Permafrost creep and rock glacier dynamics

PERMAFROST AND PERIGLACIAL PROCESSES
Permafrost and Periglac. Process. 17: 189–214 (2006)
Published online in Wiley InterScience
(www.interscience.wiley.com) DOI: 10.1002/ppp.561
Permafrost Creep and Rock Glacier Dynamicsz
Wilfried Haeberli,1* Bernard Hallet,2 Lukas Arenson,3 Roger Elconin,4 Ole Humlum,5 Andreas Kääb,1,y
Viktor Kaufmann,6 Branko Ladanyi,7 Norikazu Matsuoka,8 Sarah Springman9 and Daniel Vonder Mühll10
1
Department of Geography, University of Zurich, Zurich, Switzerland
Quaternary Research Center and Department of Earth and Space Sciences, University of Washington, Seattle,
Washington, USA
3
Department of Civil and Environmental Engineering, University of Alberta, Edmonton, Alberta, Canada
4
1734 Bannister Road, Anchorage, Alaska, USA
5
Institute of Geosciences, University of Oslo, Oslo, Norway
6
Institute of Remote Sensing and Photogrammetry, Graz University of Technology, Graz, Austria
7
Northern Engineering Center (CINEP), Ecole Polytechnique, Université de Montréal, Montréal, Canada
8
Institute of Geoscience, University of Tsukuba, Ibaraki, Japan
9
ETH Zurich, Zurich, Switzerland
10
University of Basel, Basel, Switzerland
2
ABSTRACT
This review paper examines thermal conditions (active layer and permafrost), internal composition
(rock and ice components), kinematics and rheology of creeping perennially frozen slopes in cold
mountain areas. The aim is to assemble current information about creep in permafrost and rock glaciers
from diverse published sources into a single paper that will be useful in studies of the flow and
deformation of subsurface ice and their surface manifestations not only on Earth, but also on Mars.
Emphasis is placed on quantitative information from drilling, borehole measurements, geophysical
soundings, photogrammetry, laboratory experiments, etc. It is evident that quantitative holistic
treatment of permafrost creep and rock glaciers requires consideration of: (a) rock weathering, snow
avalanches and rockfall, with grain-size sorting on scree slopes; (b) freezing processes and ice
formation in scree at sub-zero temperatures containing abundant fine material as well as coarse-grained
blocks; (c) coupled thermohydro-mechanical aspects of creep and failure processes in frozen rock
debris; (d) kinematics of non-isotropic, heterogeneous and layered, ice-rich permafrost on slopes with
long transport paths for coarse surface material from the headwall to the front and, in some cases,
subsequent re-incorporation into an advancing rock glacier causing corresponding age inversion at
* Correspondence to: Professor Wilfried Haeberli, Glaciology and Geomorphodynamics Group, Geography Department, University of Zurich, Switzerland. E-mail: [email protected]
y
Present address: Institute of Geosciences, University of Oslo, Oslo, Norway.
z
Final report of the Task Force on Permafrost Creep and Rock Glacier Dynamics. This Task Force was established in 1998 by the
International Permafrost Association (IPA) and the International Commission on Snow and Ice (ICSI; now the Commission on
Cryospheric Sciences, CCS) with Wilfried Haeberli and Bernard Hallet as co-chairs in order: (1) to define the state of knowledge
relevant to the flow and evolution of perennially frozen ice/rock mixtures; and (2) to provide an overview of ongoing studies that
obtain quantitative information from drilling, geophysical soundings, geodetic/photogrammetric monitoring and measurements of
surface conditions. For this purpose, a special half-day workshop took place on 27 March 2001 within the framework of the First
European Permafrost Conference in Rome, Italy (Rea, 2001). Further discussions followed during and after the 8th International
Conference on Permafrost in Zurich (20–25 July 2003). The present contribution is based on the corresponding results and
constitutes the final product of the Task Force. The following colleagues served as lead authors of the main subsections in this report:
thermal conditions: Ole Humlum (surface, active layer) and Daniel Vonder Mühll (boreholes); composition: Norikazu Matsuoka
(rocks) and Roger Elconin (ice); geometry/kinematics: Andreas Kääb and Viktor Kaufmann (photogrammetry, geodesy); and
rheology: Sarah Springman, Lukas Arenson (mountains) and Branko Ladanyi (lowlands).
Received 6 June 2006
Revised 20 June 2006
Copyright # 2006 John Wiley & Sons, Ltd.
Accepted 20 June 2006
190 W. Haeberli et al.
depth; and (e) the dynamics of rock glaciers, which include spatial and temporal variations in velocity
that are manifested in the ridges, furrows and other surface structures typical of rock glaciers, as well as
their down-valley motion. Copyright # 2006 John Wiley & Sons, Ltd.
KEY WORDS: permafrost; rock glaciers; cold mountains; periglacial slope movement; geomorphic landforms; Martian
ice flow features
INTRODUCTION
The growing interest in mountain permafrost has led to
rapid progress in recent years in understanding both
the gravitationally-induced creep of permafrost on
slopes in cold mountain areas and the motion of rock
glaciers (Figure 1). The planetary science community
has also become increasingly interested in diverse
features visible on Mars that resemble rock glaciers or
Figure 1 Active rock glacier ‘Giandia Grischa’ in the eastern
Swiss Alps. The field of the flow vectors is continuous and quite
homogeneous, as expected with thermally controlled occurrence of
subsurface ice/rock mixtures. Flow velocities range from a few
centimetres (marginal zones) to slightly less than a metre per year,
and the time period involved with the development of the cumulative
talus deformation—which creates the spectacular landform—is
measured in millennia (Frauenfelder et al., 2005).
Copyright # 2006 John Wiley & Sons, Ltd.
are suggestive of deformation of subsurface ice
(Head et al., 2005). The widespread occurrence of ice
just below the surface is consistent with measurements
from the gamma-ray and neutron spectrometers on
board Mars Odyssey (Feldman et al., 2004) and
theoretical expectations (Mellon and Jakosky, 1995).
In addition, conservative estimates suggest that Mars
may possess a planetary inventory of water equivalent
to a global ocean 0.5–1 km deep of which as much as
90–95% is believed to reside as ground ice and
groundwater within the planet’s crust (Clifford et al.,
2003). Hence considerable attention is focused on
Martian permafrost, especially because the continuing
interest in extraterrestrial life is concentrated on the
surface where the common ingredients to sustain
life—photons and water—occur together.
In this paper, available knowledge is reviewed
that could eventually: (1) help advance our understanding of rock glaciers and related features on
Earth and Mars; and (2) form the basis for numerically modelling the thermo-mechanically coupled
dynamics of creeping ice/rock mixtures. It summarises recent sophisticated field and laboratory experiments (drilling, borehole observations, geophysics,
photogrammetry and remote sensing, creep tests, etc.)
that have improved our understanding of the principal
processes underlying the motion of deep-seated, largescale creep of perennially frozen debris in alpine
settings. The report also refers to the large set of wellestablished methods, from simple to sophisticated,
that are currently available for probing, mapping and
monitoring permafrost, and for modelling its spatial
distribution in complex high-mountain topography
(e.g. Etzelmüller et al., 2001; Harris et al., 2001;
Hoelzle et al., 2001; Vonder Mühll et al., 2002).
The term ‘rock glacier’ has been used for
various complex landforms of cold mountain areas.
Two extremes represent the ends of a geomorphic
continuum: (a) steadily creeping perennially frozen
and ice-rich debris on non-glacierised mountain
slopes; and (b) debris-covered glaciers in permafrostfree areas. Semantically, they can be differentiated as
follows: (a) is a ‘rock glacier’ whereas (b) is commonly
called a ‘debris-covered glacier’. Terminological
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
Permafrost Creep and Rock Glacier Dynamics
problems arise in areas with continental climates
where surface ice interacts with periglacial and
subglacial permafrost (Haeberli, 1983). There, complex mixtures of buried surface ice and ice formed in
the ground can occur and all transitions from debriscovered polythermal or cold glaciers to ice-cored
moraines and deep-seated creep of perennially frozen
sediments are possible (cf. Haeberli, 2000; Kneisel
et al., 2000). Narrow definitions of landform terms are
confusing in such cases and should be replaced by a
flexible concept considering, for each individual case,
the materials involved, thermal conditions, processes,
time periods and spatial scales. The following sections
concentrate predominantly to exclusively on rock
glaciers on non-glacierised mountain slopes.
A clear distinction must be made between two
fundamental but complementary characteristics
(Haeberli, 2000):
ground thermal conditions (permafrost) that allow
for the formation and the long-term existence of
subsurface ice; and
composition of the ice/rock mixtures, that defines
the amount and distribution of ice existing below
the surface and enabling the steady-state creep
of the ice/rock mass considered.
