A conceptual model of Hiorthfjellet rock glacier, Svalbard

A conceptual model of Hiorthfjellet rock glacier, Svalbard Permafrost, Phillips, Springman & Arenson (eds) 2003 Swets & Zeitlinger, Lisse, ISBN 90 5809 582 7 R.S. Ødegård Gjøvik University College, Gjøvik, Norway K. Isaksen Norwegian Meteorological Institute, Oslo, Norway T. Eiken and J.L. Sollid Department of Physical Geography, University of Oslo, Oslo, Norway ABSTRACT: Field studies of the Hiorthfjellet rock glacier (78 15 N, 15 47 E) near Longyearbyen, Svalbard, started with surface velocity measurements in 1994. This paper presents a conceptual model for the development of this rock glacier. There is particular focus on the complexity of the front processes. It is suggested that the surface slope is mainly controlled by the accumulation of debris at the base of the active creeping part close to the front, as the rock glacier advances. This process will depend on the flow regime, and could force the front to move upward in a more mature state, assuming a strong longitudinal stress coupling. The suggested conceptual model is discussed related to the overall morphology and dynamics of the rock glacier. The proposed model is consistent with available field measurements. Considerable additional field data will be needed to validate the model. 1 INTRODUCTION The Svalbard archipelago (74 N to 81 N, Figure 1) has continuous permafrost with permafrost thickness varying from less than 150 m near sea level to more than 450 m in mountain areas (Liestøl 1977). MAAT at Svalbard Airport, Longyarbyen, is 6.7 C for the period 1961 90 (28 m a.s.l.). During the first part of the 20th century several interpretations existed on the origin of these features in Svalbard. Liestøl (1962) was the first to link Svalbard rock glaciers to slope processes. Talus terraces in Svalbard were described as having a typical flow structure some what like rock glaciers. Liestøl investigated a rock glacier in Longyearbyen in 1954 where a slip had occurred. He found irregular layers of sharp-edged stones and gravel alternated with layers of dirty ice. This is one of the few known descriptions of the inner structure of Svalbard rock glaciers. His interpretation was that rockslide material accumulates on snow in early spring/summer and gradually transforms into ice during summer. Sollid & Sørbel (1992) suggest Holocene or even pre Holocene age of some rock glaciers. André (1994) states that rock glacier formation started 3500 BP based on dating of surface material with lichenometry. The surface at the top front of Hiorthfjellet rock glacier was roughly dated to 4000 BP based on the length/surface velocity ratio (LSVR) (Isaksen et al., 2000). Due to rock glacier dynamics and long-term initial development the rock glacier probably started to develop at the onset of Holocene. The Hiorthfjellet rock glacier is 400 m long with an estimated mean thickness of approximately 35 m (Figure 2). The first interpretations of field data from the geodetic survey, DC resistivity measurements and GPR sounding are presented in Isaksen et al. (2000). Normalized horizontal surface velocities were measured to 9.0 10.3 cm/year (Isaksen et al., 2000). This paper contains further interpretation of field data and a conceptual model is suggested for the dynamics at the front of Hiorthfjellet rock glacier. This is a contribution to understanding the long debated morphology of these features in Svalbard and to the discussion of the dynamics of rock glaciers in general. Figure 1. Key map showing the location of the field site. 2 FIELD DATA AND INTERPRETATION The surface slope is decreasing from approx. 20 at the top to 10 12 near the front. On the top front of the rock glacier 20 40 m up-glacier from the steep 839

glacier from what is interpreted as the accumulation area. There is no significant trend in the gradient of the reflection layers relative to the surface gradient, however large anomalies are detected, believed to originate from the actual accumulation process (Figure 3). This implies that the layers, interpreted as rockslide deposits, reinforce the upper supersaturated part. If the layering of the Hiorthfjellet rock glacier are similar to the observations by Liestøl in 1954 on the other side of the fjord (Liestøl, 1962), the layers of rockslides are saturated with ice alternating with layers of ice. The radar data give few indications of the thickness of the rock glacier. Rough estimates of the thickness can be made assuming a basal shear stress of 100 kpa and an ice content of 75% giving an average thickness of the supersaturated layer of 35 m. This calculation is based on a strong longitudinal stress coupling with an average surface gradient of 16.6. frontal slope the surface gradient is again increasing to a maximum of 14 15. The average longitudinal surface gradient is 16.6 calculated from the lower accumulation area to the top front. There are two zones on the frontal slope where the slope