Content uploaded by Reginald L. Hermanns
Author content
All content in this area was uploaded by Reginald L. Hermanns on Dec 09, 2015
Content may be subject to copyright.
59
6 Rapid rock-slope failures
reginald l. hermanns and o d d v a r longva
Our life is short. The memory of mankind as a whole is poor. The few mountain collapses that we experience in our lifetime leave us
with the impression that these collapses are very exceptional, extraordinary events. However, that is not the case. Mountain collapses
are normal events in the mountains, especially in the high mountains, where they have an important natural role to play in helping
to form and shape the mountains; a process that continues relentlessly and steadily. In the mountains we have to expect mountain
collapses from time to time, from place to place.
(A. Heim, 1932)
ABSTRACT
Large catastrophic rock-slope failures are a serious threat to
humans because
they cannot be controlled by any physical measures
•
they can be highly mobile and reach areas several kilometers
•
from the source
they can trigger damaging secondary effects.
•
Nowadays a concerted effort is being made, using advanced
remote sensing tools, to identify and monitor slopes that might
fail catastrophically. Small accelerations that might be precur-
sors of failure can be used as indicators to help in emergency
planning. Characterization of the structure of the rock mass
of a potentially unstable slope is an essential step in assessing
the likelihood of catastrophic failure. Geologic records indicate
that rapid slope failures are not distributed evenly in time due,
in part, to climate variability. Analyses of the deposits of large,
rapid rock-slope failures in the central Andes indicate that the
proximity to active faults and the type of deformation found
on these structures control landslide distribution and size.
Many catastrophic rock-slope failures are triggered by strong
crustal earthquakes. Historic data indicate that subduction
earthquakes are less effective in triggering large landslides than
crustal earthquakes.
6.1 INTRODUCTION
The topic of this chapter is catastrophic rock-slope failures. We
apply the term “catastrophic” to rock-slope failures that involve
substantial fragmentation of the rock mass during runout and
that impact an area larger than that of a rockfall (shadow angle
of ca. 28–32° from the source). Most catastrophic rock-slope
failures are larger than 106 m3, and failure involves the devel-
opment of a continuous rupture plane between the underlying
rock mass and the sliding rock body. We exclude collapses of
volcanic edices because this topic is covered in Chapter 4.
We rst describe the failure process and summarize the con-
ditions that indicate the possibility of an imminent failure. We
then describe rock mass structures that control the failure pro-
cess and depositional features that are characteristic of pre-
historic catastrophic rock-slope failures. Next we summarize
several chronological studies of large catastrophic landslides in
the central Andes and Norway that have implications for cli-
matic conditioning of slopes for failure. Finally we discuss the
inuence of the tectonic setting based on a systematic analysis
of rock-slope failures in the central Andes.
6.2 OVERVIEW OF ROCK-SLOPE FAILURES AND
HAZARD
Rock-slope failures of 106–107 m3 have occurred nearly every
year somewhere on Earth in the past century, and there have
been two or more such failures in some years. Rock-slope failures
larger than 1 km3, on the other hand, are much less frequent; the
only recent events of this size described in the scientic litera-
ture are the 1911 Usoi landslide in Tajikistan (2.2 km3; Schuster
and Alford, 2004) and the 1974 Mayunmarca landslide in Peru
(Kojan and Hutchinson, 1978). Even larger events, however,
have been documented in the geologic record, some up to sev-
eral tens of cubic kilometers in size (e.g., Saidmarreh, Iran;
Harrison and Falcon, 1934; and Lluta Valley, Chile; Wörner
et al., 2002).
Some deposits of large rock-slope failures are today densely
populated, for example the Flims landslide deposit, which
occurred 8200 14C years BP (von Poschinger et al., 2006),
Landslides: Types, Mechanisms and Modeling, ed. John J. Clague and Douglas Stead. Published by Cambridge University Press.
© Cambridge University Press 2012.
Hermanns and Longva 60
and some cities with more than 1 million inhabitants are par-
tially built on young landslide deposits. For example, landslide
deposits only about 10,000 years old cover an area of 60 km2
within the incorporated area of La Paz, Bolivia (Dobrovolny,
1962), and relatively young rockslide deposits are present within
Caracas, Venezuela (Ferrer, 1999).
With sufcient relief (150 m, see Keefer, 1984; 400 m, see
Hermanns and Strecker, 1999), large catastrophic rock-slope
failures achieve high velocities (5–>100 m s−1) in short travel
distances; thus evacuation of the runout area, without prior
warning of failure, is impossible. Catastrophes involving such
landslides include the Huascaran landslide, which destroyed the
town of Yungay in Peru in 1970 (Plafker and Ericksen, 1978)
and a rock-slope failure that destroyed several villages dur-
ing the Khait earthquake in Tajikistan in 1949 (Evans et al.,
2009a).
Most catastrophic rock-slope failures are “rock avalanches.”
This term was coined by Hsü (1975), based on Heim’s (1932)
description of the phenomenon. Heim used the German terms
“Bergsturz,” “Trümmerstrom,” and “Sturzstrom” for streams
of rapidly moving debris resulting from the disintegration of a
failed large rock mass. The streaming behavior generally devel-
ops only when the landslide is larger than 106 m3. Synonyms for
rock avalanche include rockfall avalanche, rockfall-generated
debris stream, and sturzstrom. Runout distances of rock ava-
lanches commonly exceed several kilometers; their high mobil-
ity may be evidenced by high run-up on opposite valley slopes,
which is related to the volume of the initial failed rock mass
(Scheidegger, 1961), and superelevation of debris at bends in the
ow path (Nicoletti and Sorriso-Valvo, 1991). Mobility can be
enhanced by the entrainment of saturated soil material, snow,
or ice along the ow path (Hungr and Evans, 2004). Flow vel-
ocities can be calculated from run-up and superelevation using
the equations summarized in Crandell and Fahnestock (1965).
Rock avalanches are not the only landslides that t our def-
inition of catastrophic rock-slope failures. Others include
rock–ice avalanches and rockslides or rockfalls, including those
that enter a water body and trigger a tsunami (see Chapter
10). Rock–ice avalanches can involve a range of ratios of rock
and ice. For example, a rock–ice avalanche on November 29,
1987 at Estero Parraguirre, Chile, was initiated by the failure
of 6 × 106 m3 of rock, but an additional 9 × 106 m3 of deb-
ris and ice were entrained from a glacier onto which the rock
mass fell (Hauser, 2002). The 1970 Cerro Huascaran rock-slope
failure had an initial volume of 6.5 × 106 m3 of rock and ice,
but entrained 43 × 106 m3 snow, ice, and debris along its path
(Evans et al., 2009b). Both events transformed into debris ows
that traveled, respectively, 57 and 180 km. On September 20,
2002, 18.5–27 × 106 m3 of rock and ice dropped from a steep
rock slope onto Kolka Glacier in the Russian Caucasus. The
rock–ice avalanche removed the lower part of the glacier and
traveled down the Genaldon Valley for 20 km, at which point
it transformed into a debris ow that traveled to the entrance
of the Karmadon Gorge, killing 140 people and causing wide-
spread destruction (Huggel et al., 2005).
