Hummocks: How they form and how they evolve in rockslide-debris avalanches

Abstract

Hummocks are topographic features of large landslides and rockslide-debris avalanches common in volcanic settings. We use scaled analog models to study hummock formation and explore their importance in understanding landslide kinematics and dynamics. The models are designed to replicate large-scale volcanic collapses but are relevant also to non-volcanic settings. We characterize hummocks in terms of their evolution, spatial distribution, and internal structure from slide initiation to final arrest. Hummocks initially form by extensional faulting as a landslide begins to move. During motion, individual large blocks develop and spread, creating an initial distribution, with small hummocks at the landslide front and larger ones at the back. As the mass spreads, hummocks can get wider but may decrease in height, break up, or merge to form bigger and long anticlinal hummocks when confined. Hummock size depends on their position in the initial mass, modified by subsequent breakup or coalescence. A hummock has normal faults that flatten into low-angle detachments and merge with a basal shear zone. In areas of transverse movement within a landslide, elongate hummocks develop between strike–slip flower structures. All the model structures are consistent with field observations and suggest a general brittle-slide emplacement for most landslide avalanches. Absence of hummocks and fault-like features in the deposit may imply a more fluidal flow of emplacement or very low cohesion of lithologies. Hummocks can be used as kinematic indicators to indicate landslide evolution and reconstruct initial failures and provide a framework with which to study emplacement dynamics.

Full text

Landslides
DOI 10.1007/s10346-012-0368-y
Received: 7 August 2012
Accepted: 31 October 2012
© Springer-Verlag Berlin Heidelberg 2012
E. M. R. Paguican IB. van Wyk de Vries IA. Lagmay
Hummocks: how they form and how they evolve
in rockslide-debris avalanches
Abstract Hummocks are topographic features of large landslides
and rockslide-debris avalanches common in volcanic settings. We
use scaled analog models to study hummock formation and ex-
plore their importance in understanding landslide kinematics and
dynamics. The models are designed to replicate large-scale volca-
nic collapses but are relevant also to non-volcanic settings. We
characterize hummocks in terms of their evolution, spatial distri-
bution, and internal structure from slide initiation to final arrest.
Hummocks initially form by extensional faulting as a landslide
begins to move. During motion, individual large blocks develop
and spread, creating an initial distribution, with small hummocks
at the landslide front and larger ones at the back. As the mass
spreads, hummocks can get wider but may decrease in height,
break up, or merge to form bigger and long anticlinal hummocks
when confined. Hummock size depends on their position in the
initial mass, modified by subsequent breakup or coalescence. A
hummock has normal faults that flatten into low-angle detach-
ments and merge with a basal shear zone. In areas of transverse
movement within a landslide, elongate hummocks develop be-
tween strike–slip flower structures. All the model structures are
consistent with field observations and suggest a general brittle-
slide emplacement for most landslide avalanches. Absence of
hummocks and fault-like features in the deposit may imply a more
fluidal flow of emplacement or very low cohesion of lithologies.
Hummocks can be used as kinematic indicators to indicate land-
slide evolution and reconstruct initial failures and provide a
framework with which to study emplacement dynamics.
Keywords Hummocks .Avalanche .Large landslides .Analog
modeling .Horst and graben structures
Introduction
Hummocks are morphological features seen as mounds and ridges
that characterize large landslides and debris avalanches. Hum-
mocks are seen on most sub-aerial and sub-marine mass move-
ments on the Earth and also on other planets. They are especially
common on volcanic mass movements, for example, the Iriga
debris avalanches (Fig. 1) (Voight et al. 1981; Siebert 1984; Glicken
1986). The hummock family includes torevas, which are large tilted
and rotated blocks left within or at the foot of the failure scar.
Torevas can reach up to several kilometers in size and can disag-
gregate on their downhill sides into smaller hummocks (Lucchitta
1979; Francis et al. 1985; Wadge et al. 1995). Often steep-sided in the
downslope direction, the proximal sides of torevas are often filled
in by post-collapse material (Glicken 1991; Palmer et al. 1991).
Downslope of the torevas are smaller hummocks. These can have
radial or transverse orientation with respect to the landslide trans-
port direction. This arrangement has been explained to be due to
basal shear resistance when hummocks are either slowed and
sculpted by adjacent faster moving material parallel to the flow
direction (Glicken 1986,1991), stretched during transport
(Dufresne 2009), and compressed by deceleration when an
avalanche encounters topographic irregularities (Eppler et al.
1987) or water bodies (Siebert et al. 1995). Their shape has also
been attributed to faulting formed as the mass spreads (Shea and
van Wyk de Vries 2008), suggesting a link between hummock
shape and the spreading kinematics.
Hummocks can thus form from horst and graben structures
during lateral spreading of an avalanche (Voight et al. 1981,1983) or
by the separation of individual avalanche blocks rafted in finer-
grained material (Glicken 1986,1996; Crandell 1989). Their height
and number density often decrease away from source (Ui 1983;
Siebert 1984; Glicken 1986,1996; Crandell 1989), a fact that has
been often explained to be due to progressive disaggregation of
debris avalanche blocks (Ui and Glicken 1986; Takarada et al.
1999).
Prominent elongated, sub-parallel alignments hummock
trains have been considered to be remnants of longitudinal ridges
(Dufresne and Davies 2009). Such longitudinal ridges are probably
remnants of hummocks dissected by transport parallel strike–slip
faults related to lateral velocity changes (Shea and van Wyk de
Vries 2008; Andrade and van Wyk de Vries 2010).
Using analog models, we study the evolution and spatial
distribution of hummocks in large-scale volcanic landslides. Inter-
nal and surface structures and morphology of the mass move-
ments are characterized with the aim to understand the formation
and geometry of hummocks and to explore their use as an indi-
cator of landslide kinematics and dynamics.
Methodology
Analog models
Most avalanche analog or numerical model studies have focused
on understanding the transport and emplacement mechanisms of
landslides (Campbell 1989; Campbell et al. 1995; Pouliquen and
Renaut 1996; Staron et al. 2001; Kelfoun and Druitt 2005) and
verifying models of granular flow assumed to operate in such
events (Denlinger and Iverson 2001; Iverson and Denlinger
2001). Recent analog models by Shea and van Wyk de Vries
(2008) and Andrade and van Wyk de Vries (2010) explored the
kinematics of rockslides by describing the deposit structure and
morphology of the upper brittle layer from the early stages of
collapse towards the final phase of material runout. These models
use either a polished surface to simulate a low basal friction
contact or a ductile basal layer to simulate basal ductile deforma-
tion. The main mass of the landslide is modeled by a granular
sand-and-plaster mix.
In our models, we assume, like Andrade and van Wyk de Vries
(2010), a basal ductile layer and a frictional deformation regime of
the main body. In the model, there are basically two deformational
regimes: simple shear is concentrated in the ductile base and the
brittle body stretches or contracts in predominantly pure shear.
The analog models record the continuous movement of a low-
viscosity basal silicone layer spreading over an unconfined
Landslides
Original Paper