gradients are 40 or steeper. The upper zone is extending down to 16 22 m below the top front of the rock glacier. The second zone is at the foot of the frontal slope. The GPR profile presented in Figure 3 was obtained using a Pulse EKKO 100 (Sensors and Software Inc., Mississauga, Canada) with an antenna frequency of 50 Mhz. The profile follows the centre flow line. The georadar measurements give no indication of any cumulative shear in the upper 20 m in the area down 3 A CONCEPTUAL MODEL FOR THE FRONT PROCESSES Heaberli (1985) addressed processes at the front of a rock glacier, stating that coarse debris, which falls off the top of active rockglacier fronts, is subsequently overridden, and may then form a relatively stiff basal layer. The hypothesis was later confirmed by deformation measurements in boreholes in the Eastern Alps (e.g. Heaberli et al., 1998, Hoelzle et al., 1998). The Figure 2. The Hiorthfjellet rock glacier is in middle of the photo with the steep front free of snow (June 1998). Figure 3. Radar profile and interpretation. In the lower plot the white dots are depths down to 9 m, grey dots 10 15 m depth and black dots 16 20 m depth. 840

borehole measurements at rock glacier Murtel Corvatsch show that the basal debris layer is about 20 m thick. This mechanism has been conceptually illustrated by Haeberli et al. (1998). In a model of pure plastic flow this mechanism will be less efficient, because the front of the glacier is simply pushed forward, and there will be no gradual increase in surface gradient at the front. The details on how a permafrost ice-rock mixture deforms is far from understood (Haeberli, 2000). However, all deformation measurements in boreholes (Haeberli et al., 1998, Hoelzle et al., 1998) and geomorphological indications strongly suggest that a pure plastic model is not applicable. This assumption is made for the Hiorthfjellet rock glacier. There are no direct field measurements of strain at depth. The steep front of a rock glacier will cause a stress field that can be modelled based on well known methods devoloped for finite slopes. The stress distribution beneath a finite slope is more complex than below an infinite slope. There are no convincing field data from the Hiorthfjellet rock glacier to make definite statements about how the stress field relates to the deformation at the front. A relative increase of the surface gradient on top of the rock glacier close to the front is consistent with an increasing submerging (vertical) velocity. Similarly, the bulge at the base of the steep frontal slope can be explained by shear strain close to the base of the basal debris layer, but other explanations like active layer block creep and a different angle of repose could be suggested as well. In the suggested model it is assumed that the stress field from the steep front can be neglected for the upper supersaturated part of the rock glacier. On the other hand, compression and damped creep close to the front in the basal debris layer could be significant for the overall morphology and dynamics of Hiorthfjellet rock glacier. Sediments in front of the rock glacier consist of lichen covered blocks interpreted as early or mid Holocene dirty snow avalanche deposits from the western flank of the Hiorthfjellet rockwall. Even close to the base of the steep rock glacier front there are very few blocks originating from the active rock glacier. This is regarded as good geomorphological evidence that very few blocks leave the front slope. The reason is probably that there are few large blocks exposed on the steep frontal slope. Debris is expected to melt out in the upper area of the front slope, and will be exposed to gravitational processes. After rearrangements in the active layer these sediments will be exposed to percolating water that refreezes in the pores. It is not known if these processes cause ice saturation or not. If air voids exist as the debris is overridden by the advancing rock glacier, the basal debris might compress due to increased normal stress. Shearing of coarse material implies movements normal to the plane of shear, which causes volume expansion (dilatancy) (Arenson and Springman, 2000). There is no reason to expect that pore water pressure influences the mechanical properties of the basal debris (permafrost). The mechanical properties of coarse grained sediments with ice content at saturation level or below are known from geotechnical investigations. Pore ice will deform even at very low stresses. The long term strength of these sediments is mainly controlled by inter-grain friction like non-permafrost sediments. Ice will add strength to cold permafrost consisting of coarse grained sediments, but for simplicity the added strength is not considered in the conceptual model. Sayles (1973) suggested that for noncohesive materials in the unfrozen state, the long term strength in the frozen state would be roughly equal to that measured in triaxial tests on the freely