Geomorphology 132 (2011) 327–338
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Geomorphology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g e o m o r p h
Effects of rock avalanches on glacier behaviour and moraine formation
Natalya V. Reznichenko a,⁎, Tim R.H. Davies a, David J. Alexander b
a
b
Department of Geological Sciences, University of Canterbury, Private Bag 4800, Christchurch, New Zealand
School of Geography, Planning and Environmental Management, University of Queensland, QLD 4072, Brisbane, Australia
a r t i c l e
i n f o
Article history:
Received 10 January 2011
Received in revised form 24 May 2011
Accepted 25 May 2011
Available online 1 June 2011
Keywords:
Rock avalanches
Supraglacial debris
Ice-surface ablation
Glacier advances
Terminal moraines
Paleoclimatology and hazards
a b s t r a c t
Although large rock avalanches are infrequent, sediment production in active orogens is dominated by such
events, which may strongly influence geomorphic processes. Where rock avalanches fall onto glaciers, they
may affect glacier behaviour and moraine formation. We outline the processes of rock avalanche initiation and
motion, and show that supraglacial deposits of rock avalanche debris have distinct sedimentological and
thermal properties. Laboratory experiments on the effects of such debris on ice ablation are supplemented by
field data from two rock avalanches in the Southern Alps, New Zealand. Their effects are compared with those
of the thinner supraglacial debris that results from small rockfalls and melt-out of englacial debris.
Implications of rock–avalanche debris cover for glacier behaviour are explored using a mass-balance model of
the Franz Josef Glacier in New Zealand, demonstrating a likely supraglacial rock avalanche origin for the
Waiho Loop moraine, and considering the potential hazard of a large rock avalanche onto the present-day
glacier.
© 2011 Elsevier B.V. All rights reserved.
1. Introduction
In this paper we examine the effects of the deposits of large rock
avalanches on glacier behaviour and moraine formation. Supraglacial
rock avalanche deposits on the glacier ablation zone reduce icesurface ablation, and this can significantly alter the overall glacier
mass balance. With a sufficient proportion of the ablation zone
covered by rock avalanche debris the glacier mass balance will
become significantly less negative, and, as a result, the terminus may
advance. This advance may generate a terminal moraine that has no
climate-related cause. The much-debated Late-Glacial Waiho Loop
moraine, the Franz Josef Glacier, New Zealand (Barrows et al., 2007;
Tovar et al., 2008), appears to be the result of such a sequence of
events and we use it as an example to demonstrate how a large rock
avalanche deposit can reduce ablation causing a glacier to advance
and form a prominent frontal moraine.
Rock avalanche deposits on glaciers and their effects have received
particular attention only recently. The increasing frequency of
recognised historical and prehistorical catastrophic rock avalanche
deposits in high mountains areas worldwide (e.g., Post, 1967;
Whitehouse and Griffiths, 1983; Hewitt, 1998, 2009; Hermanns
et al., 2001; Geertsema et al., 2006; Allen et al., 2011), together with
well-defined magnitude–frequency relationships that indicate that
very large events contribute more volume over time than do mediumsized ones (Malamud et al., 2004; Korup and Clague, 2009) has
⁎ Corresponding author. Tel.: + 64 3 364 2987x7697; fax: + 64 3 364 2769.
E-mail address: [email protected] (N.V. Reznichenko).
0169-555X/$ – see front matter © 2011 Elsevier B.V. All rights reserved.
doi:10.1016/j.geomorph.2011.05.019
established the importance of such events in the sediment delivery
and geomorphology of active orogens. Where glaciers are present in
mountain valleys, some rock avalanches inevitably fall onto them;
however, supraglacial rock avalanche deposits appear to be less
frequently reported than those that do not fall onto glaciers. This is
partly because of the limited extent of present-day glaciers and also,
significantly, because supraglacial rock avalanche deposits are rapidly
reworked by glaciers and deposited as moraines whose potential
rock–avalanche origins have rarely been considered in the past. By
contrast, a rock avalanche depositing into a glacier-free river valley,
though progressively removed by river erosion, often leaves longterm recognisable traces of its initial extent in the form of remnant
large angular boulders and vestigial high terraces (Chevalier et al.,
2009; Lee et al., 2009). Thus, more likely that the frequency of
supraglacial rock avalanches and, hence, their significance to glacier
dynamics and behaviour have in the past been underestimated.