transport zone. In some models, a lubricating oil layer that simu-
lates a very-low-viscosity basal slide plane underlies the basal
viscous silicone layer. These models provide a hybrid between
the smooth slide models and the viscous basal models.
Model setup and parameters
We use as a standard model a volcanic cone made of sand and
plaster, silicone, and oil (Fig. 2) as used and scaled, for example, by
Delcamp et al. (2008), Andrade and van Wyk de Vries (2010), and
Mathieu and van Wyk de Vries (2011). Sand and plaster represent
the brittle edifice, silicone is the underlying ductile strata, and oil
(when used) is the highly lubricated sliding base. The sand–plaster
mix has cohesion and internal friction scaled to be similar to most
rocky materials of volcanic slopes. By adding plaster, cohesion is
raised as described by Donnadieu and Merle (1998). The silicone
layer simulates the low-viscosity basal layer in spreading strato-
volcanoes and flank collapses (van Wyk de Vries and Francis 1997;
van Wyk de Vries et al. 2000; Wooller et al. 2004; Cecci et al. 2005;
Andrade and van Wyk de Vries 2010). This layer often consists of
mobilized sediments, volcanoclastic, and altered rocks. The oil
placed under the silicone in some experiments decreases friction
between the sliding basal layer and the plastic sheet, thus speeding
up sliding. In such cases, simple shear is concentrated in the oil
layer and the rest deforms by pure shear stretching. This layer
represents a possibly very-low-resistance layer that could be pres-
ent at an avalanche base (e.g., Thompson et al. 2010; Siebert 1984).
The failure angle of the landslide (i.e., its initial plan shape)
and inclination of transport and depositional areas are also taken
into account in the analog models (see Table 1for the scaling
variables). Failure occurs in the edifice when the basal layer de-
formation induced by the cone load and pressure gradient creates
stresses greater than the strength of the sand or sand and plaster.
The collapsing material then slides and spreads downslope on a
plastic-covered table that can be inclined up to 10°. In each model
Fig. 1 Iriga Volcano and her two DAD (Paguican 2012; Paguican et al. 2011,2012) showing the regional faults. Water bodies: Lakes Baao, Buhi and Bato, hummocks. DAD
structures—faults and ridges
Original Paper
Landslides

experiment, initiation and deposition slopes are fixed at 0°, 3°, 5°,
or 10°.
Three sets of experiments were made. Set 1 used pure sand of
negligible cohesion (~0 Pa) with oil under 1 or 2 cm of silicone (see
Table 2for dimensionless ratios). Set 2 used different cohesion
granular layers with sand and plaster proportions of: 3:1 (250 Pa);
1:1 (500 Pa), and 1:3 (750 Pa) and has oil under 1–2 cm of silicone.
Set 3 has the same ratio of cone material as in set 2 but underlain
by 1 cm of silicone without lubrication. In some of the experi-
ments, model cones are topped with pure plaster to enhance the
visibility of structures formed after collapse.
In total, 45 experiments were completed. There were nine
experiments for set 1, 24 for set 2, and 12 for set 3. Sequential
photographs (plan view) were taken to record the development of
surface structures. Vertical sections were made for the set 1 exper-
iment to view internal deformation within the landslide field.
Repetition of the same experiments with identical initial parame-
ters showed the same geometric, morphologic, and dynamic char-
acteristics, thus demonstrating reproducibility.
Scaling
The scaling procedure used in the analog experiments is the same as
in previous models that simulate flank destabilization and cata-
strophic volcano collapses (Donnadieu and Merle 1998; Lagmay et
al. 2000; Vidal and Merle 2000; Andrade and van Wyk de Vries 2010).
Geometric, kinematic, and dynamic scaling between the analog
experiment and real stratovolcanoes is calculated based on the
parameters listed in Table 1.
Scaling determines the conditions necessary for proportional
correspondence between geometric features and forces acting in
nature and in the laboratory. According to the Buckingham Pi
theorem, ten independent dimensionless variables must be
defined and need to be similar between models and nature.
Five of the variables are geometrically and closely similar
(defined in Table 2). Five kinematic and dynamic variables
are calculated using the gravitational (F
G
), inertial (F
I
), failure
resistance (F
R
), and viscous (F
V
)forcesactingonbothnatural
and analog models, which are defined as:
(1) FG¼ρ"g"h
(2) FI¼ρ"V2, where Vis the process velocity
(3) FV¼μ
t, where tis the process time
(4) FR¼Cþh
R2FG
3$FV
!"
,Cis the edifice cohesion and Ris the
edifice radius, assuming a Navier–Coulomb failure.
The densities of sand layers comprising the model cone layers
were similar and averaged at 1,500 kgm
−3
. Our scaling procedure
considers the difference in the average bulk density of a volcanic
edifice estimated at 2,200 kgm
−3
(Williams et al. 1987), and a basal
pumice-rich ignimbrite or sediment with densities in the range of
1,400 to 2,100 kgm
−3
(Bell 2000).
The heights of the analog cones are (1 cm00.12 km in nature)
12, 11, and 11–15 cm for set 1, 2, and 3 experiments, respectively.
Each one is built with slopes of 25–30°, the angle of repose of
stratocones. The model cones for set 1 and 3 experiments have a
radius of 15 cm and for set 2 was 16 to 17.5 cm. Each model is scaled
with respect to the radius of the Socompa and Mombacho volca-
noes that are both gravitational spreading volcanoes with major
landslides with their basal perimeters defined by breaks in slope
and associated thrust-fold belts (van Wyk de Vries and Francis
1997; van Wyk de Vries et al. 2001).
Fig. 2 Analog model setup and illustration of geometrical parameters in avertical view and bplan view. Constants and scaling parameter values are given
in Tables 1and 2
Landslides