drained unfrozen material. In the following it is assumed that the internal shearing resistance of the basal debris is constant. This is certainly not valid when the normal stress increases as the material is overridden by the rock glacier. The assumption should be refined in future quantitative models, but the simplification is probably not critical considering the precision in the calculations. In a conceptual model of the rock glacier front processes the following assumptions are made: The flow regime at the front is mainly controlled by the applied stress from the creeping rock glacier. The strain caused by the steep front is neglected for the supersaturated part of the rock glacier, but is considered for the basal debris layer. The flow of the rock glacier is caused by shear strain in supersaturated sediments. Debris that melts out from the upper front area accumulates further down the front slope- not in front of the rock glacier. Zero pore water pressure in basal debris (pores filled with ice or partly filled with ice in coarse debris with permafrost) The sediments accumulating at the front are noncohesive with constant angle of shearing resistance (does not depend on the applied stress). In the initial stage of rock glacier development the front will reach the minimum angle of shearing resistance. Because of the loose state of debris in the active layer at the front of a rock glacier it is reasonable to assume that the front of an active rock glacier will be close to the minimum angle of shearing resistance. Rapp (1960) and Chandler (1973) investigated the inclination of talus and rock glacier slopes in Spitsbergen. Chandler (1973) suggested a critical value for angular rockfill material to be 39 40. When a critical angle is reached the front of the rock glacier will adjust to this angle. This will cause melting of the supersaturated upper part of the rock glacier because 841

debris will be removed from the upper front and accumulate further down the slope. When the normal stress is low the debris that accumulates on the lower frontal slope will be pushed forward by the rock glacier, and shearing will occur. As the front of the rock glacier progress, the normal stress on the debris will increase, which might cause compression and dilatancy. The shear stress at the base relative to the normal stress will eventually reach a critical level known as the coefficient of friction. If the coefficient is assumed to be 0.8, corresponding to an angle shearing resistance of 40 and an applied sharing force of 100 kpa, the critical normal stress will be 125 kpa. An ice-rock mixture with an ice-content of 75% will have a typical density of 1350 kg/m 3, which means that the critical normal stress will be obtained with an ice/rock overburden of approx. 10 m. When the critical normal stress is reached debris will start to accumulate at the base of the rock glacier under the frontal slope (see sketch in upper part of Figure 4). After the initial stage the rock glacier develops as a two-layered system. Several scenarios can be suggested for the further development. In the lower part of Figure 4 three scenarios are suggested. (1) The rock glacier base progresses in a direction controlled by the initial angle between bed slope and the base of the rock glacier. The initial accumulation of debris will cause vertical strain on the supersaturated layer, which means that the rock glacier will move forward at an angle slightly less than the bed slope. If this initial movement progresses, the accumulation of debris will get thicker and thicker as the accumulation of debris at the base is no longer directly controlled by the bed slope gradient. (2) If the process described under (1) is kind of a self repeating mechanism (substituting the original bed with the accumulated debris layer), there might be a tendency of increased accumulation of debris at the base of the rock glacier causing vertical strain on the advancing rock glacier after the (1) (2) (3) Figure 4. The upper figure is a sketch of the initial stage of debris accumulation. Lower part see text for details. initial stage. This process could force the front to move upward in a mature state. (3) The basal debris layer will increase in thickness in the initial stage until equilibrium is reached between the mechanism described under (2) and compression and damped creep of the overridden material. The latter mechanisms will probably be more significant when the basal debris layer gets thicker. The important parameters of these processes are the normal stress and shear stress near the front and mechanical properties of sediments accumulating at the front (partly controlled by permafrost temperature). The details of these processes cannot be assessed with the present knowledge of rock glacier rheology. The complexity of these thermo-mechanical processes should not be underestimated. 