Permafrost is used here to describe a specific ground
thermal condition (temperature at or below 08C for a
minimum of one—or even better two—years),
irrespective of ice content or lithology. Such
thermal conditions not only enable the formation
and preservation of a variety of ground ice but also
determine the spatio-temporal scales of permafrost
creep within ice-containing materials.
191
Surface Conditions and Energy Fluxes.
In mountain areas outside the High Arctic, many
rock glacier fronts closely approach limits of local
permafrost occurrence. As a consequence, mean
annual surface temperatures are often not far from
melting conditions. Rock glacier surfaces are often
characterised by a system of ridges and depressions,
orientated parallel or transverse to the overall flow
direction (Figure 2). Numerical modelling of energy
fluxes through such surfaces requires parameterisation
of solar radiation, sensible heat, surface albedo, heat
conduction and advection, latent heat transfer, etc., for
complex topography and extremely heterogeneous
surface characteristics (air, rock); it needs a correspondingly large amount of precisely measured or
computed data. Models of energy-balance and
thermal-offset (accounting for the difference between
the temperature at the surface and at the permafrost
table, respectively (Mittaz et al., 2000; Hoelzle et al.,
2001)) must be combined. Modelling the evolution of
the snow cover (Mittaz et al., 2002) is one of the
principal components because snow cover not only
insulates the ground from the cold air during winter
and from warm air in summer but also critically
influences the surface albedo and, hence, net radiation.
Initial results of detailed modelling (Stocker-Mittaz
et al., 2002) and measurements in coarse debris
mantling a glacier (e.g. Conway and Rasmussen,
2000) clearly show that advective subsurface fluxes
are important. Special conditions are encountered at
the head of many rock glaciers under near-perennial to
perennial patches of surface ice—often deposited by
THERMAL CONDITIONS
Surface and Active Layer
The seasonally frozen active layer above the
perennially frozen material is commonly a few
decimetres to a few metres thick (Humlum, 1997).
Energy fluxes at the surface and thermal characteristics of the active layer control the initiation, growth
and maintenance of rock glaciers. The coupling
between atmospheric and ground thermal processes
due to the non-conductive heat transfer processes is
highly complex. Increased knowledge of this coupling
and of rock glacier active-layer thermal characteristics
is, therefore, essential for understanding the age,
climatic sensitivity and palaeoclimatic significance of
rock glaciers (Humlum, 1996).
Copyright # 2006 John Wiley & Sons, Ltd.
Figure 2 Example of a rock glacier active layer surface with
3–10 m topography and associated variations in maximum grain
size. The biggest rock fragments visible measure about 1.5 m, Disko
Island, Greenland, September 2000.
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
192 W. Haeberli et al.
spring avalanches. Such perennial snowbanks or
glacierets are generally thin (metres), cold and frozen
to the bed, because ice temperature cannot rise above
zero during the summer while temperatures drop far
below zero during the winter (Haeberli, 2000). As a
consequence, they are conducive to generating
negative ground temperatures and participate in the
build-up of frozen material, which tends to creep
downslope.
Active Layer Characteristics and Processes.
The thickness of the active layer on rock glaciers is
a function of local meteorological characteristics,
thickness and duration of the winter snow cover, and
the active layer’s textural characteristics. The grain
size usually decreases with depth in the active layer.
With increasing depth, fine-grained material gradually
becomes abundant, and close to the permafrost table,
fines are often predominant (Figure 3). On pebbly rock
glaciers (Matsuoka and Ikeda, 2001), fines may be
significant through most of the active layer. Especially
large boulders may penetrate the general terrain
surface to form ‘natural cairns’ (Humlum, 1988).
Active-layer thicknesses from about 0.5 m to 3 m are
typically reported from excavations, drill-holes or
geophysical soundings (e.g. Barsch, 1996; Vonder
Mühll, 1996; Wagner, 1996; Humlum, 1997; Isaksen
et al., 2000). Depending upon deformation of the rock
glacier, the active layer may extend across zones of
longitudinal extension and compression, giving rise to
local thermo-mechanical interactions (Haeberli and
Vonder Mühll, 1996), introducing at a local scale
either degradation or aggradation of the permafrost
Figure 3 Excavation in the 1.2 m thick active layer on a rock
glacier near Longyearbyen, Svalbard, August 2000. The surface of
the active layer consists of rock fragments 10–60 cm in size. With
depth, fines become more abundant, especially close to the permafrost table (exposed at the bottom of the excavation).
Copyright # 2006 John Wiley & Sons, Ltd.
below due to mechanical thickening (compression)
and thinning (extension) of the active layer.
Automatic dataloggers can readily be used to
monitor ground temperatures, snow cover duration,
zero-curtain effects, etc. During winter periods with
low wind speed and thin snow cover, radiative heat
losses may produce very cold air, which tends to
penetrate into the active layer. For example, rock
glacier active layer temperatures in Greenland in early
October follow the surface temperature with little lag
(Figure 4). A thin snow cover during autumn and
winter contributes to rapid cooling of the active layer,
partly due to surface radiation effects (Keller and
Gubler, 1993), but primarily by not blocking surface
openings in the active layer surface (Bernhard et al.,
1998).
Cold air penetrating somewhere upslope slowly
flows downslope within the active layer, to resurface at
lower elevations. In contrast, when a thick snow cover
exists on rock glaciers, active layer temperatures are
mainly controlled by conduction (Harris and Pedersen,
1998; Humlum, 1997; Hoelzle et al., 1999). In the
latter situation, temperature variations are smooth and
follow the surface temperature with considerable lag.
In Figure 4, this characterises the period from late
November to late April. Non-conductive processes
may, however, be important if vertical holes (funnels)
exist in the winter snow, connecting the outside air
with the air beneath the snow cover. As described by
Keller and Gubler (1993) and Bernhard et al. (1998),
such funnels are mainly found along surface depressions in the rock glacier surface and provide easy
access for cold air to reach the interior of the active
layer, displacing warmer air upwards.
The presence of snow cover in late winter and early
spring introduces non-conductive heat transfer processes during periods of snowmelt, as surface meltwater percolates into the active layer and refreezes
(Humlum, 1997). Similarly, if warm air masses
deliver rain on cold snow in the winter or spring,
the refreezing of percolating water can warm the snow
and the upper part of the active layer appreciably
(Putkonen and Roe, 2002).
Such an event is illustrated in Figure 4 at the end of
May 2001, when temperatures rise rapidly at all levels
in the active layer. Radiation effects during summer
produce very warm air masses at the surface. However,
due to density differences, such warm air rises and
does not penetrate into the active layer. Corresponding
heat transfer is conductive, while cooling during clear
nights may take place much more rapidly due to
density-induced non-conductive effects (Humlum,
1997). This is essentially what has been called Balch
ventilation, defined as the insulating effect of air-filled
Permafrost and Periglac. Process., 17: 189–214 (2006)
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Permafrost Creep and Rock Glacier Dynamics
193
Figure 4 Surface and active layer temperatures (8C) and photogaphs illustrating conditions from October 2000 to June 2001 at
Mellemfjord, Disko Island, central West Greenland. Upper left photograph: 4 October 2000; upper right: 14 May 2001; lower left:
24 May 2001; lower right: 26 June 2001. The photographs were obtained by an automatic digital camera looking SW towards two cirques
with rock glaciers. The mountains on the opposite side of the fjord reach about 1000 m a.s.l.
voids in coarse deposits (Thompson, 1962; Barsch,
1996). Wind pumping generated by high wind speeds
may, however, upset the above general pattern and
drive warm surface air masses into the active layer by
way of forced ventilation, dynamically displacing
colder and denser air masses (Humlum, 1997), as is
well known from similar processes in snow (Clarke
et al., 1987; Colbeck, 1989).
Copyright # 2006 John Wiley & Sons, Ltd.
Permafrost Conditions
Boreholes.
Relatively few boreholes have been drilled into and
through rock glaciers to permit investigation of the
thermal characteristics of the deeper subsurface. In the
1970s, two boreholes were completed in the Swiss
Alps (10 m at Murtèl: Barsch, 1977a; 7 m at Gruben:
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
194 W. Haeberli et al.
Barsch et al., 1979) and one in Canada (Kluane Range:
Johnson and Nickling, 1979). In the latter, temperatures remeasured some six years after drilling were
no longer negative indicating a very rapid thaw of
the permafrost (probably assisted by water advection).