The low-viscosity ductile basal layer of the analog models is
composed of equilateral silicone pieces, 18 cm in length, and has a
varying thickness of about 15 % of the model cone height. The vertex
of the silicone is always placed at the center of the cone to ensure that
edifice collapse includesthe summit. The edges of the silicone define
the lateral limits of the subsequent collapse amphitheater.
Sand and plaster cohesion were scaled to the cohesion of volca-
nic rocks that range from 10
6
to 10
8
Pa for intact basalts and other
lavas, 10
2
to 10
5
Pa for volcanic ash and tephra (Afrouz 1992; Bell
2000), and 10
4
to 10
7
Pa for both lavas and tephra. There is no direct
measurement for the basal layer viscosity during catastrophic col-
lapses, but it may be estimated in the order of 10
7
Pas from numerical
Table 1 List of scaling variables
Definition Variable Unit Nature (N) Model (M) Ratio (M/N)
Edifice height hm 1,300 0.11 8.5×10
−5
Edifice radius Rm 4,000 0.15 3.8×10
−5
Edifice cohesion CPa 10
4
250 0.025×10
−5
Edifice density ρkg m
−3
2,200 1,500 0.7
Basal layer angle αrad pp/6 0.17
Basal layer length Dm 8,000 0.18 2.3×10
−5
Basal layer vertex distance dm? 0 –
Basal layer thickness Tm 200 0.01 5×10
−5
Basal layer viscosity μPa s 10
7
20,000 0.002
Basal layer density γkg m
−3
1,400 1,000 0.7
Failure and deposition slope β°0–45° 0°, 3°, 5° 0–0.1°
Velocity Vms
−1
100 10
−6
10
−7
Time tt 60 >3,600 >60
Gravity acceleration gms
−2
9.8 9.8 1
Table 2 Definition and values of the independent dimensionless variables (after Andrade and van Wyk de Vries (2010))
Dimensionless
variable
Definition Equation Value
Nature Model
(M1)
Model
(M2)
Model (M3) M1
N
M2
N
M3
N
π
1
Edifice (height/
radius)
h
R0.3 0.7 0.7–0.9 0.8 ~1 ~1 ~1
π
2
Basal layer thickness/
edifice height
T
h0.1–0.2 0.1–0.2 0.1 0.1–0.2 ~1 ~1 ~1
π
3
Basal layer length/
edifice
radius
D
R1.3–2 1.2 1–1.1 1.2 ~1 ~1 ~1
π
4
Basal layer vertex
distance/edifice
radius
d
R?000 –––
π
5
a, angular distance aR
h1.6–10.8 0.7 0.6–0.8 0.7 ~1 ~1 ~1
π
6
Edifice/basal layer
density
ρ
g
1–1.6 1.5 1.5 1.5 ~1 ~1 ~1
π
7
Gravitational/viscous
forces
FG
FV
168–375 291 291–397 317 ~1 ~1 ~1
π
8
Frictional/viscous
forces
FR
FV
32–36 186–276 178–361 168 ~1 ~1 ~1
π
9
Inertial/viscous
forces
FI
FV
1.3–132 3×10
−4
–3×
10
−5
3×10
−4
–3×
10
−5
3×10
−4
–3×10
−5
Low Low Low
π
10
Inertial/gravity
forces
FI
FG
0.8–
0.00352
9×10
−7
–9×
10
−8
9×10
−7
–7×
10
−8
8×10
−7
–8×
10
−8
Low Low Low
Original Paper
Landslides

analysis and simulations of debris avalanche flows (Sousa and Voight
1995; Dade and Huppert 1998; Kelfoun and Druitt 2005; Andrade and
van Wyk de Vries 2010). The analog silicone has a viscosity of ~10
4
Pa
s and the oil ~10
−1
Pas.
Results
Standard experiment
The experiments generally result in shorter runout than in the
natural counterpart as at a late stage deformation becomes very
slow; however, the morphology and structures observed are sim-
ilar to those observed in large-scale landslide deposits and are seen
to remain similar in those experiments that attained scaled natural
runouts (Fig. 1; Online Resource 1). Despite changing variables and
configurations in the sets of experiments, there are recurrent
morphological features and structures that are developed across
the experiments. The structures produced are thus general fea-
tures, and their appearance does not depend strongly on thickness
of silicone or cohesion of the brittle layer.
Deformation of the analog cones starts as soon as they are
constructed and ends when deformation is almost negligible and
overall morphology remains invariable. The rates of deformation
depend mainly on the silicon viscosity, the presence of an oil layer,
and to a lesser extent on the sand–plaster mixture cohesion.
Experiments that have long runout are stopped when the analog
landslide reaches the end of the table setup.
Model avalanche class
Model avalanches can be of three classes and can develop three
types of hummock morphology and structure (Fig. 3; Online Re-
source 1): class A avalanches have progressive spreading and
extension, class B avalanches undergo progressive spreading but
have a localized early compression phase related to the spreading
proximal zone pushing against a decelerating distal area, and class
C avalanches have late-stage compression in the deposition zone.
Individual avalanches can change from one type to another during
their development.
Surface morphology and structures
There are two major zones identified in our experiments
(Figs. 3,4,and5): the collapse zone, which is equivalent to
the toreva domain of Andrade and van Wyk de Vries (2010),
and the depositional zone, formed when the avalanche pro-
gresses out of the amphitheater limits, laterally and longitu-
dinally spreading out material at its base. These two zones are
generally separated by a major graben that Andrade and van
Wyk de Vri es (2010) also observed. This graben is an arcuate
depression perpendicular to the slide direction. It separates
the area of larger toreva-like hummocks from the zone with
smaller hummocks (Figs. 3,4,and5).
The collapse zone is further sub-divided into the upper,
the middle, and the lower collapse areas (Fig. 3). The upper
collapse area is composed primarily of summit material dom-
inated by normal faults oriented perpendicular to the slide
direction; the middle collapse area is where numerous strike–
slip faults appear; and the lower collapse zone is where
arcuate normal faults, with convex traces towards the volcanic
cone, are dominant. Irregularities of elevation profile in the
collapse zone are due to toreva tilting and rotation.
The depositional zone is sub-divided into the proximal,
medial, and distal areas. It is where torevas and blocks break
up as the landslide spreads laterally and longitudinally. Prom-
inent features in this zone are lateral levees, ridges, and
hummocks (Fig. 3). The depositional zone, like the collapse
zone, is cut by normal, strike–slip, and thrust faults. Normal
faults appear in extension-dominated areas, where they ac-
commodate basal shearing, tilting, and rotation of the sliding
materials. They also appear in the medial to distal deposition-
al zones as the deposits spread further. Thrust faults dominate
in the proximal and distal accumulation zones where the
materials approach gentler slopes or at the frontal edge where
margins resist spreading. Arcuate transtensional faults form in
the sliding direction because of the differential forward move-
ment of the deposit combined with lateral spreading.
Hummocks and torevas are observed in all experiments (Fig. 3).
Large hummock trains in set 2 and 3 experiments are parallel to the
slide direction. In set 1, they are transverse and curved towards the
frontal margins. For sets 2 and 3 with 1:1 ratio of sand and plaster,
both large and small hummocks form and are superimposed on the
same area. Large hummocks are elongated, whereas the smaller
hummocks are equant. Only small and rounded hummocks are
observed in experiments with the less cohesive, pure sand set 1
experiments.
Plan view shape
Avalanche deposits produced in set 1 and 2 experiments with oil
have long runout equaling that of natural deposits, whereas set 3
without the oil have a short runout and with a wide depositional
zone (Figs. 3and 8). The presence of the oil layer and low cohesion
of the cone edifice in the experiments influences the landslide
runout shape and length (exp. 1.6, 2.3, and 2.7 in Fig. 3). On the
other hand, the absence of oil and higher cohesion dampens
spreading of the avalanche front. There is instead enhanced lateral
spreading that generates a wide depositional zone (exp. 3.6 and 3.8
in Fig. 3). Lobes can form at the frontal and lateral ends of each
landslide deposit, and the lobe shape can vary (Fig. 3). Experi-
ments with oil produce landslide deposits with irregular lobes.
Those without lubrication have symmetrical, more rounded, and
more regular lobes. This is attributed to the presence of oil layer as
the reduced friction by lubrication of the silicone affects the basal-
and upper-layer spreading rate on both lateral and frontal
margins.
Subsurface deformation
Vertical cross-sections (Fig. 5a), parallel to the main sliding direc-
tion, reveal the avalanche internal structure. The collapse and
depositional zones are dissected by normal faults and are separat-
ed by a graben. Strike–slip faults, often trending arcuate towards
the frontal margin, are difficult to see in this view but are clear in
the plan views in Fig. 3. In the collapse zone, listric normal faults
converge into the ductile shear zone. The listric normal faults
accommodate the sliding, tilting, and rotation of the blocks in
the collapse zone.
In the depositional zone, the listric faults are more shallowly
dipping and more closely spaced. In the topmost part of the brittle
distal layer, there are also shallower and minor low-angle normal
faults that accommodate the sliding, tilting, and rotation of
smaller blocks within the brittle material. These produce the
Landslides

