 or oblique to it (Gupta 2005). Interestingly, subsequent deep-seated–shallow landslides commonly occur in the block part of the rockslide, which was in extension, but subsequent landslides are rare in the zone of compression. This could probably be seen as an analogy of this giant mass movement and the tectonic effect on landslides analysed by Gupta (2005). Our analysis of the giant rockslide has helped to define the differences between tectonic and masswasting features. In contrast to endogenous-driven tectonics, for the stress analysis of mass movements it must be possible to distinguish both extensional and compressional features within the body of gravitational slope instabilities. This observation could be very helpful during regional geological studies to help determine what is a slope failure and what is a tectonic feature.

Conclusions A giant rockslide that occurred on the southern side of an ancient volcano near Boaco (central Nicaragua) in the Holocene was the focus of a detailed structural–geological and palaeostress study. We did discover that: † It was possible to define the palaeostress regime of a very large rockslide using techniques normally used for regional tectonic analysis, such as field measurements of slickensides and a palaeostress analysis † Results of the palaeostress analysis were generally consistent with the observed rockslide morphology and known displacement directions † The techniques used in this paper could be applied to other very large landslides where slickenside data are available and where it is important to distinguish ‘slope tectonic’ features from ‘regional tectonic’ features. The field work was carried out under the guidelines of a project of geological mapping of risky processes in the area of Boaco and Santa Lucia, which was provided by the Czech Geological Survey in co-operation with the Instituto Nicaragu¨ense de Estudios Territoriales (INETER) in 2007– 2009 with financial support from the Ministry of the Environment and the Ministry of Foreign Affairs of the Czech Republic. The authors would like to acknowledge field and technical support of our Nicaraguan colleagues from the INETER, especially W. Strauch, A. Alvarez, A. Castellon, M. Echeverry, T. Obando, A. Munoz and G. Chavez, and our Czech colleagues participating on the project.

Special thanks belong to M. Jaboyedoff, the chief editor of this volume, for his recommendations and advice, and N. Harland for reviewing the English. Last, but not least, we would like to express many thanks to J. Coe from the Colorado Geological Survey and another anonymous referee for their fruitful help and advice in improving this paper.

References Abrahams, A. D. & Flint, J. J. 1983. Geological controls on the topological properties of some trellis channel networks. Bulletin of the Geological Society of America, 94, 80– 91. Caceres, D., Monterroso, D. & Tavakoli, B. 2005. Crustal deformation in northern Central America. Tectonophysics, 404, 119– 131. Castelli, M., Scavia, C., Bonnard, C. & Laloui, L. 2009. Mechanics and velocity of large landslides – Preface. Engineering Geology, 109, 1– 4. Cruden, D. M. & Varnes, D. J. 1996. Landslide types and processes. In: Turner, A. K. & Shuster, R. L. (eds) Landslides: Investigation and Mitigation. Special Report 247, Transportation Research Board, National Research Council, Washington, DC, 36– 75. Embleton, C. (ed.) 1987. Neotectonics and morphotectonics. Zeitschrift fur Geomorphologie Supplement, 63, 240. Fell, R. 1994. Landslide risk assessment and acceptable risk. Canadian Geotechnical Journal, 31, 261– 272. Garayar, J. 1972. Geologı´a y depo´sitos de minerales de la regio´n de Chontales y Boaco. Informe, 11. Archivo de INETER, Managua. Gomberg, J., Bodin, P., Savage, W. & Jackson, M. E. 1995. Landslide faults and tectonic faults, analogs?: The Slumgullion earthflow, Colorado. Geology, 23, 41–44. Gupta, V. 2005. The relationship between tectonic stresses, joint patterns and landslides in the higher Indian Himalaya. Journal of Nepal Geological Society, 31, 51–58. Hradecky, P., Baron, I. et al. 2007. Estudio geologico para reconocimiento de riesgos naturales, area de Santa Lucia, Boaco. Project Stage Report, Archive of INETER, Managua. Kernstockova, M. 2005. Palaeostress analysis of fault/slip data sets. Bachelor of Science thesis, Faculty of Science, Masaryk University, Brno [in Czech]. Kernstockova, M. & Melichar, R. 2008. Numerical palaeostress analysis – the limits of automation. In: Poblet, J., Medina, M. G., Pedreira, D. & Ferna´ndez, C. L. (eds) International Meeting of Young Researchers in Structural Geology and Tectonics. YORSGET-08, Oviedo (Spain), 1 –3 July 2008. University of Oviedo, Oviedo, 317 –321. Melichar, R. & Kernstockova, M. 2008. 9D space – the best way to understand palaeostress analysis. In: Poblet, J., Medina, M. G., Pedreira, D. & Ferna´ndez, C. L. (eds) International Meeting of Young Researchers in Structural Geology and Tectonics. YORSGET-08, Oviedo (Spain), 1 –3 July 2008. University of Oviedo, Oviedo, 327 –331.