In 1987, a 58 m deep borehole drilled through the
Murtèl-Corvatsch rock glacier initiated one of the
longest measurement time series of Alpine permafrost
temperatures (Vonder Mühll and Haeberli, 1990;
Haeberli et al., 1998; Vonder Mühll et al., 1998). At
Pontresina-Schafberg, two boreholes were drilled in
1990 (37 m and 65 m deep) as part of a project
concerning snow avalanches and debris flows (Vonder
Mühll and Holub, 1992). A borehole was cored to
10 m depth in 1995 at Galena Creek, Wyoming,
USA (Ackert, 1998). Within a joint project of
three institutes of the Swiss Federal Institute of
Technology (ETH) Zurich, several boreholes
were cored in the Muragl rock glacier (four boreholes,
ca. 70 m deep, 1999) and in the Murtèl-Corvatsch
rock glacier (two boreholes, 51 m and 63 m, 2000:
Arenson and Springman, 2000; Arenson, 2002;
Vonder Mühll et al., 2003). The main results obtained
concerning thermal conditions are summarised in
Table 1.
The rock glaciers investigated are all located near
the local boundary of permafrost distribution and
correspondingly contain warm permafrost typically
some tens of metres thick. Kudrayahov et al. (1991)
discuss technical aspects of drilling in permafrost.
Details about drilling in mountain permafrost and
parameters to be investigated are given in Vonder
Mühll (1996) and Arenson (2002).
Data Analysis and Results.
Heat transfer in permafrost largely occurs through
conduction (Lachenbruch et al., 1988), hence permafrost temperatures can be analysed using classical heat
conduction theory (Carslaw and Jaeger, 1959). Borehole
temperatures have been measured in rock glaciers
Table 1
with an absolute and relative accuracy of < 0.18C
and < 0.058C, respectively (Vonder Mühll, 1992),
and borehole measurements in other permafrost
regions and ice sheets indicate that modern instrumentation could provide considerably more precise
definition of the thermal regime of rock glaciers (e.g.
Cuffey and Clow, 1997; Dahl-Jensen et al., 1998).
The most systematic analysis of temperatures
observed in a rock glacier borehole was made for
the 1987 borehole at Murtèl 2 (Vonder Mühll and
Haeberli, 1990; Vonder Mühll, 1992; Vonder Mühll
et al., 1998). Murtèl rock glacier is located at the lower
limit of permafrost distribution and its tongue is
advancing onto permafrost-free terrain. Heat flow was
determined as about 150 mW m2, with thermal
conductivity of the frozen material being determined
(a) experimentally on core samples in the laboratory
and (b) using observed amplitude reductions and
phase lags at depths. The exceptionally high
temperature gradient and heat flow are caused by
the existence of a talik at a depth of about 50 m
(Vonder Mühll, 1992), probably maintained by
groundwater flow in summer. At the same site, the
time series of borehole data collected since 1987
(Figure 5) clearly documents some thermal characteristics of permafrost, in particular:
below the active layer, heat is transported almost
exclusively by conduction resulting in an exponential decrease in amplitude and a linear increase in
phase lag with depth;
the subsurface acts as a low-pass filter in terms of
the temperature signal introduced at the surface,
resulting in a low-amplitude sinusoidal annual signal at about 10 m depth and a temperature that is
close to the annual average;
snow strongly influences permafrost temperatures,
and
since the beginning of the measurements, temperatures ranged from 2.68C to 1.28C at 10 m
Key temperature data from boreholes through rock glaciers.
Site
Murtèl-Corvatsch (Switz.)
Gruben
RG II, Kluane Range
Pontresina-Schafberg 1/1990
Pontresina-Schafberg 2/1990
Galena Creek
Muragl
Number of
boreholes
Maximum
drilling
depth (m)
Mean
temperature at
the permafrost
table (8C)
Permafrost
thickness
(estimated) (m)
Depth of
shear zone (m)
5
1
1
1
2
1
4
63
7
17
65
37
10
72
2.0
1.0
0.6
1.5
0.5
1?
>0.5
>50, talik
80
>20
>60
30
?
20
22–30
Copyright # 2006 John Wiley & Sons, Ltd.
16
ca. 30
17–20
Permafrost and Periglac. Process., 17: 189–214 (2006)
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Permafrost Creep and Rock Glacier Dynamics
195
Figure 5 Time series of borehole temperatures at various depths (2.55, 3.55, 9.55 and 19.57 m) within the active Murtèl rock glacier, eastern
Swiss Alps.
depth, and from 2.18C to 1.48C at 20 m with a
warming trend of a few tenths of degrees celsius per
decade.
The temperature at the base of the active layer is the
upper boundary condition for the permafrost temperatures at depth. From borehole temperature data
above the depth of zero annual amplitude, the date of
the first major snowfall, the duration of positive and
negative temperatures and the occurrence of zerocurtain effects—especially during snowmelt—appear
to be the most important factors influencing overall
ground temperatures. Zones of circulating water or air
can exist within or below the permafrost body of a
rock glacier. Besides the talik already mentioned at
Murtèl rock glacier, water was observed in several
boreholes in connection with pressuremeter tests in
two boreholes (1/2000 (Arenson, 2002) and 2/2000),
and a sudden warming took place between 3 and
20 m depth on 25 April 2001 (Vonder Mühll et al.,
2003) without affecting the top 3 m of the borehole.
In the latter case, warming must have happened
from the bottom upwards, and subsequent cooling
took place from the top downwards. Indications of
temperature variations being closely related to
atmospheric temperatures and layers of frozen ground
below the main portion of the permafrost were
observed in all three boreholes from the Muragl rock
glacier. Such short-term changes at depth are
interpreted to arise from heat advection due to air,
Copyright # 2006 John Wiley & Sons, Ltd.
rather than water, movements (Vonder Mühll et al.,
2003).
COMPOSITION
Rock Component
The characteristics and origin of the rock component
of rock glaciers can be examined in the light of
lithological influences, debris supply processes and
subsurface structure. The initial formation of a rock
glacier requires thick debris accumulation under a
periglacial climate conducive to permafrost. Typical
origins of such debris accumulation include debrisladen snow avalanches, episodic rock avalanches and
long-lasting rockfall activity, which are triggered by
unloading during deglaciation, earthquakes, heavy
rains or weathering processes. Further rock glacier
development requires continuous debris input from
rock walls, which balances the removal of debris as it
is advected downslope by the motion of the rock
glacier (e.g. Burger et al., 1999; Haeberli et al., 1999;
Humlum, 2000). Modelling of rock glacier dynamics
should incorporate debris budgets in the rock wallrock glacier system. This downslope transfer of debris
has its own significance as a slope process that can
play an important role in the geomorphic evolution of
cold-climate, high-relief areas (Barsch, 1977b, 1977c,
1996; Frauenfelder et al., 2003; Humlum, 2000).
Permafrost and Periglac. Process., 17: 189–214 (2006)
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196 W. Haeberli et al.
Debris input mainly results from physical weathering in rock walls, which periodically releases blocks
of rock that fragment into coarse materials as they
fall onto talus slopes and rock glaciers. In an early
model of rock glacier evolution, Olyphant (1983,
1987) simply assumed the rate of debris input to be a
function of a topographic factor, the area of the rock
wall above the rock glacier. This simplification is
probably appropriate at a local scale where climate
and geology can be assumed to be similar. A universal
model of the rate of debris input, however, should
encompass three kinds of source rock wall characteristics: topographic (height, width, gradient and
exposure of the source rock wall), geological (grain
sizes and pore structures that affect weathering, joint
spacing, continuity and orientation of joints and other
discontinuities that dictate large-scale strength) and
climatic factors (thermal and hydrological regimes).
Lithological Controls.
Table 2 highlights geological effects on rock glacier
distribution. Four major rock types, granite, gneiss,
sandstone and limestone, are commonly found in rock
glaciers in the form of boulders or blocky debris.
Lithologies that mainly produce finer or platy debris,
including schist and shale, are more rarely sources of
material for rock glaciers (e.g. Wahrhaftig and Cox,
1959; Evin, 1987). Rock walls composed of the former
group of lithologies often feed relatively large rock
glaciers that have a blocky outer layer lacking
interstitial fine debris, which can be termed ‘bouldery’
rock glaciers. In contrast, the latter group produces
Table 2 Bedrock geology and rock glacier distribution
compiled from the literature (data after Matsuoka and
Ikeda, 2001).