Fig. 3 Development of analog avalanches. Five analog models that represent the three sets of experiments are grouped according to avalanche classes and the resulting
hummock types: one for set 1 and two each for sets 2 and 3. Avalanches can be of class A with progressive spreading and extension resulting in progressively broken,
primary, smaller hummocks, Hummock type 1a; avalanche class B for progressive spreading with internal compression in some areas resulting in a wider range of
hummock size. Hummock type 1b and 2 or avalanche class C for progressive sliding with late-stage compression resulting in increasing hummock size. Hummock type 1b.
Gray arrows indicate that during an avalanche, hummocks can start big and break during extension or they can merge sometime during their development. For each
experiment, three photos representing the avalanche evolution and hummock development are presented in the first row, with their hummocks, debris field, and
depositional zone delineations below for statistical analysis of hummock area and spatial distribution. Morphological features are labeled and delineated in some
experiments: collapse zone (CZ) and its upper (UC), medial (MC) and lower (LC) areas; depositional zone (DZ) and its proximal (PD), medial (MD), and distal (DD)
areas; the graben (G); accumulation zone (AZ) in the PD, MD, or DD, ridges (R), torevas (T), and hummocks (H); and the surface structures: normal (red), thrust (blue),
and strike–slip (yellow) faults
Fig. 4 A 3D visualization example of structural and morphological development of an avalanche and formation of hummocks during its evolution (the 3D models are
created by the ARC 3D webservice, developed by the VISICS research group of the K.U. Leuven in Belgium). Time A, formation of normal faults and movement along these
cause edifice extension downwards by sliding; time B, collapse of more edifice materials near the summit and formation of more normal faults in the collapse (CZ) and
depositional (DZ) zones. Movement along these faults spreads the materials further down, with the graben forming as the edifice materials come out of the amphitheater
limits towards a gentler slope. At time C, torevas (T) formed since the earlier stages are more evident and the materials at the base spread both to the lateral and
downslope directions. Hummocks in the DZ are well formed at this stage, and at time D, the avalanche spreads and hummocks break up
Original Paper
Landslides

secondary small hummocks that appear like detachment blocks
gliding on top of the deposit or sometimes stranded on top of the
larger hummocks.
The graben is a transition point of the normal faults between
the collapse and depositional zones. It is where sliding torevas
start to break up as the landslide spreads into a wider depositional
Fig. 5 Sub-surface (a) vertical view parallel to slide direction of an analog avalanche and (b) interpretation. aPlan view image for the locations of the cross-sections 1–6.
The different sub-areas for collapse and depositional zones and graben are shown in 1–6. The analog has an upper brittle layer where normal faults are evident and a
ductile layer underneath with varying thickness throughout the avalanche area. High-angle normal faults (HANF) in the upper brittle layer become listric and converge
into a low-angle normal fault (LANF). These faults accommodate the sliding, tilting, and rotation of edifice blocks in the collapse zone, forming the torevas and their
smaller counterpart, first-order type 1 and second-order type 2 hummocks. On the topmost part of the brittle layer are shallower low-angle normal faults that
accommodate the sliding, tilting, and rotation of minor blocks at the very top, forming the type 2 hummocks. bA cross-section interpretation of an avalanche
Landslides

zone or where torevas accumulate as they approach a gentler
slope.
In general, the dip of normal faults decreases with depth
and distance from the summit. This is shown by listric faults
that converge in the boundary of the upper brittle layer and
the underlying ductile layer in the collapse zone. Many shal-
low normal faults do not reach this boundary in the deposi-
tional zone but flatten into the lower part of the brittle
material. The high- and low-angle normal faults in the col-
lapse zone create the torevas. In the depositional zone, they
create hummocks. Figure 5b shows a cross-sectional interpre-
tation and morphology of these structures.
From the collapse zone towards the proximal deposition
zone, brittle shearing dominates as ductile shearing only
occurs in the upper part of the thinning silicone. From the
medial towards the distal deposition zone, however, ductile
shearing dominates, resulting in both the thrusting of the
silicone and thinning or thickening of the sand–plaster layer.
Hummocks
In this study, each experiment is divided into three stages: initial,
development, and final deposition. Experiments that best repre-
sent the recurrent morphology and structures are chosen and
presented in Fig. 3and Supplement A. Using statistical analysis,
hummock area, spatial distribution, and factors affecting it are
explored for patterns that aid to investigate their types and evolu-
tion throughout the landslide–avalanche development.
Characterization
We distinguish two types of hummocks formed in the experiments
(Figs. 3and 5): type 1—primary hummocks and type 2—secondary
hummocks. Type 1 hummocks are of two classes: type 1a are the
small and generally equant hummocks formed early on through
the initial development stages of faulting and may undergo minor
breakup during emplacement and type 1b are formed at the same
time but are the larger and often elongated hummocks. Type 2
forms on top and alongside the larger type 1b hummocks. Type 1b
hummocks develop from original landslide blocks that became
stretched and flattened, and they may also grow from the accretion
of other type 1 hummocks.
Exploratory statistics
Hummock area is automatically calculated from registered exper-
iment photographs managed in a GIS. Hummock number density
is generated using the center of a hummock and number per unit
area. A line graph of increasing hummock area (Fig. 6) shows that
experimental avalanches can produce either a highly contrasting
or a gradually changing population of hummock type. Small hum-
mocks, types 1a and 2 (exp. 1.8, 2.7, 2.3, 3.8) resulting from class A
and B avalanches (extension during avalanche), tend to have a
very high frequency of small and very few big hummocks.
This is seen by the very abrupt change in slope on Fig. 6.
Type 1b hummocks produced by class B and C analog ava-
lanches (with compression at some points during the ava-
lanche) (exp. 2.3, 3.8, 3.6) exhibit gently sloping graphs,
indicating that within certain stages during the avalanche the
hummocks are of restricted range of sizes. In general, hum-
mock population may decrease as their size increases or
populations increase as they progressively break up.
During an avalanche, types 1a and 2 small hummocks
break up and disintegrate if they are formed by lower-cohe-
sion material as shown by an overall decrease in hummock
size (Fig. 7a; Online Resource 2, left). However, as an ava-
lanche slows or when cone materials block the spreading of
silicone at the front of the avalanche causing compression,
hummocks can merge and grow (Fig. 7b; Online Resource 2,
right). This results in larger hummocks at the medial to distal
margins.
Spatial distribution
Hummock spatial distribution is recorded using the center of each
hummock as described in Yoshida et al. (2012). The elongation and
direction of the ellipse (Fig. 8) represents the general direction and
orientation in which all the hummocks form and separate. In
general, this ellipse follows the plan view shape of the debris
avalanche deposit itself.
The mean center (Fig. 8) shows the point where the hummock
population center moves at each stage of the experiment. In
general, these trends allow visualization of the link between central
hummocks, the overall direction of movement, and the avalanche
flow direction. The trend in the directional ellipse and mean center
of the hummock areas show that spreading occurs mostly parallel
to the long axis of the avalanche for low-cohesion material and
with a lubricated base. In contrast, spreading is dominantly lateral
for high basal friction and more cohesive experiments.
Sequence of events
The development of a rockslide debris avalanche model can be
sub-divided into three stages depending on the structures and
morphological features that form. These are the slide initiation,
development, and final emplacement. The slide initiation stage is
also described by Andrade and van Wyk de Vries (2010), but they
did not carry the models further. A discussion regarding these
stages is presented in this section with accompanying diagrams
(Figs. 4and 9) that show the development of the overall morpho-
logical and structural features and the formation and evolution of
hummocks.
Slide initiation
Fractures begin to develop at the onset of edifice collapse. As soon
as the underlying ductile layer in the experiments begins to
stretch, sliding initiates and transforms early fractures into faults.
These faults define the amphitheater walls (Figs. 4a and 9) and also
accommodate the extension of edifice blocks as the landslide–
avalanche spreads outwards. These blocks may become torevas if
they remain intact in the collapse zone or the proximal area of the
depositional zone. If broken into smaller blocks, they become type
1 hummocks in the depositional zone. Transtensional faults in the
collapse zone start to appear early on.
Development stage
The collapse, depositional zones, and horst and graben structures
become obvious during the development stage (Figs. 4b, c and 9).
More normal faults form and edifice materials slide in both zones.
Large blocks and torevas at the lower part spread more, break up,
and form the smaller hummocks (Type 1a, 2). Some torevas remain
intact and some can accumulate and combine with neighboring
ones, forming the larger type 1b hummocks.
Original Paper
Landslides