Secondary effects extend the area of catastrophic rock-slope
failures. The most important secondary effects are
• damming of river valleys, resulting in upstream ooding
behind the debris dam and potential downstream ooding
from overtopping and breaching of the dam
• landslide-triggered displacement waves (Costa and Schuster,
1988; Clague and Evans, 1994; Tappin, 2010; Evans et al.,
2011).
An assessment of the hazard posed by an unstable rock slope
must include the possible impacts of secondary phenomena.
The probability of failure is difcult to determine. Regional-
scale measures of failure probability include the annual fre-
quency of landslides of a particular size per 10,000 km2 of
mountainous terrain (Hungr, 2006) or the number of events per
thousand years per region (see below). Such measures, however,
are not helpful at the local scale, where, for example, it may be
necessary to assess the probability that a particular slope will
fail (Aa et al., 2007; Hermanns et al., 2012). In such situations,
it is essential to search for archival and other historic data on the
slope of concern (Glastonbury and Fell, 2010), as well as local
evidence of past instability such as ground cracks, recent rock-
falls, and hydrologic changes on or near the slope. Monitoring
of slope deformation may provide some warning of approach-
ing failure (Crosta and Agliardi, 2003 and references therein),
although instances where such data have been successfully used
to predict catastrophic failure are few. As a rule of a thumb, a
protracted acceleration of deformation rates is a clear sign that
failure is imminent (Fig. 6.1).
Large catastrophic rock-slope failures can rarely be prevented
or mitigated. Risk can be reduced, however, by
recognizing slopes that potentially might fail suddenly
•
slope monitoring and the formulation of warning and emer-
•
gency evacuation plans.
Two assumptions underlie these measures: rst, rock-slope
deformation that may lead to failure can be detected and
Time
Displacement rate
Failure
Pre-failure
stage
Post-failure
stage
Fig. 6.1. Schematic diagram showing different stages of slope move-
ment leading to catastrophic rock-slope failure. The solid line is for a
nonseismic failure; the dashed line is for a seismically triggered failure.
The nomenclature follows Leroueil et al. (1996).
Rapid rock-slope failures 61
monitored; and second, the deformation accelerates prior to
failure. The rst assumption is valid, given recent developments
in satellite and ground-based remote sensing, although the
application of these tools on a continuous basis is expensive.
The second assumption may not be true. In seismically active
regions, for example, slope deformation accelerates to cata-
strophic failure within seconds (Fig. 6.1). In these areas, a bet-
ter option than slope monitoring is zoning and restrictions on
land use based on numerical modeling of rock-slope stability
and runout (e.g., Welkner et al., 2010).
Regional studies have shown that catastrophic rock-slope fail-
ures do not occur uniformly in space (Abele, 1974; Hermanns
and Strecker, 1999) or time (Trauth et al., 2000). Rather, they
are controlled by lithological and structural conditions, their
setting with respect to active faults, and climate. We focus on
these controls in the following sections.
6.3 RECOGNITION OF TYPES OF ROCK-SLOPE
FAILURES AND THEIR DEPOSITS
Deposits of catastrophic rock-slope failures can be recognized,
depending on their climatic setting, for years, decades, centuries,
or even millennia. Rock avalanches leave characteristic sheets of
debris ranging from several meters to several hundred meters
thick with sharp margins. When the landslide is unconstrained
by topography, its deposits are relatively thin lobate sheets with
lateral levees, frontal rims, and blocks meters to tens of meters
in size at the surface. The coarse carapace typically overlies mas-
sive angular debris ranging in particle size through many orders
of magnitude down to sub-micron size. Many of the clasts are
densely fractured and display what have been termed “jigsaw
texture” (Yarnold, 1993). Rock-avalanche deposits derived from
a single rock type are monolithic. In cases where two or more
lithologies are present in the source area, the lowest lithologies
in the scarp are concentrated at the outer margin of the deposit,
while the highest lithologies in the scarp are concentrated in
proximal positions. Failures involving both rock and ice are
more difcult to recognize because the initial failed mass com-
monly transforms into a debris ow or mudow along its path
(Hauser 2002; Huggel et al., 2005; Evans et al., 2009b). Fauqué
et al. (2009) described deposits of late Pleistocene and early
Holocene rock–ice avalanches from the south face of Cerro
Aconcagua in South America. The deposits have long been mis-
interpreted as glacial deposits due to their inclusion of a variety
of lithologies, including glacial and alluvial materials entrained
along the ow path. The different components in the deposit
were not completely mixed; instead, zones consisting mainly of
rock-avalanche material are in contact with zones of reworked
glacial and alluvial deposits.
Prehistoric catastrophic landslides into a water body are rela-
tively easily identied using bathymetric and seismic data. The
blocky deposit of such a landslide will overlie older marine or
lacustrine sediments. If the background sedimentation rate is
sufciently high, the deposit will be buried and may be visible
only as a layer of chaotic reectors in a seismic reection record
(Strasser et al., 2006; Longva et al., 2009).
Identifying slopes that might fail catastrophically in the
future is a more complex task. Historic data, however, indicate
that large rock-slope failures are preceded by slope deformation
(Fig. 6.1; Eisbacher and Clague, 1984; Furseth, 2006), which
may make it possible to identify a potential rock-slope failure
before it happens. A systematic national program to character-
ize such slopes has been initiated in Norway (Hermanns et al.,
2012). Mapping rock slopes that are slowly deforming is the rst
step in locating the sites of future catastrophic failures. An engin-
eering geologist can then complete a detailed structural analysis
of a specic, slowly deforming slope to determine whether cata-
strophic failure is kinematically possible. Glastonbury and Fell
(2010) dened and illustrated, using schematic cross-sections,
eight structural environments in which catastrophic slope fail-
ures have occurred in the past (Fig. 6.2A–H). In Norway, struc-
tural conditions favoring failure in igneous or metamorphic
rocks have been identied (Fig. 6.2I, J; Braathen et al., 2004).
For example, two sets of joints allow both toppling and slid-
ing (Fig. 6.2J), resulting in a type of landslide referred to as a
slide–topple. An investigator should also document how much
rock deformation has already taken place and whether and how
often catastrophic failures have occurred at the site in the recent
past (Guglielmi and Cappa, 2010; Sanchez et al., 2010).