PALAEOSTRESS OF A GIANT ROCKSLIDE Ollier, C. D. 1981. Tectonics and Landforms. Longman, London. Pohn, H. A. 1983. The relationship of joints and stream drainage in flat lying rock of south central New York and northern Pennsylvania. Zeitschrift fur Geomorphologie N.F., 27, 375 –384. Reid, M. E., Christian, S. B. & Brien, D. L. 2000. Gravitational stability of three-dimensional stratovolcano edifices. Journal of Geophysical Research, 105, 6043–6056. Rogers, R. D., Mann, P. & Emmet, P. A. 2007. Tectonic terranes of the Chortis block based on integration of regional aeromagnetic and geologic data. Geological Society of America, Special Papers, 428, 65–88. Savage, W. Z. & Swolfs, H. S. 1986. Tectonic and gravitational stress in long symmetric ridges and

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valleys. Journal of Geophysical Research, 91, 3677– 3685. Scheidegger, A. E. 1980. The orientation of valley trends in Ontario. Zeitschrift fur Geomorphologie N. F. 24, 19–30. Siebert, L. 2002. Landslides resulting from structural failure of volcanoes. In: Evans, S. G. & De Graf, V. J. (eds) Catastrophic Landslides: Effects, Occurrence, and Mechanisms. Reviews in Engineering Geology, 15, 209–236. Weyl, F., 1980. Geology of Central America. Gebruder Borntrager, Berlin. Yamaji, A., Otsubo, M. & Sato, K. 2006. Palaeostress analysis using the Hough transform for separating stress from heterogeneous fault-slip data. Journal of Structural Geology, 28, 980–990.

Complex landslide behaviour and structural control: ˚ knes rockslide, Norway a three-dimensional conceptual model of A MICHEL JABOYEDOFF1*, THIERRY OPPIKOFER1, MARC-HENRI DERRON1,2, ¨ HME2 & ALINE SAINTOT2 LARS HARALD BLIKRA2,3, MARTINA BO 1

Institute of Geomatics and Analysis of Risk, University of Lausanne, Amphioˆle, 1015 Lausanne, Switzerland

2

Geological Survey of Norway, Leiv Eirikssons vei 39, 7491 Trondheim, Norway 3˚ Aknes/Tafjord Beredskap IKS, Ødega˚rdsvegen 176, 6200 Stranda, Norway *Corresponding author (e-mail: [email protected]) ˚ knes is an active complex large rockslide of approximately 30–40 Mm3 located Abstract: A within the Proterozoic gneisses of western Norway. The observed surface displacements indicate that this rockslide is divided into several blocks moving in different directions at velocities of between 3 and 10 cm year21. Because of regional safety issues and economic interests this rockslide has been extensively monitored since 2004. The understanding of the deformation mechanism is crucial for the implementation of a viable monitoring system. Detailed field investigations and the analysis of a digital elevation model (DEM) indicate that the movements and the block geometry are controlled by the main schistosity (S1) in gneisses, folds, joints and regional faults. Such complex slope deformations use pre-existing structures, but also result in new failure surfaces and deformation zones, like preferential rupture in fold-hinge zones. Our interpretation provides a consistent conceptual three-dimensional (3D) model for the movements measured by various methods that is crucial for numerical stability modelling. In addition, this reinterpretation of the morphology confirms that in the past several rockslides ˚ knes slope. They may be related to scars propagating along the vertical folioccurred from the A ˚ knes slope is presented. ation in folds hinges. Finally, a model of the evolution of the A