Lithology
Major type of
surface material
Number
of regions
Granite (incl.
granodiorite)
Gneiss
Sandstone
Limestone (incl.
dolomite)
Schist
Basalt (incl.
andesite)
Shale (incl.
mudstone, phyllite)
Quarzite
Rhyolite
Conglomerate
Porphyry
Bouldery
16
Bouldery
Bouldery
Bouldery
14
13
12
Pebbly
Bouldery
10
8
Pebbly
8
Bouldery
Pebbly
Bouldery
Bouldery
6
3
3
2
Copyright # 2006 John Wiley & Sons, Ltd.
much smaller forms (mostly less than 300 m in length)
termed ‘pebbly’ rock glaciers, which consist mainly of
pebbles and cobbles filled with fine debris (Matsuoka
et al., 2005; Figure 6A). Despite their size, the latter
group is more susceptible to weathering than the
former, so that the debris input to rock glaciers per unit
area of rock wall may actually be more rapid. At least
two conditions are responsible for small forms. First,
mechanically weak rocks can only support low rock
walls, which restrict the total debris input despite the
higher weathering susceptibility. Where such a
lithology constitutes an exceptionally high rock wall,
however, large pebbly rock glaciers can develop
(Figure 6B). Second, the fine-grained debris that is
relatively abundant in the active layer of pebbly rock
glaciers is frost-susceptible, and hence is prone to
downslope motion by solifluction or debris flows. The
resulting rapid downslope transport of debris inhibits
debris accumulation. The rounded frontal slope of
pebbly rock glaciers may also reflect such surficial soil
movements.
Environmental Controls.
Debris production processes can be classified in
terms of temporal and dimensional scales (Whalley,
1984; Matsuoka et al., 1998). Weathering associated
with diurnal freeze-thaw and thermal fatigue can only
be effective on the outermost rock a few centimetres or
decimetres thick and causes small rock falls. Weathering associated with seasonal temperature variations,
including annual freeze-thaw cycling, can be effective
to a depth of a few metres, possibly triggering larger
rock falls and potentially removing the entire active
layer. Major rock avalanches and rock wall failures
can occur progressively over longer time scales as a
result of: (1) increasing gravitational stresses by
oversteepening the rock wall or decreasing the support
provided by permafrost due to lowering of the rock
glacier surface; and (2) weakening the rock mass
either by increasing stresses inside rock fractures, due
to increasing water or ice pressure, or by localised
chemical weathering that slowly lowers the fracture
toughness of rock (Hallet et al., 1991). Rock
avalanches are also commonly observed in association
with major earthquakes in mountainous terrain (e.g.
Bull et al., 1994; Keefer, 1994; Bull and Brandon,
1998). A question is then raised: which process is most
responsible for rock glacier development?
Field observations suggest that seasonal events
mainly contribute to talus accumulation, at least in
areas that are not subject to frequent and large
earthquakes. Boulder falls associated with annual
freeze-thaw action are considered to prevail for many
rock glaciers, because they are volumetrically
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Permafrost Creep and Rock Glacier Dynamics
197
Figure 6 Pebbly rock glaciers in the Swiss Alps having a rounded frontal slope and a gravelly active layer filled with fine debris. (A) Small
forms developed below a shale rock wall 70 m high, Piz Nair, Upper Engadin. (B) Extraordinarily large forms (ca. 1 km long) developed at
the foot of a shale rock wall 300 m high, Piz Üertsch, Albula.
dominant and most common during seasonal thawing
periods (e.g. Church et al., 1979; Matsuoka et al.,
1998). Figure 7 shows a large boulder fall that
occurred above the Murtèl rock glacier in June 1997.
The debris volume is equivalent to 1 mm of retreat of
the whole rock wall, which is comparable with the
Late Holocene mean annual retreat rate for this site
(Haeberli et al., 1999). Other studies also estimate that
rock walls above active rock glaciers have been
retreating at rates in the order of millimetres per year,
based on estimates of what is needed to sustain the flux
of debris in rock glaciers (Barsch, 1977c, 1996; Frich
and Brandt, 1985; Humlum, 2000). Moisture and rock
Figure 7 A large rock fall generated from the rock wall above
Murtèl rock glacier during a thawing period (photograph taken on
15 June 1997). The measurement of large blocks on snow indicated
that the total rock fall volume was about 100 m3, which is equivalent
to the rock wall retreat of about 1 mm.
Copyright # 2006 John Wiley & Sons, Ltd.
joint distribution in the bedrock also affect the
production rate of rock fall debris, but little is known
about how these factors vary in space and time in any
area. Freeze-thaw weathering may be most intensive
in a humid zone near the base of the active layer
(Hallet et al., 1991), which suggests that the thaw
depth affects the maximum size of removable rock.
The size of the rock fragments in rock falls depends
also on the joint spacing in the bedrock (Matsuoka,
2001). Because joints in schist and shale are closely
spaced, for example, short-term temperature cycling
(e.g. Matsuoka, 1991), which affects only the outer
few centimetres of rock walls, can be important in
delivering debris to pebbly rock glaciers where these
lithologies occur. Debris is also transferred by
secondary processes that remove material already
loosened by weathering processes and temporarily
accumulated on ledges or in couloirs. These processes
include snow avalanches and debris flows, which are
of particular significance in transporting fine materials
onto rock glaciers.
Development of the Two-layer Structure.
Bouldery rock glaciers are composed of outer
blocky sediments representing the active layer and an
inner matrix-supported frozen core (e.g. Wahrhaftig
and Cox, 1959). The outer blocky layer lacks
interstitial fine materials and the blocks are generally
much larger than those comprising the talus above the
rock glacier. The inner frozen core is a mixture of
clasts, fine-grained lithic fragments and various types
of ground ice (Fisch et al., 1977; Haeberli and Vonder
Mühll, 1996; Elconin and LaChapelle, 1997). The
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198 W. Haeberli et al.
wide range of the grain size distribution reflects the
diverse processes by which material is transferred to
the rock glacier surface, including rockfalls, debris
flows and snow avalanches. Several factors probably
contribute to the formation of the outer blocky layer,
including fall sorting, by which—during a rockfall
event—the larger blocks have sufficient kinetic
energy to reach the rock glacier surface while finer
clasts are trapped on the talus (e.g. Haeberli et al.,
1998). The finer debris eventually covers the larger
blocks due to creep, and subsequently new boulders
fall and roll onto the finer frozen core. This process
develops the two-layer structure and can explain the
difference in the size of blocks. Other plausible factors
include ‘kinetic sieving’, a phenomenon in which
coarse particles in granular mixtures gradually move
upward as the mixture is shaken (which for our
purposes corresponds to disturbing the outer blocky
layer due to deformation of the rock glacier surface)
because of the likelihood that small grains can settle
further than larger ones (Rosato et al., 1987) and
washing away of fine-grained material when permafrost has thawed (e.g. Giardino and Vick, 1987).
However, neither of these processes accounts for
blocks in the outer layer tending to be larger than those
in the talus above the rock glacier and washing appears
ineffective on pebbly rock glaciers.
Understanding the genesis of the internal structure
is relevant when modelling the debris budget, because
it affects the debris flux in the rock glacier.
Ice Component
Information about the internal structure of rock
glaciers is important but very difficult to obtain.
Hence, few direct observations have been reported on
ice distribution or ice physical and chemical characteristics, including organic impurities.
Occurrences, Volumes and Structures.
The structure of the permafrost core of rock
glaciers, as determined by direct observation of
borehole cores, large outcrops and tunnels is
summarised in Table 3. These observations reveal
three general types of ice occurrences: (1) massive ice
with dispersed debris and intercalated debris-laden
ice, overlying a basal layer of rock; (2) a few metres to
tens of metres of debris-laden ice overlying massive
ice; and (3) ice/rock mélange throughout (Figure 8).
Ice crystals in massive ice in thin section and in
natural, sun-etched exposures consist of fine-to-coarse
interlocking grains with diameters of 5 mm (Barsch
et al., 1979), 5–10 mm (Wagner, 1990) and 10–30 mm
(Elconin and LaChapelle, 1997). Wagner (1990)
Copyright # 2006 John Wiley & Sons, Ltd.
measured C-axis orientations at a depth of 9.5 m in
Murtèl rock glacier; the mean dip of the basal planes
was 11.58 from horizontal, roughly parallel to the
surface slope. Spherical and elongated air bubbles are
common in rock glacier ice. Air bubbles 1–2 mm in
diameter and elongated up to 20 mm in the direction of
flow were reported by Elconin and LaChapelle (1997).
The air content in the upper 10 m of a core retrieved
from Murtèl rock glacier averaged 3% by volume
(Wagner, 1990). Alternating layers of bubbly and nonbubbly ice can also be prevalent (e.g. Barsch et al.,
1979); Elconin and LaChapelle (1997) reported
foliation, defined by layers of bubbly and bubblefree ice, which was strongly associated with other
flow-generated structures.
Geochemical and Organic Constituents for
Determining Ice Origin and Age.