Fig. 6 Increasing line plot of hummock area. Avalanches can produce either a combination of a large number of small hummocks or a very low count of big hummocks.
Small hummocks, regardless of whether they are primary or secondary, can be highly contrasting, having mostly very small and very few large hummocks. Sizes of type 2
hummocks, however, exhibit a gradual change in slope, implying that for avalanches that produce larger hummocks (compression-dominated avalanches), their
hummocks are of similar sizes during transport
Fig. 7 Line graph showing the mean areas of hummocks with respect to the distance from the summit (the greater the time, the farther the hummocks are from the
source). Type 1 (a) and type 2 (b) hummocks will always tend to break up due to spreading and extension during an avalanche emplacement. However, they can start to
merge and form together with an increase in size once they undergo compression
Landslides

Final stage
During the final stage, type 2 detachment hummocks become
evident, whereas torevas in the collapse zone remain unchanged.
In the depositional zone, hummocks can break up or remain intact
as the avalanche continues to spread. The ductile layer may build
up at the margin where a thrusted brittle upper layer forms
compression ridges.
Discussion
To understand what hummocks are and how they are formed,
their relationship with other structural and morphological features
within an avalanche field must first be recognized. From the
analog models, we observe that hummocks form in all experi-
ments but that their final spatial distribution, size, elongation,
and morphology can vary (Fig. 2).
Avalanche characteristics
The analog avalanche models constructed in this study develop a
similar and consistent set of structures. In all of the experiments,
the collapse and depositional zones are separated by a graben and
faults dissect the entire mass (Figs. 3,4, and 5). In the collapse
zone, large listric faults form the main failure planes of the ava-
lanche. From the middle towards the base of the collapse zone,
transtensional and compressional faults appear as the collapsing
mass spreads outwards (Figs. 3,4, and 5). The areas between the
faults are the hummocks, effectively horsts.
The depositional field has three areas: the proximal,
medial, and distal areas. Size, orientation, elongation, and
spatial distribution of the hummocks are different in each
of these areas and are controlled by varying emplacement
conditions.
Fig. 8 Trend of the directional ellipses and mean centers. The directional ellipses show the lateral or longitudinal formation direction of hummocks and the mean center is
the center point of the central hummock
Original Paper
Landslides

In experiments where a low-cohesion pure sand cone
stands on an inclined plane and the base is oil-lubricated,
progressive but gradual spreading and thinning of the ava-
lanche is observed (Fig. 3). On the other hand, in more
cohesive models of sand and plaster cones on a gently in-
clined plane, regardless of lubrication at the base (Fig. 3),
higher friction and resistance to slide result in an accumula-
tion area within the depositional zone. Hummocks are
formed from the broken edifice blocks that slid, tilted, and
rotated as the avalanche progresses.
Hummock description
Morphology
Hummocks can be characterized according to their basal shape,
surface morphology, and orientation with respect to flow and slide
direction. The basal shape of hummocks can be equant, circular,
rectangular, or elliptical. Hummock deformation and shape after
initial formation depends on the velocity distribution and the
strain field of the avalanche. Their initial shapes stretch along
acceleration and contract on the deceleration directions. Hum-
mock elongation is a product of initial shape and subsequent
deformation that can be either dominantly parallel, perpendicular,
or randomly oriented with respect to flow direction. The surface
morphology of hummocks can be flat-topped or pointed, ridged,
or rounded. Densely faulted hummocks can create pinnacle-
topped features (Fig. 9b).
Generally, circular-based and flat-topped hummocks are
formed in rapidly emplaced long-runout models, whereas elongat-
ed and pinnacle hummocks forminexperimentsofcohesive
avalanche material with moderate runout. The high slope angle
of collapse base, lower cohesion, and greater lubrication can in-
crease the runout length and decrease the lateral or longitudinal
spreading direction, shape, and orientation of individual
hummocks.
Types
We identify different hummock types based on their size and
development (Fig. 3): type 1a are the first-order small hummocks;
type 1b are the first-order large, elongated hummocks; and type 2
are the second-order small, polygonal hummocks that formed
after and raft over the type 1b hummocks. The extensional faulting
of avalanche deposits forms the first-order hummocks. Small-scale
shear faulting or brittle fracturing separating and detaching the
stretched upper layers of the larger primary hummocks forms the
type 2 hummocks. They are a boudinage feature of stretched and
separated competent layers, such as those seen in many avalanche
hummock cuts (e.g., Fig. 5f in Shea et al. (2008) or Fig. 4 in Bernard
et al. (2009)) or in other avalanche analog models (Shea and van
Wyk de Vries 2008).
From the vertical cross-sections, we observe that high-angle
normal faults in the upper brittle layer converge into low-angle
normal faults towards the lower part, accommodating the sliding,
tilting, and rotation of edifice blocks forming the type 1 hum-
mocks. The shallower low-angle normal faults or brittle fractures
on the topmost part of the brittle layer serve as a décollement
interface of the different lithologies forming the type 2 hummocks.
Trends in size and spatial distribution
Hummocks can have a low population density with a large size
distribution, as is generally the case when type 1b hummocks are
present, or there can be a high population but with a small size
distribution, as in the case for deposits with dominantly small
hummocks (types 1a and 2). If a model avalanche produces large
hummocks, the difference in size distribution is minimal. On the
other hand, if smaller hummocks form, their size range is broad,
with most hummocks at the lower end of the size spectrum.
As an avalanche develops, hummocks can decrease in number
but increase in size or they can increase in number but decrease in
size. At the final stage, however, hummock population is more
often higher compared to the early stages, indicating that hum-
mocks do break up during sliding development.
A progressively spreading and thinning avalanche body (class
A) is dominated by extension. The largest hummocks are confined
in the lower collapse zone and proximal depositional zone, and
type 1a hummocks spread all over the depositional zone. This is
observed in Iriga DAD2 (debris avalanche deposit) (Paguican 2012;
Paguican et al. 2011,2012). In avalanche classes B and C where
there is substantial compression due to frontal restriction, there
are hummock types 1b and 2. The bigger hummocks can be located
Fig. 9 The different stages of avalanche emplacement showing an interpretation of hummock formation (a) and degree of faulting (b)
Landslides

