The 1978 quick clay landslide at Rissa, mid
Norway: subaqueous morphology and tsunami
simulations
J.-S. L'Heureux1*, R. S. Eilertsen2, S. Glimsdal3, D. Issler3 I.-L. Solberg1, C. B.
Harbitz3
1. Geological Survey of Norway (NGU) & International Centre for Geohazards (ICG),
Trondheim, Norway. *e-mail: jean.lheureux@ngu.no
2. Geological Survey of Norway (NGU), Tromsø, Norway
3. Norwegian Geotechnical Institute (NGI) & ICG, Oslo, Norway
Abstract The 1978 landslide at Rissa is the largest to have struck Norway during
the last century and is world-famous because it was filmed. Swath bathymetry data and seismic reflection profiles reveal detailed information about the subaqueous
morphology of the mass-transport deposits (MTD). Results show that the landslide
affected nearly 20 % of the lake floor and that it exhibits a complex morphology
including distinct lobes, transversal ridges, longitudinal ridges, flow structures and
rafted blocks. The rafted blocks found at the outer-rim of the MTD travelled a distance of over 1000 m in the early stage of the landslide on an almost flat basin
floor. Simulation of sediment dynamics and tsunami modelling show that the rafted blocks most likely triggered the flood wave with a recorded maximum surface
elevation of 6.8 m.
Keywords Rissa landslide • quick clay • coastal setting • mass transport deposits •
mobility • numerical simulations • tsunami modelling
1. Introduction
Landslides are common in the sensitive marine clay deposits of Canada and
Scandinavia. When occurring along fjords, lakes or rivers these landslides can
generate destructive tsunamis. Although relatively rare, such events can cause
much damage and loss of life. The increasing societal awareness for such natural
hazards calls for a better understanding of the link between such landslides and
their potential to generate tsunamis.
The famous 1978 quick-clay landslide at Rissa is the largest to have struck
Norway in the last century. Of the 40 people caught in the landslide, one person
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died. The water waves that were generated by the landslide as it entered Lake
Botnen, caused flooding and much damage in the village at the other side of the
lake. Run-up heights of up to 6.8 m were documented along the lakeshore (Fig. 1).
In this paper we present an integrated and multi-disciplinary study focusing on
the subaqueous morphology, mobility and tsunamigenic implication of the 1978
Rissa landslide. The aim is to better understand the failure and depositional processes of the Rissa landslide, and to discuss its implication on the generation of
the tsunami waves. Additionally, the available detailed information on the Rissa
landslide and tsunami (i.e. eye-witnesses and amateur videos) is used to constrain
our tsunami simulations.
2. Setting and review of the 1978 landslide at Rissa
Lake Botnen at Rissa is a 1 km narrow and 5 km long inlet connected to the
Trondheimsfjord by a shallow river. Following the last glacial period, the area was
subject to intense glacio-isostatic rebound and a relative fall of sea level. The marine limit lies at 160 m above the present sea level (Reite 1987). The lowlands
along the shores of Lake Botnen are relatively flat and are almost entirely covered
by thick glacio-marine and marine deposits, locally overlain by littoral deposits
(Reite 1987). During their emergence in the Holocene, marine deposits were exposed to fresh groundwater flow and leaching of salts from the pore water which
resulted in the development of very sensitive clay, i.e. quick clays (Rosenquist
1953).
The 1978 landslide took place at the south-western corner of Lake Botnen (Fig.
1B). Gregersen (1981) described the landslide as a two-stage process. At first, an
initial slide was triggered by excavation and stockpiling along the lakeshore. During this initial failure, 70–90 m of the shoreline slid out into the lake, including
half of the recently placed earth-fill. The slide edges were 5–6 m high and extended 15–25 m inland. The landslide developed retrogressively in the south-western
direction over the next 40 minutes. The sediments liquefied completely during the
sliding and the debris literally poured into the lake. At this stage the landslide area
took the shape of a long and narrow pit open towards the lake (Figs. 1–2). The initial landslide evacuation area covered a surface of 25–30,000 m2 (i.e. 6–8 % of the
final slide area) (Gregersen 1981).