18
Oxygen and Deuterium: The relationship between
these stable isotopes in rock glacier ice has been
subject to in-depth analysis in two studies (Stauffer
and Wagenbach, 1990; Steig et al., 1998). Results
from both suggest that insignificant isotopic fractionation occurred before ice formation, indicating a
meteoric or short-residence groundwater source.
Nuclear Bomb Era Radionuclides: Diverse radionuclides in rock glacier ice and meltwater have been
analysed, including 210Pb and 238U (Gäggeler et al.,
1990), 3H (Wagenbach, 1990), and 3H, 36Cl, and 35S
(Cecil et al., 1998). These studies suggest that modern
precipitation does not infiltrate more than 0.5 m into
rock glacier ice, the ice is more than a century old, and
that meltwater is probably derived from snow and ice
that fell after the end of atmospheric atomic bomb
testing.
Chemical Impurities: Chemical analyses of solute
and insoluble material from 17 m of core taken from
the Murtèl rock glacier were performed by Baltensperger et al. (1990). Analysis of ion constituents and
concentrations in solution show little if any anthropogenic impact. Extracted insoluble crystalline matter
was derived from local bedrock.
Organic Impurities: Stem remains from seven
different moss species were extracted from 3 m below
the active layer in the Murtèl rock glacier (Haeberli
et al., 1999). The 14C age of the moss was 2250
(100) years BP. Twenty-three taxa of pollen and
spores extracted from the moss were present in the
region between 2000–8000 years BP. These data and
other collateral evidence indicate ice aggradation
spans the greater part of the Holocene. Konrad et al.
(1999) reported the following 14C ages: (1) 200 40
years BP of a pine needle 100 m from the headwall at
an undisclosed depth; (2) 2250 (35) and 2040 (35)
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Permafrost Creep and Rock Glacier Dynamics
199
Table 3 Summary of direct observations of active rock glacier internal structures.
Exposure type
Internal
structure–generalised1
Reference
Rock glacier
name and location
Three tunnel levels
with several
cross-cuts
admixed ice and rock
throughout, estimated
ice volume 40–70%
Bateman and
McLaughlin
(1920)
horizontal tunnel
100 m long and
30 m below
surface
admixed ice and rock
throughout, with
massive ice near
lateral margin
Brown (1925)
‘Glacier Mine’
(informal),
south-central
Alaska, USA
‘Hurricane Basin’
(informal),
Colorado, USA
borehole, 10.4 m
stratified debris-laden
ice and massive ice,
estimated average ice
volume 50–60%
Barsch (1977a)
‘Murtèl I’, Swiss Alps
outcrop, 10 m deep
Ice/rock mélange with
deformed ice lenses,
high ice content
Fisch et al. (1977)
‘Comb de Prafleuri’
(informal), Swiss Alps
borehole, 7 m deep
stratified debris-laden
ice and massive ice,
estimated average ice
volume 50–70%
Barsch et al. (1979)
‘Gruben’, Swiss Alps
Three boreholes to
bedrock, 50–60 m
deep
massive ice in upper
25 m, with ice volumes
of 90 to 100%,
debris-laden ice and
massive ice layers at
intermediate depth
(7 m thick), and rock
with interstitial? ice to
bedrock
Haeberli et al. (1988);
Haeberli (1990);
Arenson et al. (2002)
‘Murtèl I’, Swiss Alps
outcrop, longitudinal
section >80 m
long, max.
thickness 14 m
debris-laden ice upper
2–3 m overlying massive
ice 3–11 m thick with
scattered rock fragments
Moore and Friedman
(1991)
‘Berthoud Pass’
(informal), Colorado,
USA
Four boreholes
admixed rock and ice up
to 24 m thick; drilling
stopped in ice laden with
fine grained debris
Hamre and McCarty,
1996; Hamre
et al. (2000)
‘East Lone Mountain’,
Montana, USA
outcrop, complete
transverse section,
>90 m wide, max.
thickness 60 m
several boreholes
Ice/rock mélange
throughout with folded
strata, ice volume >50%
Elconin and
LaChapelle (1997)
‘Fireweed’, (informal),
south-central Alaska,
USA
massive ice up to 25 m
thick with dispersed
debris overlying unknown
thickness of rock and ice
layers of ice and rock,
estimated ice volume 40–70%
Konrad et al. (1999)
‘Galena Creek’,
Wyoming, USA
Arenson et al. (2002)
‘Muragl’, Swiss Alps
Four boreholes to
bedrock,
64–72 m deep
1
The term ‘admixed ice and rock’ is used when field evidence of a sheared matrix is not given; ‘melange’ is used when
field evidence of a sheared matrix is given.
Copyright # 2006 John Wiley & Sons, Ltd.
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200 W. Haeberli et al.
glacial ice and of short-residence groundwater can be
quite similar, however. In addition, rock glacier ice is
likely to be polygenetic and may include significant
meteoric contributions. Moreover, debris-rich ice
typically from the basal portion of glaciers shares
many characteristics with diverse forms of ground ice.
Figure 8 Small section of the full transverse cross-sectional
exposure of Fireweed rock glacier, Wrangell Mountains, Alaska.
The rock and ice mélange is representative of the composition and
structure of the entire exposure.
years BP from unspecified organic material in ice at
the surface 400 m from the accumulation zone; and
(3) 1680 70 years BP from leaf fragments in the
upper 0.2 m of ice halfway down the rock glacier.
Ice Origin: Diagnostics.
Differentiation between ice of glacial or non-glacial
origin in permafrost is usually difficult because the
two types share petrographic, stratigraphic, physical
and chemical qualities (cf. Mackay, 1989; French and
Harry, 1990). Transformation of original structures
due to flow (i.e. flow metamorphism) in rock glacier
permafrost and the close proximity of some rock
glaciers to glacierets and small glaciers compound this
problem.
Petrofabrics: Glacier-like ice crystal morphology,
orientation and air bubble content observed in rock
glaciers do not indicate glaciogenic ice, as is
sometimes assumed. Post-depositional processes,
most notably flow metamorphism, lead to glacier-like
crystal morphology and orientation but are independent of ice origin. Spherical and elongate air bubbles,
in ordered or dispersed arrangements, are common in
both types of ice. Bubble foliation in glacier ice is
caused primarily by flow metamorphism (e.g. Hambrey,
1979), and by analogy, so is bubble foliation in rock
glaciers (Elconin and LaChapelle, 1997).
Stratigraphy: Small- to large-scale depositional and
deformational structures composed of massive and
debris-laden ice, formed in glacial and non-glacial
environments, can appear very similar.
Water Quality: A common assumption is that ice
formed in the non-glacial environment is composed
of ion-rich groundwater, while glacier ice is composed
of relatively pure meteoric water. The ionic content of
Copyright # 2006 John Wiley & Sons, Ltd.
Ice Origin: Accumulation Area (‘Rooting Zone’ or
‘Source Area’).
Typically, the head of a rock glacier is located at the
base of a bedrock headwall and is the primary
accumulation site for both rock debris and ice. Due
to dynamic and diverse atmospheric, hydrologic
and geologic conditions in alpine environments, a
wide range of ice formation processes can exist in
close proximity in space and time, particularly in the
accumulation zone. Observable types of ice formation
in the accumulation zone at the surface are related to
snow metamorphism and surface icings. In the
subsurface, negative mean annual ground temperatures and abundant liquid water assure ice growth in
rock debris in the form of interstitial, segregation,
injection or vein ice (Wahrhaftig and Cox, 1959;
Wayne, 1981; Haeberli and Vonder Mühll, 1996).
Sources for the liquid water include groundwater,
rain and meltwater from snow, avalanche runout, firn
and ice.
GEOMETRY AND KINEMATICS
There is a long tradition of geodetic and photogrammetric surveying of the effects of permafrost creep
within active rock glaciers (Kääb, 2005). Available
measurements indicate that the surface movements of
rock glaciers are characteristically slow (typically
decimetres per year) and that they form a continuous
and internally coherent flow field as expected with
thermally controlled, macroscopically homogeneous
permafrost. It is also evident from the surface flow
fields that the debris transported on the surface of rock
glaciers originates from talus and other accumulations, such as moraines or other heaps of loose
debris, and most likely over time scales of millennia
(Figure 9). The precise determination of surface flow
fields in rock glaciers also provides a context for
subsurface information obtained from boreholes and
geophysical soundings, and helps to determine the
degree to which this subsurface information is
representative spatially.