Notably, many recently identified rock avalanche deposits, for
example those studied by Hewitt (2009) in the Karakoram Himalaya,
had earlier been interpreted as moraines and corresponding (incorrect) paleoclimatic inferences drawn (see also McColl and Davies,
2011; Kariya et al., 2011). One of the conclusions of the present work
is that prominent moraines can be generated by rock–avalanchedriven glacier advances; these cannot presently be distinguished from
moraines reflecting climatically driven advances, and their inclusion
in moraine age data may also have led to incorrect paleoclimate
interpretations.
In this work we highlight recently-published information about
the properties of rock avalanche deposit, their effects on glacier
ablation, and glacier response to reduced ablation. We report an
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investigation of the current state and effects of two recent rock
avalanche deposits on glaciers in Southern Alps of New Zealand, and
develop a conceptual picture of the response of mountain valley
glaciers to emplacement of extensive rock avalanche debris on the
ablation zone. Mass-balance modelling of the Franz Josef Glacier
response to severely reduced ablation also demonstrates the potential
of a large rock avalanche to substantially advance the present-day
terminus of this glacier, with following significant disruption to
society and tourism.
2. Rock avalanches onto glaciers
2.1. Causes
A rock avalanche is a large-scale rock slope failure (of the order of a
million cubic metres or more) that leads to a mass of fragmenting rock
travelling across a landscape at velocities up to 100 m/s (Voight and
Pariseau, 1978; Fort et al., 2009; Hewitt, 2009). Failures of this scale
are frequently triggered by major earthquakes, but a seismic trigger is
not necessary; for example, the 1991 Aoraki/Mt. Cook, 1991 Mt.
Fletcher, 1999 Mt. Adams, and 2007 Young River landslides in the
Southern Alps of New Zealand between 1991 and 2007 were all about
10 7 m 3 in volume and neither seismically- nor rainfall-triggered
(McSaveney, 2002; Hancox et al., 2005; Allen et al., 2011). Relatively
few large-scale rock-slope failures appear to be triggered by rainfall
(Melosh, 1987; Soldati et al., 2004). Contrary to convention, recent
work suggests that loss of ice buttressing of valley slopes during and
following deglaciation may have been overemphasised as a cause of
the apparently high frequency of large rock avalanches in the early
Holocene (McColl et al., 2010).
2.2. Processes
A rock avalanche usually initiates as the detachment of a large
mass of rock from the upper part of a mountain edifice by completion
of a through-going failure surface. To generate the highly-fragmented
debris described below, the volume needs to be in the order of 10 6 m 3
or more (Davies and McSaveney, 2009). The detached mass initially
slides under gravity down the mountainside, progressively collapsing
as it does so into joint-controlled blocks. However, at an early stage,
progressively smaller intact (i.e. uncracked) rocks will start to
comminute (or fragment; e.g., Davies and McSaveney, 2009) under
the high stresses of the fall, generating fragments of all sizes to
submicron diameter; large quantities of fine particles are expelled
from the surface to form a growing aerial dust cloud. By the time it
reaches the glacier surface, the mass is pervasively fragmenting and
travelling at high velocity. The debris then moves rapidly across the
ice surface, spreading both laterally and longitudinally as it does so
and becoming thinner in the process. A rock avalanche travelling
across a glacier continues to comminute until it halts, as demonstrated
by the frequent presence of distally located, shattered but undisaggregated clasts; these are the result of fragmentation and, if formed
more proximal, would have been disaggregated by the shear strain
field in the avalanche (e.g., McSaveney, 1978). Often evidence is found
that the ice-surface has been eroded by the rock avalanche passage,
which contributes mass to the emplaced deposit (e.g., McSaveney,
1978, 2002).
2.3. Deposits
Supraglacial rock avalanche deposits are usually much thinner
relative to their volume than their nonsupraglacial counterparts, and
thus larger in area. This is the result of the lower basal friction of rock
avalanches that travel over snow and ice (Post, 1967; McSaveney,
1975, 2002; Gordon et al., 1978; Evans and Clague, 1988; Jibson et al.,
2004, 2006; Lipovsky et al., 2008; Deline, 2009; Hewitt, 2009), which
allows a given volume to spread over a greater area. For example, the
average thickness of the 40 × 10 6 m 3 Sherman Glacier rock avalanche
of 1964 (Fig. 1) was about 1.5 m (McSaveney, 1978), whereas the
thickness of the majority of reported nonsupraglacial rock avalanches
thicknesses varies between 10 m and 60 m or more (Smith et al.,
2006). This characteristic allows a rock avalanche to cover a very large
area of glacier with a debris carpet some metres thick.