6.4 TEMPORAL ROCKSLIDE DISTRIBUTION AND
CLIMATE CHANGE
The impact of climate warming on slope stability, mainly
through thawing of alpine permafrost and debuttressing of
glacially oversteepened, unstable rock slopes due to glacier
retreat, has been the subject of much discussion (Abele, 1974;
Evans and Clague, 1994; Noetzli et al., 2007; Fischer et al.,
2010; Huggel et al., 2010). Except for the historic period, how-
ever, this causative relation can only be demonstrated by dat-
ing large numbers of catastrophic rock-slope failures. Based
on a compilation of data from the European Alps, Abele
(1997) concluded that most large catastrophic rock-slope fail-
ures occurred in late-glacial time. Some, however, occurred
much later, implying a delay in failure due to progressive
rock mass weakening. Most large landslides in the Scottish
Highlands appear to be of late-glacial or early-Holocene age
and have been attributed to glacial steepening and seismic
activity caused by rapid glacio-isostatic rebound (Ballantyne,
1997). Some large landslides in this area, however, happened
several thousand years after deglaciation, implying that pro-
gressive stress release and joint propagation, and perhaps
other time-dependent factors, have played a role (Ballantyne
et al., 1998). Cruden and Hu (1993) compiled the ages of large
landslides in the Canadian Rocky Mountains and proposed an
exhaustion model to explain the decrease in activity through
the Holocene. The premises of this model are that there are a
nite number of potential failure sites that are conditioned by
Hermanns and Longva 62
glaciation and that each of these sites fails only once. These
premises have been proven to be wrong because more than one
slope failure can occur at a single site and because completely
new instabilities can be created through the gradual fatigue of
rock masses (Hermanns et al., 2006 and references therein; Aa
et al., 2007).
Hermanns et al. (2000) presented a systematic regional
inventory of catastrophic rock-slope failures in northwest
Large rock glide
AB
CD
EF
GH
IJ
Rough translation slide
Planar translational slide
Toe buckling translational slide
Curved compound slide
Biplanar compound slide
Irregular compound slide
Toe buttress compund slide
Rock fall slide Slide topple
Fig. 6.2. Schematic diagrams show-
ing different distributions of geologic
discontinuities controlling large rock-
slope failures (after Braathen et al.,
2004; Glastonbury and Fell, 2010).
Rapid rock-slope failures 63
Argentina. The ages of 25 of the 55 mapped deposits were
determined through 14C dating, terrestrial cosmogenic nuclide
dating, pedological methods, and tephrochronology. The
ages of an additional eight deposits were reported later by
Hermanns and Schellenberger (2008). None of the 55 land-
slides in the dataset has a source on glaciated slopes, even
though alpine glaciers reached down to 4300 m asl in the east-
ernmost ranges during the late Pleistocene (Haselton et al.,
2002). Instead they have sources below 3500 m asl in deeply
incised valleys or on slopes several tens of kilometers away
from trunk streams. All landslide deposits in the deeply incised
valleys date to the late Pleistocene or Holocene, whereas those
on the slopes bordering the watershed are all more than
100,000 years old (Hermanns et al., 2000). Landslides in the
incised valleys occurred mainly during periods of wetter cli-
mate; enhanced runoff and lateral erosion of valley oors
were likely the main causes of this landsliding (Trauth et al.,
2000; Hermanns and Schellenberger, 2008). However, some
large landslides apparently occurred during what are thought
to have been dry phases of the Holocene. These exceptions
occur near active faults and may have been seismically trig-
gered (Hermanns and Schellenberger, 2008).
Deposits of 22 catastrophic landslides were identied and
characterized during a systematic study of a major trafc corri-
dor 800 km south of the region discussed above, within the cen-
tral Andes (Fig. 6.3; Rosas et al., 2007). Valley glaciers extend
down to 3800 m asl in this region today, and reached several
hundred meters lower at the maximum of the last glaciation
(Fauqué et al., 2009). Twenty of the landslides occurred in ice-
free environments; the other two fell onto a valley glacier (Fig.
6.3; Fauqué et al., 2009, Welkner et al., 2010). The ages of 21
of these deposits were determined by terrestrial cosmogenic
nuclide dating or radiocarbon dating of plant detritus recov-
ered from lacustrine deposits behind the landslide barriers
(Fauqué et al., 2008a, 2008b, 2009; Rosas et al., 2008; Welkner
et al., 2010). The landslides occur in two geomorphic settings:
(1) fault-controlled valleys that have never been glaciated, and
(2) deeply incised valleys that contained glaciers at the Last
Glacial Maximum or that drained glaciated valleys. Only one
of the 13 dated deposits within the glaciated and trunk valleys
dates to a time when glaciers were more extensive than today
(Fauqué et al., 2008b, 2009). As shown in Figure 6.3, 10 of the
13 are late-glacial to early-Holocene age, and 2 are younger.
The deposits of the two youngest landslides yielded similar
ages and are close to one another in a valley with a potentially
active fault.
Many catastrophic rock-slope failures have occurred in the
deeply incised fjords of Norway. They include disasters in
Loen in 1905 and 1936, when large blocks of rock fell into
Lake Loenvatnet from the 1493-m-high mountain Ravnefjell
(Hermanns et al., 2006 and references therein). The two land-
slides generated tsunamis with maximum wave heights of 40
and 70 m that, respectively, killed 61 and 74 people. A similar
landslide in 1934 in Tafjord triggered a tsunami that killed 40
people (Hermanns et al., 2006 and references therein).
Prehistoric rockslide and rock-avalanche deposits have been
found in fjords and valleys throughout Norway; they are best
documented in Møre and Romsdal County on the west coast
and in Troms County in northern Norway (Blikra et al., 2006;
Furseth, 2006). Events in these areas range in age from late-
glacial to late-Holocene. However, as seen in Canada, Scotland,
and in the Alps, the largest landslides apparently occurred
shortly after deglaciation. Detailed mapping and geophysical
surveying of Storfjorden revealed deposits of 107 landslides,
the largest of which occurred during deglaciation (Table 6.1;
Fig. 6.4; Longva et al., 2009). Six of the 107 landslides date
to between 12,500 and 11,000 years ago and have an average
volume of 59 million m3; three of these six have volumes of
100–200 million m3. Landslides were most frequent during the
Younger Dryas cold period between 11,000 and 10,000 14C
years BP and at the beginning of the Holocene. Since about
9000 years ago, the frequency of large landslides in Storfjorden
has been about ve per thousand years, with a slightly increased
frequency between 5000 and 1000 years ago. The volumes of
the landslides have varied over the Holocene, and there may be
both climatic and tectonic signals in the frequency distribution.
Norway lies at the western margin of the Baltic Shield, which is
thought to be a tectonically passive margin. Some researchers,
however, have reported evidence for high tectonic activity at the
end of the Pleistocene and into the Holocene in Scandinavia
(Mörner, 1996; Bungum et al., 2005). Clusters of rock-slope and
soil failures during the late-glacial period have been attributed
to earthquakes induced by rapid isostatic rebound (Bøe et al.,
2004; Blikra et al. 2006). Rapid contemporaneous uctuations
in climate may also have contributed to frequent slope failures
at that time. The Holocene climatic optimum in Norway dates
to 8000–5000 BP (Hafsten 1986). During that period, the fre-
quency of catastrophic rock-slope failures was the same as earl-
ier and later, but the events – on average – were of smaller size.
Climate began to deteriorate about 5000 years ago, and since
then the average size of landslides has increased. Systematic
regional studies on the temporal distribution of catastrophic
slope failures have been carried out in other parts of the world.