Optimal investigation of permafrost creep requires:
spatially distributed information on kinematics to
account for three-dimensional effects;
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Permafrost Creep and Rock Glacier Dynamics
201
1 m for rock glaciers in the Dry Valleys of Antarctica and
on Mt St Helens, USA (http://www.nasa.gov/vision/
earth/lookingatearth/mshelenslidar.html).
Photogrammetry is the remote sensing technology
most commonly used for rock glacier monitoring. Rock
glacier thickness changes and surface displacements
can be measured from repeated stereo-imagery (e.g.
Haeberli and Schmid, 1988; Kääb et al., 1997, 1998;
Kaufmann, 1998a, 1998b; Kääb, 2002, 2005). Recent
advances, especially in image processing, permit the
measurement of surface motion on rock glaciers with
unprecedented resolution, accuracy and spatial coverage, which has led to a number of new insights into rock
glacier development and behaviour (Kääb and Vollmer,
2000; Kaufmann and Ladstädter, 2002; Kääb et al.,
2003; Figure 10). Space-borne differential synthetic
aperture radar interferometry (DInSAR) is also able to
measure rock glacier surface displacements, especially
those in the vertical plane, with an accuracy of a few
millimetres to centimetres (Rott and Siegel, 1999;
Kenyi and Kaufmann, 2003; Nagler et al., 2001; Rignot
et al., 2002; Strozzi et al., 2004).
Figure 9 Horizontal surface velocities on Murtèl rock glacier for
1987–96. The dotted lines represent selected flow paths interpolated
from the velocity field over this period. The time step between two
adjacent dots is 100 years. The dashed transverse lines indicate 1000
year-isochrones interpolated from the flow paths. A white rectangle
at the central flow line and at about 2670 m a.s.l. indicates where
moss macrofossils dating from 2300 years BP were found in
the permafrost. Data computed from airphotos taken by the Swiss
Federal Office of Cadastral Survey (after Kääb et al., 2003).
Dynamic Processes
The main processes related to the evolution of a
rock glacier can be identified from high-resolution
precise high-resolution techniques capable of resolving the slow displacement over reasonable time
periods; and
long-term monitoring for documenting slow
temporal changes at a sufficient level of accuracy.
Monitoring Techniques
Ground-based approaches as well as air- and space-borne
ones have been applied for monitoring rock glacier
kinematics with high resolution. Ground-based surveys
utilise triangulation and laser ranging (e.g. Haeberli,
1985; Zick, 1996; Sloan and Dyke, 1998; Koning and
Smith, 1999; Konrad et al., 1999; Krainer and Mostler,
2000), or satellite geodesy (e.g. GPS; Berthling et al.,
1998). Other methods for deformation measurements,
such as steel tapes or strain wires, are seldom used (e.g.
Haeberli, 1985; White, 1987). Terrestrial laser scanning
provides local digital terrain models (DTM) with very
high resolution and accuracy (Bauer et al., 2003), and
airplane-based laser scanning systems have recently been
used to generate a DTM with a resolution approaching
Copyright # 2006 John Wiley & Sons, Ltd.
Figure 10 Horizontal surface displacements on a section of the
lowermost part of Outer Hochebenkar rock glacier, Austrian Alps,
for 1971–77. The terminus of the rock glacier showed a creep
instability indicated by missing targets for displacement measurements and crevasse-like transverse ruptures (after Kaufmann and
Ladstädter, 2002).
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202 W. Haeberli et al.
measurements of three-dimensional surface velocity
fields. The fundamental process is creep within the
permafrost. Nearly all high-resolution studies on
the deformation of (visually) active rock glaciers have
been able to detect creep. Whereas maximum speeds
are in the order of several metres per year, the
detection of minimum speeds seems mostly restricted
by the available measurement accuracy. From
comparison of available studies on rock glacier speed,
it is clear that differences in slope, thickness,
temperature, or internal composition are not able to
explain differences in speed at individual sites in any
straightforward way. On the other hand, within a
single rock glacier, the spatial pattern of surface
speeds often seems simply related to the pattern of
surface slopes (Konrad et al., 1999). ‘Cold’ polar rock
glaciers, in general, creep more slowly than ‘warm’
Alpine ones (Kääb et al., 2002). Relatively low speeds
at the rock-glacier root zones and margins probably
reflect thinner deforming layers (Kaufmann, 1998a;
Kaufmann and Ladstädter, 2002). Separation of
individual influences on surface speeds is difficult
due to the usual lack of knowledge about the internal
structure (thickness and composition).
The advection of surface micro-topography by
creep may result in a pattern of local positive and
negative thickness changes. A pattern of heaving in
front of individual transverse ridges and corresponding lowering behind them has been observed from
high-resolution studies (Kääb et al., 1998; Kääb and
Vollmer, 2000; Kääb and Weber, 2004). Comparing
the observed vertical rates with ones calculated from
creep speed and surface slope clearly confirms these
local patterns of surface elevation change. Threedimensional straining may lead to local heaving or
thinning. For an incompressible medium such as pure
ice, compression and extension should equal out.
Horizontal compression of the frozen debris is
accompanied by vertical extension, that is by surface
heaving and vice-versa. Such relationships between
horizontal and vertical strain rates can be applied for
rock glaciers where ice contents exceed saturation
(Kääb et al., 1998; Kääb and Vollmer, 2000; Rignot
et al., 2002) and where ground ice is not highly aerated
or crevassed.
General thaw settlement and frost heave as an
expression of climate forcing affect large parts of a
rock glacier in a similar way. Surface heaving or
settlement may also reflect the internal composition.
Thickness loss due to ice melt may be significant for
zones in the vicinity of perennial ice patches or for
dead ice remains (Kääb et al., 1997; Figure 11). Such
pronounced thermokarst phenomena in massive
ground ice or frost heave might, therefore, be an
Copyright # 2006 John Wiley & Sons, Ltd.
Figure 11 Thickness changes of Muragl rock glacier measured for
1981–94 from repeated aerial stereo-photography. Data computed
from airphotos taken by the Swiss Federal Office of Cadastral
Survey (after Kääb et al., 2003).
expression of thermal disequilibrium (active-layer
thickness > debris cover thickness) rather than a
climate signal (Haeberli and Vonder Mühll, 1996;
Kääb and Haeberli, 2001; Kääb and Reichmuth,
2005). According to the available monitoring series,
such small-scale differential melting or frost heave is
clearly an exception and usually represents rather a
disturbance of the general thermal equilibrium.
Widespread differential melting likely occurs only
at the front where the ice content melts out (Kääb
et al., 1997) or in areas of preferential water flow. In
fact, for most monitored rock-glacier mass changes,
no clear signal of overall mass gain or loss has been
observed. Only the study of the Murtèl rock glacier
clearly showed an overall mass loss of a few
centimetres per year (Kääb et al., 1998).
Little is known about temporal changes in rock
glacier creep rates. For some rock glaciers or parts of
them, current surface velocity fields clearly differ from
those of the past (Frauenfelder and Kääb, 2000). Over
shorter time scales, surface speeds vary both cyclically
and non-cyclically. Clear seasonal variations in speed
have been observed (Haeberli, 1985; Arenson et al.,
2002; Kääb et al., 2003). Non-cyclic pluriannual
changes in speed might result from external (climate?)
forcing (Zick, 1996; Kääb et al., 1997; Kääb and
Frauenfelder, 2001; Schneider and Schneider, 2001).
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Permafrost Creep and Rock Glacier Dynamics
Monitoring of rock glacier velocities to date has
revealed several possible causes, mostly connected to
impacts of climate variations at the scale of decades
and longer time periods.
Perspectives
The technology for monitoring rock glacier dynamics
is comparably well developed. Research limitations
primarily result from the limited availability of advanced techniques, the difficulty of developing and
implementing systematic monitoring strategies, and
the inherent difficulty of making subsurface measurements of rock glacier motion. Current research
tendencies point towards an increase in accuracy
and temporal and spatial resolution, with an improvement in measurement automation, and towards
enhanced application of space-borne techniques in
order to cover remote areas (Kääb, 2005). The
knowledge of geometry and surface kinematics of
rock glaciers is quite detailed, and the primary process
responsible for the creeping motion of rock glaciers is
relatively well understood to be due to creep
deformation in the ice, whereas analyses of the
dynamics of rock glaciers remain elusive. Displacement rates cannot be quantified because, contrary to
ordinary glaciers that are comprised of relatively
homogeneous ice, rock glaciers tend to consist of
inter-layered debris-rich and ice-rich domains, and in
general, information about the spatial extent and
rheological behaviour of these domains is insufficient
to permit detailed flow calculations. Moreover
boundary conditions equivalent to specifying accumulation and ablation rates are difficult to define,
especially in view of the multiplicity of inputs of
ice and debris in rock glaciers. The general lack of
information on the internal composition and architecture of rock glacier together with the complicated
and poorly-defined rheological behaviour of icy debris
and debris-rich ice limit attempts to quantitatively
understand velocity variations, sensitivity to external
forcing, rock glacier advance mechanisms, and the
development and dynamic interpretation of microtopography. A large number of globally distributed
monitoring series could substantially improve our
understanding of the basic processes of rock glacier
creep with their local and temporal variabilities.