anywhere in the depositional zone and can form the accumulation
zones. This is clearly observed at the Tetivicha avalanche, Bolivia
(Shea and van Wyk de Vries 2008) and Iriga debris avalanche
deposits (Paguican 2012; Paguican et al. 2011,2012).
Avalanche stages and hummock formation
From the analog models, it is clear that hummock formation is
intimately linked to the development of the avalanche. At the first
stage, normal listric faults accommodate the extension by sliding
and rotation of blocks. These blocks are the initial hummocks.
Then, during transport, the avalanche may spread freely or may
encounter barriers and topographic constraints such as other
topography. During this stage, the hummock population develops
and evolves.
From landslide to hummocks
The low-angle normal faults at the base and the high-angle normal
faults at the upper layer create the sliding plane. This plane
facilitates most of the sliding, tilting, and rotation of edifice blocks
as the body spreads down the amphitheater. The major normal
faults limit the boundaries of the collapse amphitheater. Near the
amphitheater base and outside of its limits, edifice materials will
start to spread laterally as well as longitudinally. As the sliding
plane activates and initiates sliding, brecciation and faulting in the
edifice occurs and blocks are formed by the main normal faults. As
the collapsing edifice mass moves down, extension of the upper
layers above the slide plane induces normal faulting, generating
horsts and grabens. The faulted horst blocks become the building
blocks of hummocks as the whole mass spreads.
As the edifice materials come out of the amphitheater and
into the depositional zone, debris will spread and extend not only
downslope but also laterally. As the materials approach the gentler
slope at the piedmont, proximal materials accumulate, creating
ridges oriented perpendicular to the flow. As more materials slide
down from the edifice, these materials are pushed to spread
further, where extensional or compressional structures may dom-
inate depending on the topography and morphology of the depo-
sitional zone. Naturally, extension happens on unconfined
depositional environments with regular topography. On the other
hand, compressional features such as ridges parallel, sub-par-
allel, or perpendicular to the flow direction may develop in
confined environments. These ridges may form distal accu-
mulation zones if frontal debris is compressed in the distal
front. Distal accumulation zones may be remobilized into
lahars if saturated or could remobilize as secondary slides,
as at Socompa (Kelfoun et al. 2008; Glicken 1986). High-angle
normal faults and tension fractures are not clearly rooted to
the basal layer in the depositional zone; their strain is taken
up by diffuse deformation in the lower brittle zone or by
structures too small to be observed at the model scale.
Spreading and increasing extension by block sliding, tilt-
ing, and rotation is the stage when hummocks are developed.
As spreading proceeds, hummocks can break and move far-
ther apart, resulting in increased type 1a and 2 hummocks.
Conversely, confinement can move hummocks closer and
from larger type 1b structures. Figure 4shows the juxtaposi-
tion of type 1b as well as the breakup of the type 1a and 2
hummocks. Such a situation was described at Parinacota
(Clavero et al. 2002).
The role of normal faults
The existence of listric normal faults at the collapse zone is a
consequence of the geometry of the sliding mass and the brittle
and ductile layers present. The main failure plane is the contact
where the base of the brittle blocks slides on a non-slip or fused
interface with the ductile basal layer or the stable edifice. Relative
movements of the blocks are accommodated by pervasive defor-
mation in the ductile layer under this interface. High-angle normal
faults in the collapse zone are generally rooted in the basal layer,
although this is not the case for those in the shallower depositional
zone that do not clearly extend down to the silicone, rather they
pass into a lower brittle horizontal stretching area.
Layering within the avalanche, with a sliding or slip layer at
the top such as the upper ignimbrite layer or the lower block and
ash layer at Iriga DAD2 (Paguican 2012; Paguican et al. 2011,2012),
generates an internal décollement causing minor low-angle nor-
mal faulting over the discontinuity. If there is a pre-existing dis-
continuity between layers allowing the upper layer to glide over
the lower layer, this décollement may take up the deformation.
Rotational movement of the upper layer could also generate slid-
ing at the interface. These internal avalanche structures exist as a
result of layer interfaces that could be caused by a major internal
structure or by a contrasting stratigraphy.
Degree of stretching
During the development of the landslide and evolution of hum-
mocks, the scale variation of stretched portions is observed. Post-
collapse lengths are mostly twice its pre-collapse basal radii (see
Online Resource 1). The width, thickness, and volume of the active
area change during stretching. The normal listric faults in the
collapse zone decreases in dip as they approach the depositional
zone. In the medial to distal areas, they become more horizontal
by instantaneous stretching (Fig. 9a).
Based on the increasing degree of stretching (Fig. 9a) and
normal faulting (Fig. 9b), the degree of stretching (Fig. 9) changes
from stage 1 to 3 (from modified main structural occurrences of
stretching in the crust by Brun and Choukroune (1983)). Stage 1 is
when the faults begin to form discrete blocks, with more intense
deformation between blocks. At stage 2, changes in the thickness
of the upper brittle and basal ductile layers begin and are more
evident and widely observable in stage 3. Stage 3 is when hum-
mocks are fully formed and tend to spread instead of breaking
apart with deformation affecting the surrounding zones. The basal
ductile layer becomes thin towards the summit area, and most of it
is extruded and spread under the avalanche. The ductile layer may
come to the surface due to the high degrees of stretching (the same
thing as observed at Socompa (Shea et al. 2008)). Structures shown
in the cross-sections (Fig. 5) are those of late stage 4 when the
avalanche is nearly fully stretched.
Conclusions
The analog models presented here allow the visualization of the
development of landslide–avalanche hummocks and provide a
structural and kinematic sequence of their formation. This can then
be used to understand natural avalanche hummock structure and
distribution and thus obtain the emplacement history. Hummocks
are a useful tool for understanding avalanche emplacement.
Hummocks are the morphological expression of brittle layer
deformation due to spreading in landslides and avalanches. They are
Original Paper
Landslides