For example, Bookhagen et al. (2005) and Dortch et al. (2009)
dated 16 catastrophic slope failures in the Himalaya of north-
ern India. Fourteen of these events occurred during two peri-
ods, the rst a period of increased monsoon activity 40,000 to
30,000 years ago and the second during the most intense mon-
soon phase of the Holocene, from 8400 to 7200 years ago. On
the other hand, Hewitt et al. (2011) concluded that earthquakes
may have played a greater role than climate in causing the large
Holocene rock avalanches in the Karakoram Himalaya that
they dated. Soldati et al. (2004) documented a cluster of large
landslides in the European Alps between about 11,500 and
8500 years ago; many of them were reactivated during the Sub-
Boreal period, about 5800–2000 years ago. Prager et al. (2008)
found that 12 of 14 large (>108 m3) rockslides in the Austrian
Alps and surrounding area are Holocene in age, with a minor
cluster in the early Holocene and about 4200–3000 years ago in
the Sub-Boreal period.
Hermanns and Longva 64
6.5 CATASTROPHIC LANDSLIDES AND
NEOTECTONICS
Abele (1974) discussed the importance of tectonic activity as
a preparatory mechanism for landslides. He suggested that
intense fracturing of the rock mass adjacent to faults creates
conditions conducive to slope failure, and that earthquakes trig-
ger landslides. Many case studies from around the world have
demonstrated the importance of tectonic activity both as a con-
ditioning factor and a trigger mechanism.
Tectonic activity contributes to slope instability in three
ways. First, it creates zones of weak rock along the fault. All
types of faulting break down the rock mass along the fault
trace (Brideau et al., 2005, 2009), and folding can produce
extension cracks along the hinge zone of anticlines. Second,
tectonic activity, operating over long periods, produces relief.
Normal and reverse faulting and folding are most efcient in
generating important relief (Fig. 6.5). Strike–slip faults, how-
ever, can also generate relief in transpressional zones at kinks
in the fault trace. Third, tectonic activity can translate inher-
ited structures within the rock mass into positions that are more
favorable to failure, for example by producing inclined bedded
planes. Folding changes the orientation of discontinuities over
the entire structure. In contrast, the effect is more localized in
the case of reverse and normal faulting; unfavorably oriented
discontinuities in the rock mass can be exposed in the hang-
ing wall of a normal or reverse fault. Rockslides capitalizing
on these exposed inherited discontinuities have been reported
from around the world (Eisbacher and Clague, 1984; Hermanns
and Strecker, 1999; Martino et al., 2004, von Poschinger et al.,
2006).
The association of normal faulting and slope collapse has
been documented at many sites (Dobrovolny, 1962; Martino
et al., 2004; Brideau et al., 2005; Redeld and Osmundsen,
2009), but the number of catastrophic rock-slope failures
related to those structures is small. In reverse fault settings
(Fig. 6.5B), the situation is generally different, and multiple
large landslides have been documented along these structures
(Hermanns and Strecker, 1999; Jackson, 2002; Penna et al.,
2011). Reverse faulting also produces relief contrasts (Fig.
6.5B) and weakening of rocks along the fault. Thus, in the long
term, movement along the fault leads to oversteepening and
slope collapse. Some of the world’s largest nonvolcanic rock-
slides have occurred in northern Chile in this tectonic setting
(Wörner et al., 2002). Hermanns et al. (2001) showed that activ-
ity on reverse faults in a mountain range in Argentina caused
slope steepening until about 150,000 years ago, when the locus
of deformation shifted away from the range front to the fore-
land. Between about 400,000 and 150,000 years ago, one large
rock avalanche occurred, on average, once every 30,000 years.
In contrast, there have been no large rock avalanches at the
range front during the past 150,000 years. Tectonically inactive
mountain fronts built of the same rocks and with the same relief
lack any evidence of large rock-slope collapses (Hermanns and
Cerro Aconcagua
6962 m
R
í
o
A
c
o
n
c
a
g
u
a
R
í
o
L
a
s
C
u
e
v
a
s
R
í
o
M
e
n
d
o
z
a
Argentina
Chile
Rock-avalanche deposit
Rock-ice avalanche deposit
14,300±1,000 yr 9,600±700 yr
>13.670±220 yr
11,200±1,400 yr
>18,660±330 yr
> 8,640±200 yr
> 9,640±130 yr
46,600±3,300 yr
110,000±22,600 yr
192,000±42,300 yr 112,900±14,600 yr
159,500±17,700 yr
12,600 ± 950 yr
4,500 ± 300 yr
4,100±500 yr
14,800±2,100 yr
11,500±800 yr
9,000±1,400 yr
133,700±21,300 yr
Atlantic Ocean
Argentina
Pacific Ocean
Fig. 6.3. Satellite image of part of the Andes in the vicinity of Cerro Aconcagua, showing the distribution and ages of deposits of catastrophic
rock-slope failures. Ages are from Fauqué et al. (2008a, 2008b, 2009), Rosas et al. (2008), and Welkner et al. (2010), and include (1) mean 36Cl
exposure ages (black) of 2–6 samples adjusted for an erosion rate of 2.2 mm ka−1 (italicized number is an age based on a single sample), and
(2) calibrated radiocarbon ages (gray) of plant material recovered from lacustrine deposits behind rockslide barriers. Arrow in inset shows loca-
tion of the study area.
Rapid rock-slope failures 65
Strecker, 1999). Strecker and Marrett (1999) attribute the high
concentration of rock avalanches in this area to a reorgan-
ization of tectonic deformation in the Neogene, when former
strike–slip faults were reactivated as reverse faults. Catastrophic
rock-slope failures occurred also along other strike–slip faults.
Sepulveda et al. (2010) documented catastrophic rockslides
along a strike–slip fault in the Patagonian Andes of Chile, but
at those sites the fault intersected deeply incised valleys and the
relief required for slope failure was not produced by the fault
itself (Fig. 6.5C).
Table 6.1. Overview of temporal distribution of number and volume of 107 rock avalanches in Storfjorden, Norway.
Period (14C ka BP)a12.5–11 11–10 10–9 9–8 8–5 5–1 1–0
Number of events 6 19 26 5 16 30 5
Events per 1000 years 4 19 26 5 5 8 5
Total volume (Mm3) 354 168 31.7 6.7 3.9 15.5 7.6
Average volume (Mm3) 59 8.8 1.2 1.3 0.2 0.5 1.5
Volume per 1000 years 236 168 31.7 6.7 1.3 3.9 7.6
a The ages of events are based on one radiocarbon-dated core and the regional seismic stratigraphy.
Fig. 6.4. Morphological and seismic interpretation, and estimated age of landslide deposits in Geirangerfjorden, Norway. (A) Shaded relief image;
note that deposits of younger landslides have a sharper, more irregular, morphology than those of older, deeper, buried ones. The number and size
of each arrow indicates the source, age, and size of the landslide. (B) Seismic lines (white) and depositional areas of landslides. (C) Interpretation
of seismic line. Letters X and Z in Parts A and B mark the same locations.