CREEP OF ROCK GLACIERS
The creation of a unified constitutive model for
permafrost found in alpine rock glaciers, to represent
creep and failure in variable temperature regimes,
Copyright # 2006 John Wiley & Sons, Ltd.
203
remains a distant goal. The only reliable way at
present is to carry out analyses of individual
components of rock glacier with simplified geometry
and assumptions. Examples include separate consideration of: (1) deformation within the active layer
which is very strongly linked to pore pressures (e.g.
Arnold et al., 2005); (2) creeping of the main body
of the rock glacier, which can be approximated
using appropriate creep parameters for the relevant
stress state and temperature (Arenson and Springman,
2005b); or (3) rheological response to time-dependent
changes in temperature and hence state. All of these
interact and suitable parameters must be selected
accordingly. We may also assume that the slope is
infinite, or focus on more local viscous features, for
example on the surface of the rock glacier, depending
on our specific area of interest.
The rheological behaviour of rock glaciers is
highly sensitive to the character and temperature of
their deforming layers, which usually comprise
mixtures of ice and rock elements, the latter ranging
from silty sand up to large boulders. In addition,
unfrozen water may be present and influence
responses, at permafrost temperatures close to the
melting point of the ice, in particular at the base of the
rock glacier. As in the mechanics of rock masses,
large-scale behaviour of such mixtures is difficult to
evaluate either by direct in situ measurements or by
taking samples, but it can be estimated by a
constitutive modelling approach.
Modelling the entire creep process, following
instantaneous elastic and plastic deformation due to
initial load application, including primary (e.g. see
Vyalov et al., 1962; Domaschuk et al., 1991;
Wijeweera and Joshi, 1991, for specific models),
secondary (e.g. Andersland and AlNouri, 1970) and
tertiary creep, is now possible for ice or related
homogeneous mixtures according to developments by
Goughnour and Andersland (1968), Fish (1983, 1984),
Ting (1983), Zhu and Carbee (1983), or Gardner et al.
(1984). However these models are less suitable for
rock glaciers because of their heterogeneity. Moreover, some models only acknowledge tertiary creep
and for others, the secondary (steady state) creep
phase is reduced to an inflection point, that is the
decrease of the strain rate with time is directly
followed by acceleration, without an interval at a
constant minimum strain rate. While all these
deformation phases are important for certain practical
problems, such as foundations in permafrost, and are
well covered in the literature (e.g. Andersland and
Ladanyi, 2004), they are of less importance in
connection with rock glaciers, where secondary creep
(at a minimum strain rate) should dominate.
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204 W. Haeberli et al.
Glen’s (1955) original equation for secondary creep
of polycrystalline ice is often used for the creep of
permafrost and other polycrystalline solids:
"_ ¼ Bs n
(1)
where "_ is strain rate, s is a corresponding stress, and
the coefficients B and n are material constants. B
contains the effect of temperature, grain size, fabric
and impurity content; it can be written explicitly as
(e.g. Andersland and Ladanyi, 2004):
B¼
"_ c
s ncu
(2)
where:
u w
s cu ¼ s co 1 þ
uc
(3)
Here, "_ c denotes the reference strain rate, s cu is the
reference stress or creep modulus, for temperature
u ¼ T, and uc is an arbitrary reference temperature
(see data for polycrystalline ice from Morgenstern
et al. (1980) "_ c ¼ 105 h1 , sco ¼ 0.103 MPa, and
w ¼ 0.37 valid for T ¼ 18C, and 107 "_ 102 h1 ). Many authors (Vyalov et al., 1962;
Fish, 1983, 1984; Ting, 1983; Zhu and Carbee, 1983;
Gardner et al., 1984; Wijeweera and Joshi, 1991) have
reported results consistent with Equation (1) or similar
expressions from uniaxial creep tests on cylindrical
samples. Goughnour and Andersland (1968), Andersland and AlNouri (1970) and Domaschuk et al. (1991)
have done so based on triaxial test data, which have
the potential to represent the in situ stress state
more faithfully, although size effects (particle sizes
related to sample size) and sample disturbance must
be considered and minimised where possible, by
increasing sample size, using appropriate cooling
fluids, and cooling the drill bit during drilling and
extraction.
Figure 12 shows two different samples post-creep
(volumetric ice content: 55% (no. 27) and 74%
(no. 43)). The initial sample diameter was 74 mm and
the initial sample height was approximately 150 mm.
Significant compression, that is volume decrease,
was recorded for sample no. 43, which also had
a higher volumetric air content (26%, Arenson,
2002).
The homogeneous approach presented above, with
averaged creep parameters derived from small-scale
tests, works very well for uniform ground conditions
found for some arctic permafrost (in certain areas—
e.g. Morgenstern et al., 1980; Ladanyi, 2002; Sego
et al., 2003). Recent geotechnical and geophysical
investigations and monitoring at Muragl, MurtèlCorvatsch and Pontresina Schafberg based on earlier
work (e.g. Haeberli et al., 1998; Kääb and Vollmer,
2000) have confirmed, however, that the internal
structure of a rock glacier is extremely heterogeneous
(Arenson, 2002; Musil, 2002; Maurer et al., 2003),
leading to significant variations in the parameters
back-fitted to standard creep expressions (Arenson
and Springman, 2005a, 2005b), or requiring explicit
theoretical simulation. For example, Ladanyi (2002,
2003) and Ladanyi and Archambault (2003) attempted
to simulate the creep behaviour of ice/rock mixtures
and ice-filled irregular rock joints under variable stress
and temperature conditions.
Figure 12 (A) Sample no. 27 and (B) no. 43 after testing (Arenson, 2002). Both samples were initially tested under constant stress
(creep) and after a strain of 10.3% for (A), and 19.6% for (B) they were sheared with a constant strain rate (A: 0.72 105s1;
B: 0.80 105s1).
Copyright # 2006 John Wiley & Sons, Ltd.
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
Permafrost Creep and Rock Glacier Dynamics
Pressuremeter tests allow creep parameters to be
obtained in situ (e.g. in arctic permafrost or ice;
Ladanyi and Johnston, 1973; Kjartanson et al., 1988;
Ladanyi and Melouki, 1993; Ladanyi, 1996), as well
as data on load controlled (relaxation) tests. They are
promising, even for alpine permafrost (Arenson et al.,
2003b), because the internal conditions (in situ stress
and temperature) can approach reality if the pressuremeter probe is pre-cooled prior to insertion in the
borehole shortly after completion of drilling to test
depth using pre-cooled rotary boring techniques rather
than percussion methods. Nonetheless, the volume of
permafrost investigated by this technique in the
annulus around the probe is still very limited when
compared to the volume of a rock glacier. In addition,
temperature disturbances due to the drilling have to be
considered and the creep parameters derived can vary
significantly.
Temperature probably has the most significant
influence on creep behaviour. The minimum strain rate
approached during the secondary creep phase
increases with increasing temperature. The Arrhenius
equation may represent this behaviour acceptably at
temperatures colder than about 108C (e.g. Mellor
and Testa, 1969; Hooke et al., 1980). At temperatures
close to the melting point of ice, which are important
in terms of degrading permafrost, the unfrozen
water changes the behaviour (Barnes et al., 1971).
Voytkovskiy (1960) presents an empirical relationship
that was successfully applied (e.g. Hooke et al., 1980;
Arenson and Springman, 2005b).
The rheological behaviour of sand/ice mixtures
also depends on the applied normal stress. The
effect can be estimated using the methods presented,
for example, in Andersland and Ladanyi (2004).
Arenson and Springman (2005a, 2005b) report
results from triaxial constant strain rate and constant
deviatoric stress tests on ‘undisturbed’ samples from
Muragl and Murtèl-Corvatsch (Figure 12) dependent upon a constant confining cell pressure (s3),
deviatoric stress (s1 s3), temperature and volumetric solids content (Figure 13). The data confirm
that warm, ice-rich permafrost under higher normal
stresses will exhibit greater steady-state strain
rates.