The main landslide started almost immediately after retrogressive sliding had
reached the boundaries of stage 1 (Fig. 1B). At this point, large flakes of dry crust
(150 m  200 m) started moving towards the lake, not through the existing gate
opening, but in the direction of the terrain slope onshore (see arrows; Fig. 1B).
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Fig. 1 A) Location of Rissa in mid Norway and B) shaded relief image of Lake Botnen. Registered run-up heights from the generated tsunami following the 1978 landslide are shown around
Lake Botnen in yellow (data from Larsen 1978). Also shown are the location of the initial slide
and the flakes A and B that caused the tsunami. The black arrows indicate direction of movement. C) Picture of the Rissa landslide in 1978 (Photo: Aftenposten 1978).
The velocity was initially moderate (flake A; Fig. 1B), of the order of 10–20
km/h, and increased to 30–40 km/h (flake B; Fig. 1B). Houses and farms can be
seen floating on the sliding masses on the amateur videos. A series of smaller and
retrogressive slides followed over a short period of time. The sliding process
propagated to the mountain side where it stopped. The main sliding stage lasted
for approximately 5 minutes and covered 92–94% of the total slide area (0.33
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2
km ). The total volume of mobilized sediment has been estimated in the range of
5-6 x106 m3.
3. Data and methods
Bathymetry data were acquired in 2010 using a 250 kHz interferometric sonar
system (GeoAcoustics) mounted onboard of a 7 m long craft. Sound velocity in
the water adjacent to the transmitters/receivers was measured with a Valport 650
Sound Velocity Profiler (SVP). These measurements were performed systematically to obtain optimal velocity profiles for calculation of water depths. Positioning was performed by means of differential GPS with an accuracy of ± 1 m. The
bathymetric data are presented with a grid spacing of 1 m. During this survey, a
grid of seismic data was also collected using a 3.5 kHz parametric sub-bottom
profiler. Two-way travel time was converted to water depth and sediment thickness using a constant sound velocity of 1470 ms−1.
4. Subaqueous morphology
Bathymetry data from Lake Botnen reveal a gently dipping basin with steep
shoreline slopes (up to 35°). The basin reaches a maximum water depth of 38 m in
the central part of the lake. The 1978 mass-transport deposits (MTD) are clearly
identified in the southern part of the basin and described below.
4.1. Initial failure area
Outside the area of the initial landslide, the bathymetry data reveals a 135 m
long and 4 m high shore-parallel escarpment (Fig. 2). The escarpment extends for
more than 100 m to the north in the basin and is draped by a thin veneer of sediments with scattered hummocks (up to 10 m wide). The source of this drape is ascribed to the sediment from the SW-NE oriented, narrow, landslide pit that developed in the first stage of the landslide (Fig. 1).
A morphological high is present to the west of the initial slide (Fig. 2). This
high, interpreted as bedrock from seismic data, constrained the displacements in
the initial stage and forced the main landslide event to open a new gate to the west
(see main gate; Fig. 2). The bathymetry data also reveal a ridge and trough morphology between the morphological high and the initial slide escarpment (Fig. 2).
These morphological features show a staircase pattern and could be indicative of
lateral spreading off the lakeshore in the early stage of the landslide.
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Pockmarks are identified at several locations on the lake floor (Figs. 1–3).
These crater–like features, which are up to 60 m wide and 4 m deep, occur at the
foot of the shoreline slopes and could testify to high pore-pressure gradients in the
near-shore sediments. It is therefore likely that naturally driven excess pore pressures contributed to the onset of failure in 1978.
Fig. 2 Shaded relief image with morphological interpretation of the 1978 landslide deposits.