The unstable slopes located in the Norwegian Fjord represent a major threat for populations because they can induce devastating tsunamis. Three major events are recorded in the last century in Norway (Blikra et al. 2005). The Storfjorden area in western Norway suffered several of these catastrophes in the past and, therefore, active landslides are an important concern in this region (Blikra et al. ˚ knes rockslide (Figs 1 & 2) is the 2006b). The A most active and thus most likely to provoke a catastrophic event. It is a large complex rockslide of approximately 30 –40 Mm3 (Derron et al. 2005; Ganerød et al. 2008), and is one of the most extensively monitored slope instabilities worldwide as it ˚ knes–Tafjord project that started belongs to the A in 2004. The final goal of the project is to set up an early warning system against rockslide-triggered tsunamis for the whole region (Blikra 2008). The understanding of the rockslide mechanism is crucial implementing a suitable monitoring system (Jaboyedoff et al. 2004; Blikra 2008; Froese et al. 2009). New data, such as airborne and terrestrial laser scanning digital elevation models (ALS-DEM and TLS-DEM), differential GPS and

high-resolution borehole inclinometers, allow the improvement of the conceptual model of this rockslide. In a conceptual view, the main types of rockslide mechanisms were described by Varnes (1978), Hoek & Bray (1981), Hutchinson (1988) and Cruden & Varnes (1996), but they all present idealized models. There are only few conceptual models integrating complex structures (Zaruba & Mencl 1969; Sjo¨berg 2000; Agliardi et al. 2001; Sartori et al. 2003; Eberhardt et al. 2004; Stead et al. 2006; Willenberg et al. 2008; Brideau et al. 2009a; Jaboyedoff et al. 2009). The modelling of rock failure mechanisms is still mainly performed in two dimensions (Eberhardt et al. 2004; Stead et al. 2006) and rarely in three (Ambrosi & Crosta 2006; Brideau & Stead 2010). Nevertheless, the development of a site-specific conceptual instability model is a requirement for reliable numerical modelling. Complex movements in rock slopes may be explained by mechanisms that involve pre-existing structures and the formation of new failure surfaces by breaking rock bridges that generate several

From: Jaboyedoff, M. (ed.) Slope Tectonics. Geological Society, London, Special Publications, 351, 147– 161. DOI: 10.1144/SP351.8 0305-8719/11/$15.00 # The Geological Society of London 2011.

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˚ knes rockslide and its limits on a hillshade model derived from ALS-DEM. The eastern Fig. 1. Location of the A rockslide limit corresponds to fault F1 and the SW limit of F2.

imbricated instabilities (Agliardi et al. 2001) with different movement directions (Brideau et al. ˚ knes rockslide is a good example of 2009). The A a complex slide (Blikra 2008; Ganerød et al. 2008; Kveldsvik et al. 2008). The structures, that is faults, joints and folds, are controlling the slope ˚ knes rockslide movements and the division of the A into several blocks. Even small variations of the main structural features have a strong influence on the slope tectonics and its morphology. The present study explores the impact of preexisting structures and the appraisal of new failure surfaces, and aims to show that complex slope movements need a good conceptual model in order to link observed movements to the model and design the accurate monitoring.