As a result of the intense fragmentation that occurs during fall and
runout, the debris of rock avalanches is predominantly extremely fine.
For example, McSaveney and Davies (2007) found that 90% by weight
(N99% by number) of the surface debris of the 1991 Aoraki/Mt. Cook
rock avalanche, New Zealand, was b10 − 5 m in diameter. The coarser
(metre-scale) “carapace” layer is usually at the debris surface (caused
by less intense fragmentation at the surface because of lower
confining pressure); so the surface appearance is often blocky, but
most of the underlying debris is extremely fine, with larger clasts of all
sizes forming a fractal grain size distribution (Fig. 2). The bulk density
of the subsurface debris is high caused by the wide grain size
distribution, and the voids ratio and permeability are correspondingly
low, compared to those of other types of supraglacial debris. As a
Fig. 1. The Sherman Glacier, Alaska, as seen on 27 August 1963 (left) and with rock avalanche deposit on 24 August 1964 (Post, 1965).
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Fig. 2. Examples of rock avalanche deposits: (A) basal exposure of the nonsupraglacial 1929 Falling Mountain Rock avalanche, NZ; and (B) rock avalanche debris about 2 m thick on
the terminal surface of the Mueller Glacier, Southern Alps, NZ.
result, rock avalanche debris has high thermal mass and inertia that is,
it requires a large quantity of heat to increase its temperature, so it
heats up slowly compared to debris with large volumes of air-filled
voids. This gives rock avalanche debris high insulation capacity,
compared to the thin melt-out debris common on valley glaciers, as
we now describe.
3. Supraglacial debris: Effects on ablation
The major direct effect of a rock avalanche deposit on a glacier is to
protect the ice surface beneath the debris from solar radiation,
resulting in reduced surface ablation beneath the deposit (Østrem,
1965; McSaveney, 1975; Lundstrom et al., 1993; Mattson et al., 1993;
Nakawo and Rana, 1999; Brock et al., 2010; Reid and Brock, 2010;
Reznichenko et al., 2010). The heat flux through the debris layer
depends on its thickness and physical properties (lithology, bulk
density or porosity, albedo, moisture content, and permeability; e.g.,
Nakawo and Young, 1981, 1982; Kayastha et al., 2000; Nicholson and
Benn, 2006), and determines the ice-surface melt rate. Where a large
rock avalanche deposit covers a sufficient proportion of the ablation
zone with debris of the order of metres thick, surface-ice ablation will
almost completely cease and, as a result, the dynamics of the covered
glacier will be significantly affected (Shulmeister et al., 2009).
Ablation within or beneath the ice will be unaffected by surficial
debris, but this is usually a small component of total ablation (of the
order of 10%; Alexander et al., in review).
3.1. Melt-out debris
The ablation zones of many mountain valley glaciers are covered
with extensive supraglacial debris. This results mostly from the melt-out
of englacially or subglacially transported debris originating from
subglacial or extraglacial sources (Rogerson et al., 1986; Winkler,
2009); it is commonly several to a few tens of centimeters in thickness
and characterized by angular, coarse clasts with a small proportion of
fines and, as a result, high permeability (Drewry, 1986), low bulk
density and relatively low thermal inertia. Melt-out debris increases in
thickness toward the terminus from lateral displacement and continued
emergence of englacial debris and if it achieves sufficient thickness can
reverse the increasing rate of clear-ice ablation toward the terminus
caused by reduction in glacier surface elevation. The high debris porosity
favours air and water circulation, reducing the insulation effect of the
debris; for example, Sharp (1949) found that coarse debris 15 cm deep
has the same insulating effect on ice as sand 3 cm deep. Debris-covered
glacier snouts usually are very rough, so a several-centimetres-thick
debris layer forms a mosaic with thin layers of debris, clear patches of
the ice, and exposed ice cliffs. Differential melting amplifies the
topographic unevenness of the glacier and complicates the estimation
of total melting rates under this cover (Fig. 3).