Hermanns and Longva 66
In most normal and reverse fault settings, it is difcult to sep-
arate the roles of tectonically induced rock deformation, tec-
tonic oversteepening of slopes, and slope steepening by erosion
along valleys in conditioning slopes for failure. A region where
tectonic oversteepening plays a minor role is the transition
between the Central and the Patagonian Andes in Argentina
(Penna et al., 2011). The main factors responsible for landslides
in this region are erosion along valleys and tectonically induced
rock deformation. A total of 19 large landslides have occurred
in volcanic and volcaniclastic rocks of Plio-Pleistocene age
underlying a plateau that was deformed in the Quaternary by
faulting and folding (Fig. 6.6). Local relief of 15–400 m devel-
oped on the plateau, but none of the landslides were sourced
on the faults or folds responsible for this relief. The absence of
an association of the local relief and landslides suggests that
tectonic oversteepening of slopes is not responsible for the fail-
ures. Local erosional relief of 200–1200 m was created by rivers
and glaciers crossing the plateau, and all 19 rockslides occurred
in these valleys. Relief is greatest in valley sections eroded by
glaciers (Fig. 6.6), and about 80 percent of the landslide depos-
its occur there (Fig. 6.7). The evidence thus suggests that glacier
erosion and debuttressing were important factors in condition-
ing the slopes for failure. However, because all of the landslides
are more than 10,000 years younger than the time of maximum
glaciation in the area, glacial erosion and debuttressing were
not triggering factors, only conditioning ones. Neotectonically
induced rock deformation appears as important as glaciation:
more than 95 percent of the volume of the landslide depos-
its are localized along neotectonic structures. Some of the
volcanic and volcaniclastic rocks are intensely fractured, but
there are no large discontinuities that dip toward the valley and
might form signicant sliding planes. In the case of landslides
localized along neotectonic structures, more than 85 percent of
their volume is associated with Neogene folds; only 15 percent
is associated with faulted rock. Thus folding seems the more
efcient process for weakening rocks to the point that large-
scale landsliding can occur, at least in this part of the Andes
(Fig. 6.6).
Montandon (1933) suggested that catastrophic rock-slope
failures in the Alps were triggered by earthquakes and, since
then, numerous researchers have described specic events in
detail (Shreve, 1966; Plafker and Ericksen, 1978; Adams, 1981;
Jibson et al., 2006; Owen et al., 2008; Dai et al., 2011). Keefer
(1984) and Rodríguez et al. (1999) summarized observations on,
respectively, 40 and 36 earthquakes that triggered landslides.
Most of the earthquakes that triggered large rock avalanches
were crustal; only four were great subduction earthquakes.
This statistic agrees with evidence gathered in recent years that
shallow earthquakes in the continental crust commonly trigger
large landslides along or near the surface rupture (Jibson et al.,
2006; Sepulveda et al., 2010; Dai et al., 2011), whereas sub-
duction earthquakes only occasionally trigger large landslides.
For example, the 1970 Nevado Huascaran rock–ice avalanche
(Plafker and Ericksen, 1978) was the only historic catastrophic
rock-slope failure in the Cordillera Blanca of Peru to have been
triggered by a subduction earthquake. Cerro Huascaran is
A
B
C
D
Area of rock deformation
related to tectonic activity
C
Fig. 6.5. Schematic block diagrams showing the impact of tectonic
deformation on rock-slope stability. A simple example is shown here,
with horizontally bedded sedimentary rocks. The gray dotted lines
represent a fault plane (normal fault in A) or the projection of the fault
plane (reverse fault in B). (A) Normal faults produce local relief and
localized rock deformation. In this case, the fault zone is parallel to the
slope, and sliding can occur along the fault plane. (B) Reverse faults also
produce local relief and rock deformation. The area of rock deform-
ation can be especially large in the hanging wall of listric reverse faults,
where the rock mass is compressed and tilted. The fault plane is not a
possible sliding plane because it dips into the rock mass. Gray dashed
line depicts the orientation of the fault in the failed rock mass; arrows
indicate slope adjustment by collapse and erosion. (C) Strike–slip faults
can enhance local relief by offsetting sloping ground. More import-
antly, they deform rock along the fault plane; slope failures can occur
where the fault crosses a valley. The fault is too steep to be a sliding
plane. (D) Folding produces surface relief and can create dipping planes
of weakness along which sliding may occur. Rock deformation occurs
over a large area and is related to tilting and extension.
Rapid rock-slope failures 67
bordered by a normal fault that had been active in the Holocene
(Schwartz, 1988). Only eight years earlier, a similar rock–ice ava-
lanche happened on the same rock face, but without any obvi-
ous trigger (Plafker and Ericksen, 1978), and in 1946, a M7.3
crustal earthquake triggered several rock avalanches close to
its epicenter, about 60 km from Cerro Huascaran (Heim, 1949;
Kampherm et al., 2009). This example illustrates that the coin-
cidence of a catastrophic landslide deposit with a neotectonic
fault cannot be taken as proof of seismic triggering. Thus
landslide deposits should only be used to supplement other
independent evidence for earthquakes, such as fault offsets
and seismically induced soft-sediment deformation structures
(Hermanns and Niedermann, 2011). Condence in a seismic
trigger may be increased if slope stability models indicate that
the failed rock slope could not have been unstable under aseis-
mic conditions (Jibson, 2009).
Headscarp
Rock avalanche
Rotational slide or
toppling
Reverse fault
Direct fault Syncline
Anticline Lake
Cerro Moncol
Piche Moncol
Chacayco I
Lauquén Mallín I
Lauquén Mallín II
Trohunco lake
El Convento fault
Moncol anticline
Guañacos I
Cerro Guañacos
Chochoy Mallín
Guañacos II
El Convento
Chacayco II
Chacayco fault
Guañacos fault
Chochoy Mallín fault
Chochoy I
03 km
37º35’S
71º00’W 70º50’W
37º25’S
37º15’S
Picún Leo
Vilú Mallín anticline
21º 15º 12º 13º
45º
43º
6º
43º 48º
Maximum extent of ice
Laguna
Negra
Trocomán NW
Trocomán SE
Atlantic Ocean
Argentina
aciic Ocean
Fig. 6.6. Distribution of neotectonic structures and landslide deposits in plateau basalts in the transitional area between the Central and Patagonian
Andes. The map is draped over a digital elevation model; the plateau and high valleys are light gray and white; mountain tops and low valleys are
dark gray (modied from Penna et al., 2011). Arrow in inset shows location of the study area.
Hermanns and Longva 68
REFERENCES
Aa, A. R., Sjåstad, J., Sønstegaard, E. and Blikra, L. H. (2007). Chronology
of Holocene rock-avalanche deposits based on Schmidt-hammer,
relative dating and dust stratigraphy in nearby bog deposits, Vora,
inner Nordfjord, Norway. Holocene, 17, 955–964.
Abele, G. (1974). Bergstürze in den Alpen: Ihre Verbreitung, Morphologie
und Folgeerscheinungen. Munich: Deutscher und Österreichischer
Alpenverein.
Abele, G. (1997). Rockslide movement supported by the mobilization
of groundwater-saturated valley oor sediments. Zeitschrift für
Geomorphologie, 41, 1–20.