4 THE DEVELOPMENT OF HIORTHFJELLET ROCK GLACIER From the georadar profile it is apparent that layers interpreted at rockslide deposits experience rotation and flexure in the accumulation area (Figure 3). Apart from the area close to the rockwall, limited shear strain seems to occur. However, even limited shear strain could change the mechanical properties of the original sediments. The differences in apparent resistivity values from the upper to the lower parts (Isaksen et al., 2000) could be explained by limited shear, bringing the individual particles in closer contact. Work hardening will eventually strengthen the material resulting in practically no shear strain in the upper part of the supersaturated layer a distance down glacier from the accumulation area. The model consists of a layer of a few meters below the upper rigid part with more randomly distributed blocks in a matrix of ice, which might be modelled according to the flow law of ice (Figure 5). This is consistent with borehole measurements in the Alps (Haeberli et al., 1998). The basal layer of the supersaturated rock glacier originates from the upper accumulation area. More randomly distributed blocks could be due to sorting in the accumulation process, or shear strain in the accumulation area could disturb the original layering. A rigid ice/rock mixture in the upper supersaturated layer implies a strong longitudinal stress coupling, which significant implications for the overall dynamics. The rigid layer will resist longitudinal compression. Maximum surface shear strain is calculated to 0.7 10 4 a 1 in the mid-zone based on surface velocity measurements published by Isaksen et al. (2000). The measured surface strain indicates a cumulative thickening of the mid area on the order of a few meters. This zone probably also have transverse 842

Figure 5. A conceptual model of of Hiorthfjellet rock glacier. expansion. In the top frontal area of the rock glacier no significant surface strain are measured. A small positive strain might occur near the front as a response to the stress field of the steep front. The longitudinal stress coupling is significant in the discussion on how the rock glacier responds to changes in basal friction as the length of the rock glacier increases. Assuming scenario 2 (Figure 4) a lowering of the basal gradient as the rock glacier progress will reduce the overall basal shear stress. If accumulation input was constant the dynamic response would be a gradually thicker rock glacier from the front upwards to maintain the basal shear stress. Based on this conceptual model the longitudinal surface profile of the Hiorthfjellet rock glacier is interpreted as a combination of increasing thickness of bed debris, a small longitudinal compression in the central zone and probably a transverse expansion in the mid and front zone and a dynamic feedback to maintain the overall basal shear stress. The suggested conceptual model is consistent with all available field data from the Hiorthfjellet rock glacier. Additional field data will however be needed to actually validate the model. 5 DISCUSSION AND CONCLUSIONS Accumulation of debris at the base of the front slope of a rock glacier can simply be explained by conservation of debris mass and shear strain in sediments supersaturated with ice. As long as the debris accumulates on the front slope, and not in front of the rock glacier, rock glaciers will advance as a two-layered system. The progress of the front will be controlled by the extent of this debris layer. The process is discussed by Barsch (1996) and Kääb et al. (1998) with respect to estimates of rock glacier age. At Hiorthfjellet rock glacier the present front advance rate is probably about 2 3 cm pr. year based on a simple 2D calculation using a thickness of 35 m of the supersaturated layer at the front (75% ice content). The suggested conceptual model opens the possibility that the basal debris layer gets progressively thicker towards the front. If the model is generally applicable in Svalbard, age estimates will be even more sensitive to the exact extent of the basal debris layer. Most of the Svalbard rock glaciers exhibit a decreasing longitudinal surface gradient towards the front. There are several examples of an inverted gradient close to the steep front slope. There are at least four suggestions on the origin of these depressions : 1. Liestøl (1962) stresses the protalus rampart process as most important (in situ accumulation). He does not rule out deformation due to high ice content. 2. Swett et al. (1980) suggests a rotational movement like a cirque glacier causing the depressions. 3. Humlum (1982) points out some shortcomings in Swett s argumentation, and argues that the inner depressions are mainly a degradation form. 4. Berthling et al. (1998) states that these depressions are most likely flow-related features related to compressive and extending flow. The processes described by Liestøl (1962) and Humlum (1982) have been verified by field observations. There are however several examples of Svalbard 843