3.2. Rock avalanche debris
The effect of rock avalanche debris on underlying glacier ice is
distinctly different from that of melt-out debris owing to its greater
Fig. 3. The debris-covered lower Tasman Glacier, Southern Alps, NZ. (A) The rough topography of the debris-covered glacier about 8 km from the modern terminus with welldeveloped glacier ‘karst’; the Hochstetter icefall is the steep tributary glacier in the background. (B) Differential ablation from discontinuous debris cover where the Hochstetter ice
stream flows into the Tasman Glacier. Note the general thinness of the debris cover.
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thickness (meters) and thermal inertia (because of the fractal grain
size distribution and high proportion of very fine grains). The effect of
high thermal mass is to require a large quantity of heat to be input
before the mass can achieve steady-state heat transfer, so this takes a
long time. If this time exceeds the diurnal radiation cycle time, steadystate heat transfer to the underlying ice-surface is never achieved
(Reznichenko et al., 2010). The major effect of a thick rock avalanche
deposit on a glacier ablation zone is thus to dramatically reduce the
surface ablation beneath the debris by comparison with adjacent bare
or thinly-debris-covered ice (see next section). This leads to
thickening of ice (in fact, a reduction of ice thinning) under the
deposit relative to neighbouring clear ice, which begins immediately
and is often clearly evident a year after emplacement of the deposit. If
the deposit thickness exceeds 2–3 m, the underlying ice-surface
ablation reduces effectively to zero (Bozhinskiy et al., 1986). In this
situation the rate of relative ice thickening beneath the deposit equals
the surface ablation rate of clear ice at that location.
A rock avalanche deposit also contributes a large mass of sediment
to the debris transportation system of the glacier almost instantaneously, whereas the gradual exhumation of melt-out debris cover
takes place slowly and quasicontinuously without adding mass to the
system.
If the glacier surface is crevassed, part of the rock avalanche
deposit may enter the sub- and englacial transportation systems of
the glacier (Hewitt, 2009). However, most of the debris remains
supraglacial and will eventually be carried to the terminus by glacier
flow (e.g., the Sherman Glacier, Alaska; Figs. 1 and 4).
Our study showed that sandy supraglacial cover 5 cm thick
(representing melt-out debris) reduced melting to 50% of that with
bare ice under same conditions, and 9 cm of sand reduced ablation to
25% (Fig. 5A). Using rock avalanche debris instead of sand under
diurnally-cyclic radiation (Reznichenko et al., 2010) increased the
insulating effect substantially. The presence of rock avalanche debris
also significantly reduced rainfall-induced ice-surface ablation,
whereas this effect did not occur with sand (Fig. 5B).
In summary, rock avalanche debris cover on a glacier ablation zone
dramatically reduces total ablation beneath the debris; the reduction
is much less under normal (melt-out) debris cover.
3.4. Field studies
In order to compare the laboratory results of Reznichenko et al.
(2010) with effects in the field, we surveyed rock avalanche deposits
from 1991 Aoraki/Mt. Cook and 2004 Mt. Beatrice on the Tasman and
Hooker Glaciers, respectively (Southern Alps of New Zealand, Fig. 6).
Both avalanches were typical of aseismic slope mass failures in the
Southern Alps (McSaveney, 2002; Cox et al., 2008).
Ground penetrating radar (GPR) surveys of these deposits were
carried out to measure the effect of rock avalanche debris on the
surface-ice ablation. GPR images of two perpendicular profiles along
the rock avalanche deposit and adjacent clear or debris-covered ice
provided information on the thickness of the rock avalanche deposits,
the elevation of the underlying ice surfaces, the modification of the
deposits since emplacement, and the interaction of the rock avalanche
deposits with adjacent clear or debris-covered ice.
3.3. Insulation mechanisms
In order to study the role of the debris cover in insulating the
underlying ice from melting, we conducted laboratory experiments in
which ice blocks with clear surfaces and identical sizes covered with
either medium sand (representing typical highly permeable melt-out
debris) or rock avalanche debris were exposed to identical radiation,
both steady and diurnally-cyclic (Reznichenko et al., 2010). These
experiments showed that the insulating effect of debris on ice
depends on the occurrence of cyclic radiation input. For example,
under specific laboratory conditions with continuous, steady-state
radiation, sand cover greater than 5 cm thick delays the onset of icesurface melting by more than 12 h as the sand layer warms up, but
after that heat transfer through the debris is steady and ablation rates
become similar to those for bare ice. By contrast, when realistic
diurnal cycles of radiation are imposed the debris cover continues to
reduce the rate of ice melt over the long term (Fig. 5A).