Adams, J. (1981). Earthquake-dammed lakes in New Zealand. Geology,
9, 215–219.
Ballantyne, C. K. (1997). Periglacial trimlines in the Scottish Highlands.
Quaternary International, 38–39, 119–136.
Ballantyne, C. K., Stone, J. O. and Field, L. K. (1998). Cosmogenic
Cl-36 dating of postglacial landsliding at the Storr, Isle of Skye,
Scotland. Holocene, 8, 347–351.
Blikra, L. H., Longva, O., Braathen, A. et al. (2006). Rock slope fail-
ures in Norwegian fjord areas: Examples, spatial distribu-
tion and temporal patterns. In Landslides from Massive Rock
Slope Failure. Proceedings of the NATO Advanced Research
Workshop on Massive Rock Slope Failure: New Models for
Hazard Assessment, Celano, Italy, 16–21 June 2002, ed. S. G.
Evans, G. Scarascia Mugnozza, A. Strom and R. L. Hermanns.
NATO Science Series IV, Earth and Environmental Sciences 49.
Dordrecht, Netherlands: Springer, pp. 475–496.
Bøe, R., Longva, O., Lepland, A. et al. (2004). Postglacial mass move-
ments and their causes in fjords and lakes in western Norway.
Norwegian Journal of Geology, 84, 35–55.
Bookhagen, B., Thiede, R. and Strecker, M. R. (2005). Late Quaternary
intensied monsoon phases control landscape evolution in the
northwest Himalaya. Geology, 33, 149–152.
Braathen, A., Blikra, L. H., Berg, S. S. and Karlsen, F. (2004). Rock-
slope failures in Norway: Type, geometry, deformation mecha-
nisms and stability. Norsk Geologisk Tidsskrift, 84, 67–88.
Brideau, M., Stead, D., Kinakin, D. and Fecova, K. (2005). Inuence
of tectonic structures on the Hope Slide, British Columbia,
Canada. Engineering Geology, 80, 242–259.
Brideau, M., Yan, M. and Stead, D. (2009). The role of tectonic damage
and brittle rock fracture in the development of large rock slope
failures, Geomorphology, 103, 30–49.
Bungum, H., Lindholm, C. and Faleide, J. I. (2005). Postglacial seis-
micity offshore mid-Norway with emphasis on spatio-tempo-
ral-magnitudal variations. Marine and Petroleum Geology, 22,
137–148.
Clague, J. J. and Evans, S. G. 1994. Formation and Failure of Natural
Dams in the Canadian Cordillera. Geological Survey of Canada,
Bulletin 464.
Costa, J. E. and Schuster, R. L. (1988). The formation and failure of
natural dams. Geological Society of America Bulletin, 100,
1054–1068.
Crandell, D. R. and Fahnestock, R. K. (1965). Rockfalls and Avalanches
from Little Tahoma Peak on Mount Rainier, Washington. US
Geological Survey, Bulletin 1221-A.
Crosta, G. B. and Agliardi, F. (2003). Failure forecast for large rock slides
by surface displacement measurements. Canadian Geotechnical
Journal, 40, 176–191.
Cruden, D. M. and Hu, X. Q. (1993). Exhaustion and steady state
models for predicting landslide hazards in the Canadian Rocky
Mountains. Geomorphology, 8, 279–285.
Dai, F. C., Xu, C., Yao, X. et al. (2011). Spatial distribution of land-
slides triggered by the 2008 Ms 8.0 Wenchuan earthquake,
China. Journal of Asian Earth Sciences, 40, 883–895.
Dobrovolny, E. (1962), Geologia del Valle de La Paz. Departamento
Nacional de Geología, Ministerio de Minas y Petróleo, La Paz,
Bolivia.
Dortch, J. M., Owen, L. A., Haneberg, W. C. et al. (2009). Nature and
timing of large landslides in the Himalaya and Transhimalaya of
northern India. Quaternary Science Reviews, 28, 1037–1054.
Eisbacher, G. H. and Clague, J. J. (1984). Destructive Mass Movements
in High Mountains: Hazard and Management. Geological Survey
of Canada, Paper 84–16.
Evans, S. G. and Clague, J. J. (1994). Recent climatic change and cata-
strophic geomorphic processes in mountain environments.
Geomorphology, 10, 107–128.
Evans, S. G., Roberts, N. J., Ischuk, A. et al. (2009a). Landslides trig-
gered by the 1949 Khait earthquake, Tajikistan, and associated
loss of life. Engineering Geology, 109, 195–212.
Evans, S. G., Bishop, N. F., Fidel Smoll, L. et al. (2009b). A re-examina-
tion of the mechanism and human impact of catastrophic mass
ows originating on Nevado Huascarán, Cordillera Blanca, Peru
in 1962 and 1970. Engineering Geology, 108, 96–118.
Evans, S. G., Delaney, K. B., Hermanns, R. L., Strom, A. L. and Scarascia
Mugnozza, G. (2011). The formation and behaviour of nat-
ural and articial rockslide dams: Implications for engineering
Related to folds
Related to faults
0.1
1
10
100
1000
10000
0 200 400 600 800 1000
Relief (m)
Volume (x10
6
m
3
)
Rock avalanche in
glaciated valley section
Rotational slide or topple in
glaciated valley section
Rock avalanche in
glaciated valley sections
Rotational slide in
glaciated valley sections
Rotational slide or topple in
glaciated valley sections
Rock avalanche in
non glaciated valley section
Rotational slide or topple in
non glaciated valley section
Not related to tectonic structures
Rotational slide or topple in
non glaciated valley sections
Fig. 6.7. Plot of landslide volume
versus relief in the transitional area
between the Central and Southern
Andes (modied from Penna et al.,
2011).
Rapid rock-slope failures 69
performance and hazard management. In Natural and Articial
Rock Slide Dams, ed. S. G. Evans, R. L. Hermanns, A. L. Strom
and G. Scarascia Mugnozza. Berlin: Springer, pp. 1–74.
Fauqué, L., Cortés, J. M., Folguera, A. et al. (2008a). Edades de las ava-
lanchas de rocas ubicadas en el Río Mendoza aguas abajo de
Uspallata. Actas del XVII Congreso Geológico Argentino, Jujuy,
1, 282–283.
Fauqué, L., Hermanns, R. L., Wilson, C. et al. (2008b).
Paleorepresamientos del Río Mendoza entre Polvaredas y
Punta de Vacas, Mendoza, Argentina. Actas del XVII Congreso
Geológico Argentino, Jujuy, 1, 274–275.
Fauqué, L., Hermanns, R. L., Hewitt, K. et al. (2009). Mega-
deslizamientos de la pared sur del Cerro Aconcagua y su relación
con depósitos asignados a la glaciación pleistocena. Revista de la
Asociación Geológica Argentina, 65, 691–712.
Ferrer, C. (1999). Represamientos y rupturas de embalses naturales
(lagunas de obstrución) como efectos cosísmicos: Algunos ejem-
plos en los Andes venezolanos. Revista Geográca Venezolana,
40, 109–121.