rock glaciers where these processes are unlikely to control the overall surface morphology. This is discussed by Berthling et al. (1998), who argue that the depressions are caused by extending and compressive flow, adapting a conceptual model suggested by Haeberli and Vonder Mühll (1996) for the Alps. The magnitude of measured surface strain is not convincing with respect to the importance of this process for the Hiorthfjellet rock glacier. The overall extending-compressive flow regime could explain a cumulative thickening of the supersaturated part. However, a cumulative thickening reduces the shear stress, resulting in a dynamic feedback causing the rock glacier to grow thicker in the accumulation area. A hypothesis is therefore launched stating that the low and sometimes inverted surface gradients towards the front of the Svalbard rock glaciers could be controlled by an increasing accumulation of debris at the base and a strong longitudinal stress coupling. The dynamic feedback will have the opposite effect, causing a steeper longitudinal gradient and a gradually thinner supersaturated layer towards the front. Considering the previous suggestions on the origin of this particular surface morphology, an increasing thickness of the basal layer could explain that the front moves upward. The hypothesis remains to be tested with field data of surface velocities near the front. ACKNOWLEDGEMENTS The University Courses on Svalbard (UNIS) and the Department of Physical Geography, University of Oslo, supported this study. Two anonymous reviewers and the editor gave constructive comments. The authors extend their thanks to the persons and institutions mentioned. REFERENCES André, M.F. 1994. Rock glaciers in Svalbard: tentative dating and inferred long-term velocities. Geografiska Annaler, 76A (4): 235 245. Arenson, L. and Springman, S. 2000. Slope stability and related problems of Alpine permafrost. Int. Workshop Permafrost Engineering. Longyearbyen, Svalbard, June 2000. 183 196. Barsch, D. 1996. Rockglaciers. Indicators for the Present and Former Geoecology in High Mountain Environments. Springer-Verlag Berlin Heidelberg. Berthling, I., Etzelmüller, B., Eiken, T. and Sollid, J.L. 1998. Rock glaciers on Prins Karls Forland, Svalbard. (I): internal structure, flow velocity and morphology. Permafrost and Periglacial Processes, 9: 135 145. Chandler, R.J. 1973. The inclination of talus, arctic talus terraces and other slopes composed of granular materials. The Journal of Geology, 81: 1 14. Haeberli, W. 1985. Creep of mountain permafrost: internal structure and flow of Alpine rock glaciers. Versuchsanstalt für Wasserbau, Hydrologie und Glaziologie, der Eidgenössischen Technischen Hochschule Zürich, Mitteilungen 77. Haeberli, W. and Vonder Mühll, D. 1996. On the characteristics and possible origins of ice in rock glacier permafrost. Zeitschrift für Geomorphologie, Supplementband, 104: 43 57. Haeberli, W., Hoelzle, M., Kääb, A., Keller, F., Vonder Mühll, D. and Wagner, S. 1998. Ten years after drilling through the permafrost of the active rock glacir Murtèl, Easterm Swiss Alps: Answered questions and new perspectives. In: Proceeding 7th International Permafrost Conference, Yellowknife, Canada. 403 410. Haeberli, W. 2000. Modern Research Perspectives Relating to Permafrost Creep and Rock Glaciers: A Discussion. Permafrost and Periglacial Processes, 11: 290 293. Hoelzle, M., Wagner, S., Kääb, A., Vonder Mühll, D. 1998. Surface movement and internal deformation of icerock mixtures within rock glaciers at Pontresina- Schafberg, Upper Engadin, Switzerland. In: Proceeding 7th International Permafrost Conference, Yellowknife, Canada. 465 471. Humlum, O. 1982. Rock glaciers in northern Spitsbergen: a discussion. Journal of Geology, 90: 214 218. Isaksen, K., Ødegård, R.S., Eiken, T. and Sollid, J.L. 2000. Composition, Flow and Development of Two Tongue- Shaped Rock Glaciers in the Permafrost of Svalbard. Permafrost and Periglacial Processes, 11: 241 257. Kääb, A., Haeberli, W. and Gudmundsson, G.H. 1998. Analysing the creep of mountain permafrost using high precision aerial photogrammetry: 25 years of monitoring Gruben Rock Glacier, Swiss Alps. Permafrost and Periglacial Processes, 8: 409 426. Liestøl, O. 1962. Talus terraces in Arctic regions. Norsk Polarinstitutts Årbok 1961: 102 105. Liestøl, O. 1977. Pingos, springs and permafrost in Spitsbergen. Norsk Polarinstitutts Årbok 1975, 7 29. Rapp, A. 1960. Talus slopes and mountain walls at Tempelfjorden, Spitsbergen. Norsk Polarinstitutts skrifter. 119. Sayles, F. H. 1973. Triaxial and creep tests on frozen Ottawa Sand. In: Proceeding 2nd International Conference on Permafrost, Yakutsk, USSR, North American Contribution. 384 391. Sollid, J. L. and Sørbel, L. 1992. Rock glaciers in Svalbard and Norway. Permafrost and Periglacial Processes, 3, 215 220. Swett, K., Hambrey, M. J. and Johnson, D. B. 1980. Rock glacier in northern Spitsbergen. Journal of Geology, 88: 475 482. 844