4.2. Main slide body
Deposits from the main landslide cover a large portion of the basin floor (up to
20 % of the lake). The total area covered by the landslide deposits is 0.76 km 2 and
the maximum length of 1.2 km. Seismic data show an average thickness of 6–8 m
and thus a total volume of 4.4–5.8  106 m3 (Fig. 3). Two types of MTDs are identified on the shaded relief image (Fig. 2). The first type consists of well-
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constrained sediment lobes of fine grained sediments and containing randomly deposited small blocks/hummocks (less than 20 m in diameter). The lobes show
many flow structures, whereas the blocks have a semicircular to triangular footprint. The distal portion of the lobes is characterized by arcuate pressure ridges up
to 100–200 m long (Fig. 2). Such ridges are usually compressional and are mainly
associated with flow deceleration combined with continued sediment flow piling
up (Posamentier and Kolla 2003). A 400 m long and up to 1 m high longitudinal
ridge also occurs on the basin floor in the vicinity of the lobes (Fig. 2).
Fig. 3 Seismic reflection profile NGU-1008123 showing the thickness of the 1978 landslide
deposits. Note also the pockmarks in front of the landslide deposits.
Part of the sediment lobes can be traced back to the open gate west of the morphological high (Fig. 2). The character of the lobes suggests that they were deposited in the later stage of the landslide when the clay and debris completely liquefied. The longitudinal ridge may result from simultaneous flows of different
velocity, combined with eddies generated to the north of the morphological high
since large swirls were witnessed in the lake during the event. Additionally, some
of the hummocks on the basin floor could represent houses which sank into the
lake several days after the landslide event.
The second type of MTDs is found in the frontal part of the landslide deposits.
Here, much larger blocks are found (Figs. 2–3). The blocks show rectangular and
elongated forms and are typically 80–200 m long and 30–100 m wide. Some of
these blocks have travelled up to 1150 m from the shoreline. The blocks rise 1–2
m above the lake floor and are oriented transverse to the flow direction.
The grouping of the blocks at the outer rim of the deposit suggests that these
features are related. Rafted blocks are frequently observed in translational landslides (Mulder and Cochonat 1996) and often go together with debris flows (Ilstad
et al. 2004). As such the larger blocks appear to be remnants of the large flakes
sediment that rafted towards the lake in the early stage of landsliding. Seismic data
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also shows that some of the blocks were covered by thin sediment plumes in the
final stages of the landslide.
5. Post-failure analysis of the landslide
In order to create the input for comprehensive tsunami simulations, a dynamical back-calculation of the landslide evolution from the point where it enters the
lake until it stops is needed. The eyewitness reports suggest that the main waves
were generated in the second stage of the landslide when the large flakes of dry
crust entered the lake. The flakes were of similar size, but the first was only half as
fast as the second due to obstructions in the path (i.e. during the opening of the
main gate). Here we concentrate on the second flake (i.e. flake B; Fig. 1B). For
this exploratory study, we opted for the numerical model BING (Imran et al.,
2001), which simulates the flow of a Herschel–Bulkley fluid with a plug layer riding on top of the shear layer using 1D depth-averaged equations. The rheological
exponent n was set to 0.5 to reflect the shear-thinning property of clay. Ideally, the
yield strength τy and viscosity coefficient ν would be determined from laboratory
tests of the slide material. However, remoulding and incorporation of ambient water or very soft seafloor sediments may drastically change the material properties
in the shear layer (De Blasio et al., 2004). The values of τy and ν were adjusted to
reproduce the observed run-out distance. In all simulations, values of τy between
0.1 and 0.3 kPa (depending on the value chosen for ν) were needed to attain the
observed run-out. These values are well inside the range expected for the remoulded shear strength of quick clay (i.e. <0.5 kPa).
Fig. 4 Frontal flow velocity and final sediment deposit for block B using the BING model.
See Fig. 1B for the location of block B.