Fischer, L., Amann, F., Moore, J. and Huggel, C. (2010). Assessment
of periglacial slope stability for the 1988 Tschierva rock ava-
lanche (Piz Morteratsch, Switzerland). Engineering Geology,
116, 32–43.
Furseth, A. (2006). Skredulykker i Norge. Oslo: Tun Forlag.
Glastonbury, J. and Fell, R. (2010). Geotechnical characteristics of large
rapid rock slides. Canadian Geotechnical Journal, 47, 116–132.
Guglielmi, Y. and Cappa, F. (2010). Regional-scale relief evolution
and large landslides: Insights from geomechanical analyses in
the Tinée valley (Southern French Alps). Geomorphology, 117,
121–129.
Hafsten, U. (1986). The establishment of spruce forest in Norway, traced
by pollen analysis and radiocarbon datings. Striae, 24, 101–105.
Harrison, J. V. and Falcon, N. L. (1934). Collapse structures (Kuhgalu
district, Persia). Geological Magazine, 71, 529–539.
Haselton, K., Hilley, G. and Strecker, M. (2002). Average Pleistocene
climatic patterns in the southern Central Andes: Controls on
mountain glaciations and paleoclimate implications. Journal of
Geology, 110, 211–226.
Hauser, A. (2002). Rock avalanche and resulting debris ow in Estero
Parraguirre and Río Colorado, region Metropolitana, Chile.
In Catastrophic Landslides, ed. S. G. Evans and J. V. DeGraff.
Geological Society of America, Reviews in Engineering Geology
15, pp. 135–148.
Heim, A. (1932). Bergsturz und Menschenleben. Zurich: Beiblatt zur
Vierteljahresschrift der Naturforschenden Gesellschaft.
(1949). Observaciones geológicas en la región del terremoto de
Ancash de Noviembre de 1946. Sociedad Geológica del Perú, 25,
2–21.
Hermanns, R. L. and Niedermann, S. (2011). Late Pleistocene–Early
Holocene paleoseismicity deduced from lake sediment deform-
ation and coeval landsliding in the Calchaquíes valleys, NW
Argentina. In Geological Criteria for Evaluating Seismicity
Revisited: Forty Years of Paleoseismic Investigations and the
Natural Record of Past Earthquakes, ed. F. A. Audemard, A.
Michetti and J. P. McCalpin. Geological Society of America,
Special Paper 479, pp. 181–194.
Hermanns, R. L. and Schellenberger, A. (2008). Quaternary tephro-
chronology helps dene conditioning factors and triggering
mechanisms of rock avalanches in NW Argentina. Quaternary
International, 178, 261–275.
Hermanns, R. L. and Strecker, M. R. (1999). Structural and lithological
controls on large Quaternary rock avalanches (sturzstroms) in
arid northwestern Argentina. Geological Society of America
Bulletin, 111, 934–948.
Hermanns, R. L., Trauth, M. H., Niedermann, S., McWilliams, M.
and Strecker, M. R. (2000). Tephrochronologic constraints on
temporal distribution of large landslides in northwest Argentina.
Journal of Geology, 108, 35–52.
Hermanns, R. L., Niedermann, S., Villanueva Garcia, A., Sosa Gomez,
J. and Strecker, M. R. (2001). Neotectonics and catastrophic
failure of mountain fronts in the southern intra-Andean Puna
Plateau, Argentina. Geology, 29, 619–623.
Hermanns, R., Blikra, L., Naumann, M. et al. (2006). Examples of mul-
tiple rock-slope collapses from Köfels (Ötz Valley, Austria) and
western Norway. Engineering Geology, 83, 94–108.
Hermanns, R. L., Blikra, L. H., Anda, E. et al. (2012). Systematic map-
ping of large unstable rock slopes in Norway. In Proceedings of
the 2nd World Landslide Forum, Rome.
Hewitt, K., Gosse, J. and Clague, J. J. (2011). Rock avalanches and the
pace of late Quaternary development of river valleys in the
Karakoram Himalaya. Geological Society of America Bulletin,
123, 1836–1850.
Hsü, K. J. (1975). Catastrophic debris streams (sturzstroms) gener-
ated by rockfalls, Geological Society of America Bulletin, 86,
129–140.
Huggel, C., Zgraggen-Oswald, S., Haeberli, W. et al. (2005). The 2002
rock/ice avalanche at Kolka/Karmadon, Russian Caucasus:
Assessment of extraordinary avalanche formation and mobility,
and application of QuickBird satellite imagery. Natural Hazards
and Earth System Sciences, 5, 173–187.
Huggel, C., Fischer, L., Schneider, D. and Haeberli, W. (2010). Research
advances on climate-induced slope instability in glacier and
permafrost high-mountain environments. Geographica Helvetica,
65, 146–156.
Hungr, O. (2006). Rock avalanche occurrence, process and modelling.
In Landslides from Massive Rock Slope Failure. Proceedings
of the NATO Advanced Research Workshop on Massive Rock
Slope Failure: New Models for Hazard Assessment, Celano, Italy,
16–21 June 2002, ed. S. G. Evans, G. Scarascia Mugnozza, A.
Strom and R. L. Hermanns. NATO Science Series IV, Earth and
Environmental Sciences 49. Dordrecht, Netherlands: Springer,
pp. 243–266.
Hungr, O. and Evans, S. G. (2004). Entrainment of debris in rock ava-
lanches: An analysis of a long run-out mechanism. Geological
Society of America Bulletin, 116, 1240–1252.
Jackson, L. E., Jr. (2002). Landslides and landscape evolution in the
Rocky Mountains and adjacent foothills area, southwestern
Alberta, Canada. In Catastrophic Landslides, ed. S. G. Evans
and J. V. DeGraff. Geological Society of America, Reviews in
Engineering Geology 15, pp. 325–344.
Jibson, R. W. (2009). Using landslides for paleoseismic analysis. In
Paleoseismology, ed. J. P. McCalpin. Burlington, MA: Academic
Press, pp. 565–601.
Jibson, R. W., Harp, E. L., Schulz, W. and Keefer, D. K. (2006). Large
rock avalanches triggered by the M7.9 Denali fault, Alaska,
earthquake of 3 November 2002. Engineering Geology, 83,
144–160.
Kampherm, T. S., Evans, S. G. and Valderrama Murillo, P. (2009).
Landslides triggered by the 1946 Ancash earthquake, Peru.
Geophysical Research Abstracts, 11, EGU2009–13820.
Keefer, D. K. (1984). Landslides caused by earthquakes. Geological
Society of America Bulletin, 95, 406–421.
Kojan, E. and Hutchinson, J. N. (1978). Mayunmarca rockslide and
debris ow, Peru. In Rockslides and Avalanches. 1. Natural
Phenomena, ed. B. Voight. Amsterdam: Elsevier Scientic
Publishing Company, pp. 315–361.
Leroueil, S., Locat, J., Vaunat, J., Picarelli, L. and Faure, R. (1996).
Geotechnical characterization of slope movements. In
Proceedings of the International Symposium on Landslides: 1.
Rotterdam: Balkema, pp. 53–74.