Geological and morphological settings ˚ knes rockslide is located on the NW flank of The A Sunnylvsfjorden, a branch of the Storfjorden in western Norway. During the Caledonian orogeny

(c. 400–500 Ma ago), Proterozoic granitic intrusions were metamorphosed to medium-grained, mica-rich orthogneisses with the creation of a welldeveloped mineral banding (Tveten et al. 1998). These orthogneisses make up the major part of the Storfjorden area. The area suffered several folding phases, which can be observed in several places. As most of the folds are tight– isoclinal, they are only observable ˚ knes area fold interferences in hinge areas. In the A of type 3 (Ramsay 1967) can be observed. Most of the instabilities in the Storfjorden area are caused by dip slopes along the main foliation planes that develop large potential slides (Braathen et al. 2004; Henderson et al. 2006). The topography of the Storfjorden area has been deeply cut by glacial erosion along weaker zones in the bedrock, such as foliation, faults and fractures. ˚ knes site itself has been described and anaThe A lysed in detail by Derron et al. (2005) for the main structural features, and by Ganerød et al. (2008) using geophysical data, structural analysis, displacement data and borehole data. The rockslide ranges

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˚ knes rockslide with predicted limits (from Derron et al. 2005). Three historical rockslide events Fig. 2. View of the A are outlined, including the timing of the events (modified from Kveldsvik et al. 2008).

from the toe at 150 m above sea level (m.a.s.l.) to the back scarp at 1300 m.a.s.l., and the average slope angle is 308–358 (Ganerød et al. 2008). But it is steepened locally with subvertical east –westtrending cliffs ranging from several metres to tens of metres in height. Borehole measurements reveal that the main foliation dips between 278 and 348, with an average value of 328 (Kveldsvik et al. 2006), which indicates that the gneiss foliation makes up the slope surface (Ganerød et al. 2008). This range of dip angles is related to the different

fold phases. The top of the rockslide is limited by a back-scarp zone that is identified by an open crevasse with a width of approximately 1 m in its eastern part to 30 m in its western part. Ganerød et al. (2008) and Kveldsvik et al. (2009a) showed that the landslide body is divided into several blocks by subvertical north–south fractures. From the surface morphology and geophysical data they deduced that the landslide is divided in four main bodies separated by sliding surfaces that overthrust each other and daylight at different

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levels (Blikra 2008). In that context, the water has been demonstrated to be flowing through a complex fractured zone (Lo¨w et al. 2008). The landslide is mainly moving south and downwards towards the fjord. The surface displacements are measured by several methods: GPS, photogrammetry, tachyometers, permanent laser distancemeter, extensometers, ground-based and satellite radar interferometry, and terrestrial laser scanning (Blikra et al. 2006a). These measurements show velocities of the order of a few centimetres per

year and up to 15 cm year21 in the upper western part (Fig. 3) (Eiken 2005; Kveldsvik et al. 2006; Blikra 2008; Ganerød et al. 2008), and confirm the division of the rockslide into several blocks moving in different directions at different velocities (Fig. 3). Three rockslides occurred in the recent past in ˚ knes landslide (Kveldsvik the western part of the A et al. 2008): (1) in the upper part in the second half of the nineteenth century; (2) in the lower part, in 1940; and (3) a 100 000 m3 rockslide in

˚ knes rockslide, displaying the measured displacements (data from Blikra 2008) and the Fig. 3. Hillshade model of the A main structural features. The two lateral faults (F1 and F2) are represented as blue lines, the JNS joint set as white lines and fold axis A2 are shown in green. The displacement vectors reveal south and downwards movement of the upper western part, with up to 10 cm year21, movement towards the SSE in the middle eastern part, and small upwards movements (,2 cm year21) in the toe zone.

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1960 in the mid western part. The latter event avalanched down to the sea (Fig. 2).

more complete structural description see Ganerød et al. (2008).