3.4.1. Mt. Beatrice rock avalanche deposit
The 2004 rock avalanche from Mt. Beatrice (2528 m) was
emplaced onto the debris-free part of the ablation zone of the Hooker
Glacier, close to an earlier deposit that presumably originated from
the same source (Fig. 6A). Satellite images show that between 2004
and 2009 the deposit moved about 800 m down the valley, indicating
an average ice flow velocity of about 160 m/y at that location, which is
similar to previously estimated velocities at the same altitudes
(higher than 1200 msl.) on the Tasman Glacier (Quincey and Glasser,
2009).
In December 2009 we obtained the longitudinal and cross-profiles
of the Mt. Beatrice deposit shown in Fig. 6A. Profiles revealed that five
years after emplacement an elevated ice platform, varying between 20
and 30 m above the adjacent upstream and downstream ice surfaces,
respectively, was present under the 3–7 m thick rock avalanche
deposit (Figs. 6 and 7), indicating an average clear-ice ablation rate of
Fig. 4. The 1964 Sherman Glacier rock avalanche deposit 44 years after emplacement on 2008: (A) the rock–avalanche cover travels supraglacially; and (B) and is carried to the
terminus by glacier flow, where it is dumped to form a moraine (photos by M.J. McSaveney).
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331
Fig. 5. (A) Example of surface lowering of bare ice and ice under 1, 5, 9, and 13 cm of debris cover under diurnal-cycle conditions. (B) Comparison of the surface lowering of the bare
ice, ice under debris cover of sand and rock avalanche material 9 cm deep with diurnal-cycle conditions and diurnal rainfall effect (Reznichenko et al., 2010).
Fig. 6. Location of the two rock avalanche deposits where GPR studies were conducted (red lines indicate GPR profiles, black arrows indicate ice flow directions, satellite image of 2001,
LandSat): (A) Mt. Beatrice (2004) rock avalanche deposit on the Hooker Glacier, 2007 (aerial photo by S. Winkler); and (B) Aoraki/Mt. Cook rock avalanche (1991) deposit on the Tasman
Glacier with Hochstetter and Tasman ice indicated (satellite image 2009, ASTER).
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Fig. 7. GPR profile along the 2004 Mt. Beatrice rock avalanche in December 2009; the rock avalanche deposit thickness reaches about 7 m, and relative ice thickening is about 20 and
35 m at the upstream and downstream sides of the deposit, respectively. (A) The Mt. Beatrice rock avalanche on the Hooker Glacier deposit 5 years after emplacement; and (B)
The ~ 20-m-high upvalley slope of the deposit looking downvalley.
4–6 m/y. The average debris thickness we estimated to be 3–5 m,
giving a volume of 1.4–2 × 10 5 m 3. However, as the debris is carried
downglacier, the increasing relative ice thickness causes steep slopes
to form along the sides of the deposit. These slopes grow in height and
backwaste from melting over time, redistributing the deposit on the
edges. As a result, the elevated area of the avalanche deposit cover
reduces with time, and the estimated volume decreases as material is
shed to the adjacent glacier surface. The GPR data also indicate
shearing of the thickened underlying ice owing to the differential ice
flow (Fig. 7).
3.4.2. Aoraki/Mt. Cook rock avalanche deposit
The 1991 Aoraki/Mt. Cook rock avalanche, whose volume was
originally estimated at 11.8 ± 2.4 × 10 6 m 3 (McSaveney, 2002), travelled across the Grand Plateau, down the Hochstetter icefall, and
spread onto and across the Tasman Glacier (Fig. 6B). The motion of the
avalanche deposit following emplacement was complicated. Shortly
after deposition, the Hochstetter ice stream deformed the original
avalanche deposit into two main parts: that on the Hochstetter ice had
moved more than 2200 m down the valley by 2009, whereas that on
Tasman ice had moved only 500 m (Fig. 6B). On the latter deposit we
obtained two perpendicular GPR profiles through the avalanche and
the adjacent upglacier, debris-covered ice.
The Aoraki/Mt. Cook rock avalanche deposit (Fig. 8) is up to 10 m
in thickness and a 25-m-high debris-covered ice ridge has formed at
the upstream edge of the deposit; this was not present immediately
after emplacement (McSaveney, 2002). The GPR survey showed
meltwater channels under the deposit, which can cause local
collapses. Therefore, caused by differential ice flow velocities in the
deposition area, complications from tributary ice flow and these
meltwater channels, the relative ice thickening from reduced ablation
is hard to estimate.