Within many of the rock glaciers examined to date,
significantly higher local strain rates occur within a
thin ‘shear’ zone, which may be 1–3 m or more in
thickness. The depth of this zone and the magnitude of
strain rate are strongly dependent on the internal
structure (% ice, % soil, including typical particle
sizes) and local temperature, which is often very close
to 08C (e.g. Haeberli, 1985; Wagner, 1992; Haeberli
et al., 1998; Arenson et al., 2002). Usually there is
Copyright # 2006 John Wiley & Sons, Ltd.
205
virtually no deformation below this shear zone and
relatively little creep deformation above it.
Strain rates in shear zones within 20 m of the
surface can be influenced by seasonal temperatures, so
that seasonal changes in the surface velocity have been
detected for a rock glacier with changes exceeding
0.2 m/a, which reflect temperature variations over the
seasonally affected zone (Arenson et al., 2002; Kääb
et al., 2003). Seasonal changes in ground temperature
are not only generated from changes in surface air
temperature, but also from convective heat flow within
permeable taliks or layers at the permafrost base.
Shear deformation is generally accompanied by either
dilation or contraction depending on the structure of
the permafrost, but certainly does not occur at constant
volume as is often assumed in rock glacier dynamics.
Structural hindrance and hence dilation will occur if
the debris content exceeds about 35% (e.g. Goughnour
and Andersland, 1968; Ting et al., 1983) and the
matrix is virtually air-free, whereas contraction will
occur if there is a significant amount of air, for
example > 5% present, which is not untypical for
some rock glaciers (Wagner, 1990; Arenson et al.,
2003a, 2004; Yasufuku et al., 2003). This will affect
the volume of permafrost lying in and above the shear
zone and the predictions of vertical and horizontal
strain.
Numerical modelling of the four-phase interaction
between solids, water, ice and air may be achieved
using discrete elements (Arenson and Sego, 2005) in
the future but the constitutive models will need to be
developed at a particulate level while computer
capacity must increase by another order of magnitude
to make this feasible. Attempts to back-fit average
creep parameters to a time series of deformation data,
for example by treating the rock glacier as a
homogeneous medium (e.g. Wagner, 1992), gave a
useful perspective at the global scale but inevitably
failed to represent the detailed deformation field that is
necessary for in-depth understanding of rock glacier
dynamics and for risk assessments in combination
with permafrost degradation.
CONCLUSIONS AND RECOMMENDATIONS
Research on permafrost creep and the kinematics
of active rock glaciers in cold mountainous areas
has advanced rapidly during recent decades and
continues to expand. This results from the application
of sophisticated technology in comprehensive field
and laboratory experiments, and especially from
increasing cooperation among international specialists
in geomorphology, microclimatology, glaciology,
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
206 W. Haeberli et al.
Figure 13 Triaxial creep tests at confining pressure s3 ¼ 200 0.4 kPa for different volumetric ice contents, wi, and temperatures,
T; samples 1, 2, 4, 9 and 11 were artificially frozen; (A) wi ¼ 15–33%; (B) wi ¼ 61–78%; and (C) wi ¼ 85–100% (Arenson and Springman,
2005b).
geophysics, geodesy/photogrammetry and engineering. The resulting information provides a clearer
picture concerning the basic materials, processes,
phenomena and spatio-temporal scales involved in the
motion of rock glaciers and associated features that
also result from the gravitationally-induced deformation of subsurface ice.
Thermal conditions remain the fundamental
boundary condition at the surface and define spatial
scales (extent, distribution patterns and thickness of
Copyright # 2006 John Wiley & Sons, Ltd.
permafrost) as well as temporal dimensions (formation, preservation and melting of ground ice under
conditions of Holocene climates and recent atmospheric warming). Surface temperatures affect the
permafrost through a highly heterogeneous active
layer commonly exhibiting large, wide-open voids,
making conductive, convective and advective energy
fluxes extremely complex and strongly dependent on
snow-cover characteristics. The coarse blocks of this
active layer, with its special thermal conditions, result
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
Permafrost Creep and Rock Glacier Dynamics
from weathering, rockfall activity and grain-size
sorting at the rock headwall of the creeping ice/rock
mixtures. Ground ice is most likely to be polygenetic
and generally far exceeds the pore space of the talus
within which it forms. Massive subsurface ice
sometimes occurs and may have formed below the
surface or could have originated as surface ice from
avalanche cones, perennial snowbanks, glacierets or
small (mainly cirque) glaciers, and was subsequently
buried at negative ground temperatures in debris either
derived from valley walls or accumulating on the rock
glacier surface from ablating ice. The coherent pattern
of the surface motion evident in the continuous surface
structures commonly seen on rock glaciers reflects
high ground-ice contents. Such motion continues over
millennia as indicated by high-precision geodetic/
photogrammetric surveys and radiocarbon dating. Due
to the striking heterogeneity of the material and
differences in ice content and internal architecture
between rock glaciers, however, the rheology generally remains difficult to define universally. In addition,
a number of aspects of the flow and evolution of
creeping permafrost and rock glaciers are poorly
understood; these include short-term velocity fluctuations, the occurrence of thin but effective shear zones,
extrusion flow and time-dependent volume changes.
There is a clear need for integrated numerical
modelling of rock glacier development and evolution,
as well as dynamics. However, this will require better
characterisation of rock glacier composition and
internal structure, and better understanding and more
rigorous modelling of specific processes. Such
quantitative holistic treatment is a difficult challenge
and must be based on diverse data available from
photogrammetry, data logging, borehole observations
and laboratory experiments. It should include and
combine partial models for:
(a) rock weathering, snow avalanches and rockfall
events with grain-size sorting and mass wasting
on scree slopes;
(b) freezing processes and ice formation in sub-zero
scree, with abundant fine material as well as
coarse blocks;
(c) coupled thermohydro-mechanical modelling of
creep and failure processes in frozen rock debris,
including the entire range of particle sizes from
fine to coarse, and the entire range of ice contents
from little or no ice in a dense interlocking matrix
to massive subsurface ice with highly dispersed
debris and significant air content;
(d) kinematics of non-isotropic, heterogeneous and
layered, ice-rich permafrost on slopes with long
transport paths for coarse surface material from
Copyright # 2006 John Wiley & Sons, Ltd.
207
the headwall to the front and, in some cases
subsequent re-incorporation into the advancing
rock glacier causing corresponding age inversion
at depth; and
(e) dynamics of rock glaciers, which include spatial
and temporal variations in velocity that are manifested in the ridges, furrows and other surface
structures typical of rock glaciers, as well as their
down-valley motion.
It was the clear goal of the IPA Task Force to
concentrate on purely periglacial features and
processes, which predominate in most regions. As a
next step, however, interactions and feedbacks
between snow, glaciers and frozen scree on mountain
slopes should be considered in order to understand
more complex landforms. In particular, the question of
a gradation or continuum between rock glaciers and
debris-covered glaciers should be examined in view of
ground thermal conditions and the common contrast in
surface character of these features: the transverse
ridges, deep furrows and the steep, near-angle of
repose terminus, typical of perennially frozen rock
glaciers and the ablation pits and lakes typical of
debris-covered glaciers.
Future studies of rock glaciers should be closely
linked to other aspects of permafrost science, and
should benefit from major recent advances in understanding the physical and chemical processes that
operate in all permafrost regions (e.g. Hallet et al.,
2004). These have resulted from diverse approaches
that range from examining fundamental periglacial
phenomena at the microscopic scale to recognising the
importance of permafrost processes on climate change
at the global scale. We firmly believe that significant
advance in understanding most periglacial phenomena
requires the close integration of field and theoretical
studies, and will greatly benefit from ideas and
techniques in other disciplines within the geosciences
and beyond. Much is to be gained from crossfertilisation between scientists and engineers as well.
For example, the transfer of energy through a surface
layer of coarse rock debris is central to understanding
the thermal regime of rock glaciers. It is, however,
poorly understood and little has yet been gained from
the extensive studies of a very similar engineering
problem associated with the use of artificial rock cover
to avoid warming or melting permafrost. Another
example is using what has been learned through
extensive empiricism in engineering studies about
frost resistance of concrete and other porous materials
to provide a basis for further studies of the physics of
frost weathering, which is one of the controls on the
debris supply to rock glaciers.
Permafrost and Periglac. Process., 17: 189–214 (2006)
DOI: 10.1002/ppp
208 W. Haeberli et al.
Future studies should also appeal to the broad
scientific audience that is now aware and interested in
permafrost as a critical element of the Earth’s
cryosphere particularly in light of its vulnerability
to atmospheric warming and potential feedbacks to
global climate. In addition, they should capitalise on
the growing interest from the engineering and
planetary communities, respectively, in geotechnical
issues characteristic of cold regions and in diverse
aspects of permafrost on Mars and other icy planets
(e.g. Whalley and Azizi, 2003; Head et al., 2005).
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