Methods

Main foliation

The morphological and structural analysis of the ˚ knes rockslide has been performed using an airA borne laser scanning digital elevation model (ALSDEM) acquired in 2004 from using a helicopter with a Fli-Mapw ALS equipment. The flight height was 50–60 m and the density of measurements was 15– 25 points per horizontal m2. The elevation accuracy on rock is 4– 5 cm. The reflections from the vegetation have been removed in order to obtain the ground elevation. A grid with a cell size of 1 m was then produced from the data points. Field surveys identified the principal structures and made structural measurements. Several orthophotographs were used coupled with the ALSDEM in order to make a precise link between field works and the ALS-DEM, and to confirm or discard the interpretations based on the ALS-DEM. The structural analysis based on ALS-DEM was made using the software Coltop3D (Jaboyedoff et al. 2007, 2009). This software allows the visualization and characterization of the structures by an orientation-specific colouring of the DEM. The full orientation of a topographical surface is displayed merging, in one colour, its dip angle and dip direction (Fig. 4). To do so, a lower-hemisphere Schmidt stereonet is combined with a Hue– Saturation–Intensity (HSI) wheel, which provides a unique colour for each spatial orientation. The colours are then attributed to each DEM cell by its spatial orientation of pole in the stereonet. This colour representation of the topography has the advantage of displaying the same slope orientations with the same colours. Assuming that the morphology of rock slopes is shaped by discontinuities, their orientation can be measured on the Coltop3D representation of the DEM. A terrestrial laser scanning DEM (TLS-DEM) was used to identify and characterize the main structures involved in the mass movement. It was also used to monitor the displacements of the upper part of the rockslide (Oppikofer et al. 2008, 2009). The surface displacements of the whole rockslide were extracted from the works of Eiken (2005), Kveldsvik et al. (2006), Ganerød et al. (2008) and Blikra (2008).

The main foliation (S1) is a fabric characterized by a mineralogical layering subparallel to the topography. This inherited structural heterogeneity controls most of the basal sliding surfaces of rockslides in the Storfjorden area (Henderson et al. 2006; Oppikofer & Jaboyedoff 2008; Oppikofer et al. in press). The variations in orientation of S1 induce changes in sliding direction, as can be seen by comparing Figures 3 & 4 (see the next subsection on ‘Fold hinges’). The foliation S1 undulates with a hectometric wavelength similar to large, gentle folds with the axis close to 1208/208. These undulations are termed U in the following section, ‘Fold hinges’. ˚ knes, At Rundefjellet hill, some 2.5 km SW of A the dip-slope along the gneiss foliation forms the topographical surface (Fig. 5). This surface displays a morphology in steps and is probably the basal sliding surface of a former rockslide (Oppikofer & Jaboyedoff 2007; Longva et al. 2009; Oppikofer et al. in press). Large, gentle folds or undulations (U) affect the gneiss foliation (S1) and may have led to changes in the sliding direction of the ancient rockslide. ˚ knes slope, the fold axis of the undulaIn the A tions U is subparallel to the SE-plunging axis of the isoclinal folds. Figure 6 shows that small orientation changes of the foliation direction lead to variations in the potential sliding directions.

Description of the relevant observed structures In this section the main structures that control the slope movements and which are relevant to the present conceptual model are described. For a

Fold hinges The fold hinges (with axis A2: 1208/208; Fig. 6) ˚ knes rockslide (Fig. 3) form ridges crossing the A in the slopes because they are more massive. The main foliation varies considerably in direction, thus avoiding a persistent weakness orientation. The S1 orientation changes direction over a short distance and is, therefore, is less prone to slip; hence, there is no possibility for the foliation to act as a sliding surface over long distances. The spacing between two successive ridges is fairly regular (approximately 250 m) and, therefore, reflects the wavelength of the folds A2. The folds A2 are upright or steeply inclined (Fig. 7). However, in these hinges zones the foliation orientation (S1) is very variable, including some locations where S1 is subvertical (Fig. 7b). Such zones are more prone to developing tension cracks or back scars than areas where sliding along the foliation is possible, provided that the foliation is roughly slope-parallel and fjord-dipping (Figs 7 & 8). The main back ˚ knes has a graben-like scarp on the western side of A geometry. It follows a fold hinge with an axis A2

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Fig. 4. Orientation-specific colouring of the