principally the stretched remains of tilted and rotated blocks
of the original failure volume. In general, both landslide–
avalanches and hummocks initiate and evolve in a similar
way. Landslides and avalanches start with fracturing and
faulting of the failure area (or, for a volcano, the edifice).
As strain increases, the upper landslide layer can split apart,
forming grabens and horsts and creating steeply dipping
escarpments parallel to the slide direction. Normal and
strike–slip faults accommodate the sliding, tilting, and rota-
tion of blocks that later evolve into hummocks. These faults
also accommodate the extension and spreading within hum-
mocks. Therefore, hummocks are formed by extension of the
avalanche, with the hummocks spreading and moving apart as
the avalanche spreads. They can also move together or com-
press, or new ridge-like hummocks can form due to avalanche
constriction by topography.
The morphology and spatial distribution of hummocks
are dependent on the interplay of the number density of
normal, thrust, and strike–slip faults, which is a function of
strain and material properties such as cohesion and the vis-
cosity of a ductile sliding layer. Hummock shape and size
depends primarily on the original position in the initial land-
slide and the first-formed structures. The hummock shape
changes with subsequent spreading, breakup, or merger. Dur-
ing landsliding, hummocks can break up, resulting in a higher
number density; on the other hand, hummock merger during
compression can reduce number density. At the last stages of
emplacement, hummocks can either break or remain intact
while spreading farther apart from each other and can spread
themselves, increasing their surface area.
The structures within avalanche deposits and hummocks
indicate that they form by extension and are consistent with a
general slide model where pure shear stretching dominates in
the upper brittle part and simple shear is concentrated at the
base. When hummocks and fault-like features are not formed,
a more fluidal flow type of emplacement could be possible.
Lastly, we call primary hummocks in large landslides and
avalanches type 1, which can be divided into two main
groups: the type 1a small hummocks, which are generally
found toward the front, and the type 1b large hummocks that
are initially formed from the initial landslide toreva blocks
and are found in the proximal and medial zones. Type 1b
hummocks can also form by coalescence during compressive
or arrest stages.
The type 2 hummocks are secondary features formed by a
boudinage-like process during layer parallel stretching. They
can exist on or around the type 1b hummocks and are
generally seen as single blocks of more consolidated material.
Avalanches that spread freely (class A) have large numb-
ers of small type 1a and few large type 1b hummocks. Ava-
lanches that become constrained (class B and C) tend to
preserve the larger type 1b and tend to have accompanying
type 2 secondary ones.
Acknowledgments
This study is a result of the cotutelle Ph.D. program between the
Université Blaise Pascal and the University of the Philippines.
Funding was provided by the French Embassy in Manila and the
EIFFEL Excellence Scholarship.
References
Afrouz AA (1992) Practical handbook of rock mass classification systems and modes of
ground failure. CRC, London
Andrade D, van Wyk de Vries B (2010) Structural analysis of the early stages of
catastrophic stratovolcano flank-collapse using analogue models. Bull Volcanol
72:771–789. doi:10.1007/s00445-010-0363-x
Bell FG (2000) Engineering properties of soils and rocks, 4th edn. Blackwell, Oxford
Bernard B, van Wyk de Vries B, Leyrit H (2009) Distinguishing volcanic debris avalanche
deposits from their reworked products: the Perrier sequence (French Massif Central).
Bull Volcanol 71:1041–1056. doi:10.1007/s00445-009-0285-7
Brun J-P, Choukroune P (1983) Normal faulting, block tilting, and décollement in a
stretched crust. Tectonics 2(4):345–356
Campbell CS (1989) Self-lubrication for long runout landslides. J Geolog 97:653–665
Campbell CS, Cleary PW, Hopkins MJ (1995) Large-scale landslide simulations: global
deformation, velocities and basal friction. J Geophys Res 100:8267–8283.
doi:10.1029/94JB00937
Cecci E, van Wyk de Vries B, Lavest J-M (2005) Flank spreading and collapse of weak-
cored volcanoes. Bull Volcanol 67:72–91. doi:10.1007/s00445-004-0369-3
Clavero JE, Sparks RSJ, Huppert HE (2002) Geological constraints on the emplacement
mechanism of the Parinacota avalanche, northern Chile. Bull Volcanol 64:40–54
Crandell DR (1989) Gigantic debris avalanche of Pleistocene age from ancestral Mount
Shasta Volcano, California, and debris-avalanche hazard zonation. US Geolog Survey
Bull 1861:31
Dade WB, Huppert HE (1998) Long run-out rockfalls. Geology 26:803–806
Delcamp A, van Wyk de Vries B, James MR (2008) The influence of edifice slope and
substrata on volcano spreading. J Volcanol Geotherm Res 177(4):925-943. http://
dx.doi.org/10.1016/j.jvolgeores.2008.07.014
Denlinger RP, Iverson RM (2001) Flow of variably fluidized granular masses across three-
dimensional terrain. 2. Numerical predictions and experimental tests. J Geophys Res
106(B1):553–566. doi:10.1029/2000JB900330
Donnadieu F, Merle O (1998) Experiments on the indentation process during
cryptodome intrusions: new insights into Mount St. Helens deformation.
Geology 26:79–82
Dufresne A (2009) The influence of runout path material on rock and debris avalanches:
field evidence and analogue modeling. Dissertation, Canterbury University,
Canterbury
Dufresne A, Davies TR (2009) Longitudinal ridges in mass movement deposits. Geomor-
phology 105:171–181
Eppler DB, Fink J, Fletcher R (1987) Rheologic properties and kinematics of emplacement
of the Chaos Jumbles rockfall avalanche, Lassen Volcanic National Park, California. J
Geophys Res 92:3623–3633
Francis PW, Gardeweg M, Ramirez CF, Rothery DA (1985) Catastrophic debris avalanche
deposit of Socompa volcano, northern Chile. Geology 13:600–603
Glicken H (1986) Rockslide debris-avalanche of May 18, 1980, Mount St. Helens volcano.
Dissertation, University of California, Santa Barbara
Glicken H (1991) Sedimentary architecture of large volcanic-debris avalanches. In: Fisher
RV, Smith GA (eds) Sedimentation in volcanic settings. SEPM Special Publication 45:
pp 99-106
Glicken H (1996) Rockslide-debris avalanche of May 18, 1980, Mount St. Helens Volcano,
Washington. US Geol Surv Open-File Rep 96–0677:1–90
Iverson RM, Denlinger RP (2001) Flow of variably fluidized granular masses across three-
dimensional terrain. 1. Coulomb mixture theory. J Geophys Res 106(B1):537–552.
doi:10.1029/2000JB003758
Kelfoun K, Druitt T (2005) Numerical modeling of the emplacement of Socompa rock
avalanche, Chile. J Geophys Res 110:B12202. doi:10.1029/2005JB003758
Kelfoun K, Druitt T, van Wyk de Vries B, Guilbaud M-N (2008) Topographic reflection of
the Socompa debris avalanche, Chile. Bull Volcanol 70:1169–1187. doi:10.1007/
s00445-008-0201-6
Lagmay AMF, van Wyk de Vries B, Kerle N, Pyle DM (2000) Volcano instability induced by
strike–slip faulting. Bull Volcanol 62:331–346. doi:10.1007/s004450000103
Lucchitta BK (1979) Landslides in Valles Marineris. J Geophys Res 84:8097–8113
Mathieu L, van Wyk de Vries B (2011) The impact of strike–slip, transtensional and
transpressional fault zones on volcanoes. Part I: scaled experiments. J Structural Geol.
doi:10.1016/j.jsg2011.03.002
Paguican EMR (2012) The structure, morphology, and surface texture of debris avalanche
deposits: field and remote sensing mapping and analogue modelling. Dissertation,
Université Blaise Pascal, Clermont-Ferrand
Paguican EMR, van Wyk de Vries B, Lagmay AMF (2011) Hummocks in large avalanches:
how they form and what they mean. Geophys Res Abstract. Vol. 13, EGU2011-3437
Landslides