Longva, O., Blikra, L. H. and Dehls, J. F. (2009). Rock avalanches:
Distribution and Frequencies in the Inner Part of Storfjorden,
Hermanns and Longva 70
Møre og Romsdal County, Norway. Norwegian Geotechnical
Institute, Report 2009.002.
Martino, S., Moscatelli, M. and Scarascia Mugnozza, G. (2004).
Quaternary mass movements controlled by a structurally
complex setting in the central Apennines (Italy). Engineering
Geology, 72, 33–55.
Montandon, F. (1933). Chronologie des grands éboulements alpins du
debut de l’ére chrétienne à nos jours. Matériaux pour i’étude des
calamités, 32, 271–340.
Mörner, N.-A. (1996). Liquefaction and varve deformation as evidence
of paleoseismic events and tsunamis: The autumn 10,430 BP
case in Sweden. Quaternary Science Reviews, 15, 939–948.
Nicoletti, P. G. and Sorriso-Valvo, M. (1991). Geomorphic controls of
the shape and mobility of rock avalanches. Geological Society of
America Bulletin, 103, 1365–1373.
Noetzli, J., Gruber, S., Kohl, T., Salzman, N. and Haeberli, W.
(2007). Three-dimensional distribution and evolution of
permafrost temperatures in idealized high-mountain top-
ography. Journal of Geophysical Research, 112, F02S13.
doi:10.1029/2006JF000545.
Owen, L. A., Kamp, U., Khattak, G. A. et al. (2008). Landslides triggered
by the 8 October 2005 Kashmir earthquake. Geomorphology, 94,
1–9.
Penna, I., Hermanns, R. L., Folguera, A. and Niedermann, S. (2011).
Multiple slope failures associated with neotectonic activity in the
southern central Andes (37°–37°30′S). Patagonia, Argentina.
Geological Society of America Bulletin, 123, 1880–1895.
Plafker, G. and Ericksen, G. E. (1978). Nevados Huascaran avalanches,
Peru. In Rockslides and Avalanches, ed. B. Voight. Amsterdam:
Elsevier, pp. 277–314.
Prager, C., Zangerl, C., Patzelt, G. and Brandner, R. (2008). Age dis-
tribution of fossil landslides in the Tyrol (Austria) and its sur-
rounding areas. Natural Hazard and Earth System Sciences, 8,
377–407.
Redeld, T. F. and Osmundsen, P. T. (2009). The Tjellefonna fault
system of Western Norway: Linking late-Caledonian exten-
sion, post-Caledonian normal faulting, and Tertiary rock col-
umn uplift with the landslide-generated tsunami event of 1756.
Tectonophysics, 474, 106–123.
Rodríguez, C. E., Bommer, J. J. and Chandler, R. J. (1999). Earthquake-
induced landslides: 1980–1997. Soil Dynamics and Earthquake
Engineering, 18, 325–346.
Rosas, M., Baumann, V., Videla, A. et al. (2007). Estudio Geocientíco
Aplicado al Ordenamiento Territorial, Puente del Inca, Provincia
de Mendoza. Servicio Geológico Minero Argentino, Informe
Final.
Rosas, M., Wilson, C., Hermanns, R. L., Fauqué, L. and Baumann,
V. (2008). Avalanchas de rocas de las cuevas una evidencia de
la destabilisación de las laderas como consecuencia del cambio
climático del Pleistoceno superior. Actas del XVII Congreso
Geológico Argentino, Jujuy, 1, 313–314.
Sanchez, G., Rolland, Y., Corsini, M. G. et al. (2010). Relationships
between tectonics, slope instability and climate change: Cosmic
ray exposure dating of active faults, landslides and glacial sur-
faces in the SW Alps. Geomorphology, 117, 1–13.
Scheidegger, A. E. (1961). Theoretical Geomorphology. Berlin: Springer.
Schwartz, D. (1988). Paleoseismicity and neotectonics of the Cordillera
Blanca fault zone, northern Peruvian Andes. Journal of
Geophysical Research, 93(B5), 4712–4730.
Schuster, R. L. and Alford, D. (2004). Usoi landslide dam and lake Sarez,
Pamir Mountains, Tajikistan. Environmental and Engineering
Geoscience, 10, 151–168.
Sepulveda, S., Serey, A., Lara, M., Pavez, A. and Rebolledo, S. (2010).
Landslides induced by the April 2007 Aysén fjord earthquake,
Chilean Patagonia. Landslides, 7, 483–492.
Shreve, R. L. (1966). Sherman landslide, Alaska. Science, 154,
1639–1643.
Soldati, M., Corsini, A. and Pasuto, A. (2004). Landslides and climate
change in the Italian Dolomites since the Late Glacial. Catena,
55, 141–161.
Strasser, M., Anselmetti, F. S., Fäh, D., Giardini, D. and Schnellmann,
M. (2006). Magnitudes and source areas of large prehistoric
northern alpine earthquakes revealed by slope failures in lakes.
Geology, 34, 1005–1008.
Strecker, M. R. and Marrett, R. A. (1999). Kinematic evolution of fault
ramps and role in development of landslides and lakes in north-
western Argentine Andes. Geology, 27, 307–310.
Tappin, D. R. (2010). Mass transport events and their tsunami hazard.
In Submarine Mass Movements and Their Consequences, ed.
D. C. Mosher, L. Moscardelli, J. D. Chaytor, C. D. P. Baxter, H. J.
Lee and R. Urgeles. Berlin: Springer, pp. 667–684.
Trauth, M. H., Alonso, R. A., Haselton, K. R., Hermanns, R. L. and
Strecker, M. R. (2000). Climate change and mass movements in
the NW Argentine Andes. Earth and Planetary Science Letters,
179, 243–256.
von Poschinger, A., Wassmer, P. and Maisch, M. (2006). The Flims
rockslide: History of interpretation and new insights. In
Landslides from Massive Rock Slope Failure. Proceedings of the
NATO Advanced Research Workshop on Massive Rock Slope
Failure: New Models for Hazard Assessment, Celano, Italy,
16–21 June 2002, ed. S. G. Evans, G. Scarascia Mugnozza, A.
Strom and R. L. Hermanns. NATO Science Series IV, Earth and
Environmental Sciences 49. Dordrecht, Netherlands: Springer,
pp. 329–356.
Welkner, D., Eberhardt, E. and Hermanns, R. L. (2010). Hazard inves-
tigation of the Portillo rock avalanche site, central Andes, Chile,
using an integrated eld mapping and numerical modelling
approach. Engineering Geology, 114, 278–297.
Wörner, G., Uhlig, D., Kohler, I. and Seyfried, H. (2002). Evolution of
the west Andean escarpment at 18°S (N. Chile) during the last 25
ma: Uplift, erosion and collapse through time. Tectonophysics,
345, 183–198.
Yarnold, J. C. (1993). Rock-avalanche characteristics in dry climates
and the effect of ow into lakes: Insights from mid-Tertiary
sedimentary breccias near Artillery Peak, Arizona, Geological
Society of America Bulletin, 105, 345–360.