The pronounced curvature at several points along the slope profile shown in
Fig. 4 required substantial artificial damping. Even then, the run-out distance
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J.S. L’Heureux et al.
proved to be highly sensitive to the input parameter values (e.g. reducing τy by as
little as 1% can make the run-out distance increase by more than 20%). The deposit exhibits a thick head about as long as the block at release (Fig. 4). This is
similar to the morphological observations presented above. In all simulations, the
flow velocity increases sharply to 15–18 m/s on the 10–15° steep shoreline slope
(Fig. 4).
6. Tsunami modelling
In the numerical modelling, the slide (or block) is described as a flexible box.
The box slide is slightly rounded to avoid numerical instability. The length and
width of the block are described from eye-witness observations and they are set to
200 m and 150 m, respectively. Simulations are made for a 7 m high block, using
an analytical slide velocity progression formula as well as the velocity extracted
from the BING simulations. For the analytical approach the initial and maximum
velocity is set to the one observed on the video (i.e. 11 m/s) and decays smoothly
following a sinusoidal velocity profile until the run-out distances are reached (~
1000 m). The velocity profile from the BING simulations follows much the same
pattern, but has a maximum underwater velocity close to 18 m/s (Fig. 4). The approach for the analytical slide progression velocity follows closely the description
by Løvholt et al. (2005). The generation phase is modelled by a tsunami model
called GloBouss (Pedersen and Løvholt 2008). For the run-up heights and maximum water levels, the model called MOST is applied (Titov and Gonzales 1997).
Results from the tsunami simulations are presented in Fig. 5 for both landslide
velocity models. Compared to the observed run-up heights (Fig. 1), the results using the velocity model from BING shows a significant overestimation of the maximum water level in the southernmost part of Botnen (Fig. 5A). Better matching is
found in the northern part due to the extensive wave breaking and wave height reduction in these shallower waters. Tsunami modelling gives a better match when
the analytical slide velocity progression is used (Fig. 5B). Along the western
shoreline of Botnen the run-up heights are overestimated by 0.5–1.5 m. Simulated
heights are close to the observed ones in the northern part (at Leira) and along the
eastern margin of the basin (± 10–30 %). The southernmost part of the lake is not
evaluated here, since the run-up model cannot be initiated directly by a slide motion but must read input from the tsunami model along a boundary outside the
slide area. In Fig. 5C we compare the surface elevation at a synthetic gauge at the
northern side of the lake. As described above, the two different scenarios give
similar results here, including wave breaking close to the shoreline, and the duration of the first (main) wave peak is 2–3 min. This is comparable to the real wave
which was described as one large breaking wave.
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Fig. 5 Maximum surface elevation and run-up heights (m) from tsunami simulations when
using velocity profiles from A) the Bing model and B) the analytical model (i.e. from video
observations). C) Time series from a synthetic gauge (marked with a red bullet in panels
(A) and (B)) located in the northern part of Botnen near Leira and showing the surface elevation for both slide velocity progressions as explained in the text.
7. Conclusions
The available detailed information on the 1978 landslide at Rissa in mid Norway
offers a unique possibility to study the landslide development, the post-failure dynamics and the generation of a tsunami. In this study, the integrated data set indicates that the water waves in Lake Botnen were generated by movement of large
flakes towards the lake. Shaded relief imagery and numerical simulations show
that parts of these flakes, or blocks, remained relatively intact and travelled for
more than 1000 m on the basin floor at velocities up to 18 m/s. The transportation
of large flakes of dry crust is frequently observed in quick clay landslides (TerStepanian 2000). Therefore, in coastal areas where quick clays are found, it is important to assess the possibility for such flakes to develop prior to any tsunamigenic hazard assessments. Finally, comparison of modelled and registered tsunami
run-up heights shows the importance of correctly estimating the landslide velocity
profile when assessing the tsunami-generating capacity of coastal landslides.
Acknowledgments
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We thank O. Totland and J.A. Dahl for their help during the acquisition of the geophysical data
and are grateful to G. Corner and M. Vanneste for their constructive reviews. This is contribution
no. 357 of the International Centre for Geohazards.
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