Paguican EMR, van Wyk de Vries B, Lagmay AMF (2012) Volcano-tectonic controls and
emplacement kinematics of the Iriga debris avalanches (Philippines). Bull Volcanol 74
(9):2067–2081. doi:10.1007/s00445-012-0652-7
Palmer AB, Alloway BV, Neall VE (1991) Volcanic debris avalanche deposits in New Zealand
—lithofacies organization in unconfined wet-avalanche flows. In: Fisher RV, Smith GA
(eds) Sedimentation in volcanic settings. SPEM Special Publication 45, pp 89–98
Pouliquen O, Renaut N (1996) Onset of granular flows on an inclined rough surface:
dilutancy effects. J Physique II 6:929–935. doi:10.1051/jp2:1996220
Shea T, van Wyk de Vries B (2008) Structural analysis and analogue modeling of the
kinematics and dynamics of rockslide avalanches. Geosphere 4(4):657–686.
doi:10.1130/GES00131.1
Shea T, van Wyk de Vries B, Pilato (2008) Emplacement mechanism of contrasting debris
avalanches at Volcan Mombacho (Nicaragua), provided by structural and facies
analysis. Bull Volcanol 70:899–921. doi:10.1007/s00445-007-0177-7
Siebert L (1984) Large volcanic debris avalanches: characteristics of source areas,
deposits, and associated eruptions. J Volcanol Geotherm Res 22:163–197
Siebert L, Beget JE, Glicken H (1995) The 1883 and late prehistoric eruptions of
Augustine volcano, Alaska. In: Ida Y, Voight B (eds) Models of magmatic processes
and volcanic eruptions. J Volcanol Geotherm Res 66, pp 367-395
Sousa J, Voight B (1995) Multiple-pulsed debris avalanche emplacement at Mount St.
Helens in 1980: evidence from numerical continuum flow simulations. J Volcanol
Geotherm Res 66:227–250
Staron L, Vilotte JP, Radjai F (2001) Friction and mobilization of contacts in granular
numerical avalanches. In: Kishino Y (ed) Powders and grains. Swets and Zeitlinger,
Lisse, pp 451–454
Takarada S, Ui T, Yamamoto Y (1999) Depositional features and transportation mecha-
nism of valley-filling Iwasegawa and Kaida debris avalanches, Japan. Bull Volcanol
60:508–522
Thompson N, Bennet MR, Petford N (2010) Development of characteristic volcanic debris
avalanche deposit structures: new insights from distinct element simulations. J
Volcanol Geotherm Res 192:191–200. doi:10.1016/j.volgeores.2010.02.021
Ui T (1983) Volcanic dry avalanche deposits—identification and comparison with non-
volcanic debris stream deposits. J Volcanol Geotherm Res 18:135–150
Ui T, Glicken H (1986) Internal structural variations in a debris-avalanche deposit from
ancestral Mount Shasta, California, USA. Bull Volcanol 48:189–194
van Wyk de Vries B, Francis PW (1997) Catastrophic collapse at stratovolcanoes induced
by gradual volcano spreading. Nature 387:387–390
van Wyk de Vries B, Kerle N, Petley D (2000) Sector collapse forming at Casita volcano,
Nicaragua. Geology 28:167–170
van Wyk de Vries B, Self S, Francis PW, Keszthelyi L (2001) A gravitational spreading
origin for the Socompa debris avalanche. J Volcanol Geotherm Res 105:225–227
Vidal N, Merle O (2000) Reactivation of basement faults beneath volcanoes: a new model
of flank collapse. J Volcanol Geotherm Res 99:9–26
Voight B, Glicken H, Janda RJ, Douglas PM (1981) Catastrophic rockslide avalanche of
May 18. In: Lipman PW, Mullineaux DR (eds) The 1980 eruptions of Mount St. Helens,
Washington. US Geol Surv Prof Pap 1250:347-377
Voight B, Glicken H, Janda RJ, Douglass PM (1983) Nature and mechanics of the Mount
St Helens rockslide-avalanche of 18 May 1980. Geotechnique 33:243–273
Wadge G, Francis PW, Ramirez CF (1995) The Socompa collapse and avalanche event. In:
Ida Y, Voight B (eds) Models of magmatic processes and volcanic eruptions. J Volcanol
Geotherm Res 66, pp 309-336
Williams DL, Abrams G, Finn C, Dzurisin D, Johnson DJ, Denlinger R (1987) Evidence from
gravity data for an intrusive complex beneath Mount St. Helens. J Geophys Res
92:10207–10222
Wooller L, van Wyk de Vries B, Murray JB, Rymer H, Meyer S (2004) Volcano
spreading controlled by dipping substrata. Geology 32(7):573–576. doi:/
10.1130/G20472.1
Yoshida H, Sugai T, Ohmori H (2012) Size–distance relationships for hummocks on
volcanic rockslide-debris avalanche deposits in Japan. Geomorphology 136(1):76–87.
doi:10.1016/j.geomorph.2011.04.044, ISSN 0169-555X
Electronic supplementary material The online version of this article (doi:10.1007/
s10346-012-0368-y) contains supplementary material, which is available to authorized
users.
E. M. R. Paguican ()):B. van Wyk de Vries
Laboratoire Magmas et Volcans, Université Blaise Pascal,
Clermont Université,
5 rue Kessler, 63038, BP 10448, 63000 Clermont-Ferrand, France
e-mail: engiellepaguican@gmail.com
E. M. R. Paguican :B. van Wyk de Vries
UMR 6524, LMV,
CNRS,
63038 Clermont-Ferrand, France
E. M. R. Paguican :B. van Wyk de Vries
R 163, LMV,
IRD,
63038 Clermont-Ferrand, France
E. M. R. Paguican :B. van Wyk de Vries :A. M. F. Lagmay
National Institute of Geological Sciences, College of Science,
University of the Philippines,
Diliman, Quezon 1101, Philippines
Original Paper
Landslides