Similar ice thickening has been reported on other glaciers, with the
formation of platforms up to 30 m above surrounding ice surfaces
from reduced ice-surface ablation (Post, 1968; Molnia, 2008). Hewitt
(2009) reported that the rock avalanche deposit onto the Bualtar
Glacier, Karakoram, increased the surface elevation by 5–15 m in one
year by comparison with adjacent ice. Field data, thus, confirm the
experimental result that rock avalanche debris dramatically reduces
ice-surface ablation.
4. Effects of rock avalanches on glacier dynamics
4.1. Mass balance
The net mass balance of a glacier is determined by ice-mass gain in
the accumulation zone and ice-mass loss in the ablation zone. A
positive net mass balance over a certain period will cause glacier
thickening and frontal advance, while negative net mass balance (i.e.
ablation exceeds accumulation) will usually cause the glacier to thin
along its profile and the terminus may retreat. Both advance and
retreat have response times specific to each glacier. However, in some
circumstances (e.g., a long, low-gradient glacier tongue with
extensive debris cover), negative net mass balance can cause vertical
surface lowering while the terminus remains stationary, as with the
downwasting Tasman Glacier (Kirkbride, 1993).
Two main changes in the glacier system result from the
emplacement of a rock avalanche deposit on the ablation zone
surface:
(i) the net mass balance becomes less negative or more positive by
reduction of ice-surface ablation in the ablation zone, assuming
there is no change in accumulation rate; and
(ii) additional mass is provided (a) immediately, by the rock
avalanche deposit, and (b) gradually, by the increasing relative
ice thickness caused by reduced ablation under the rock
avalanche deposit.
Rock avalanche debris deposited in the glacier accumulation zone
rapidly becomes buried by snow and ice (and entrained in the
englacial transport pathway), so it only affects the ablation rate for a
short time. It will most likely emerge in the ablation zone as melt-out
englacial or subglacial debris after being transported and dispersed in
the en/subglacial transport pathways.
The reduction of total ablation under a rock avalanche deposit can
be large (e.g., Reid, 1969; McSaveney, 1975; Reznichenko et al., 2010;
Alexander et al., 2011). Ice-surface ablation usually reduces to close to
zero beneath a rock avalanche deposit. While sub- and englacial
ablation by meltwater heat flow, heat conductivity and other
processes will be unaffected by the surface debris; they contribute
only in a minor way (~ 10–20%; Alexander et al., in review) to total
ablation. Thus, if a sufficient area of the ablation zone is covered, the
glacier net mass balance can be significantly altered. Because the
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Fig. 8. GPR profile along the 1991 Aoraki/Mt. Cook rock avalanche and adjacent debris-covered glacier surface in February 2009, the Tasman Glacier, Southern Alps, NZ. The thickness
of the rock avalanche deposit is up to 10 m, and the relative ice thickening at the upglacier edge of the deposit is about 25 m. (A) The ridge formed at the upvalley edge of the deposit
with the Hochstetter icefall in the left background; and (B) GPR measurements on the upvalley slope of the Aoraki/Mt. Cook rock avalanche deposit.
clear-ice ablation rate increases toward the terminus (Schytt, 1967)
and the rock avalanche deposit will be progressively carried downglacier, the reduction in surface ablation per unit area will increase as
the debris approaches the terminus. At the same, time the upglacier
limit of the debris cover is moving downvalley and clear-ice ablation
thereby increasing, so the net variation of mass balance with time is
not simple. For the present we note only that when the rear of the
debris cover has advected downvalley of the original terminus
position, the glacier mass balance is re-established and the debriscovered ice downvalley thereafter behaves independently of the
glacier.
4.2. Glacier response to positive mass balance
A rock avalanche covered 50% of the ablation zone of the Sherman
Glacier to an average depth of 1.5 m during the Alaska earthquake of
1964, and ablation beneath the deposit was reduced by about 80%
(McSaveney, 1975). The resulting alteration in the net mass balance
changed previous glacier retreat to steady-state conditions and, after
some years, a slow advance became evident (Post, 1968; McSaveney,
1975), which still continued in 2009 (M.J. McSaveney, GNS Science,
PO Box 30368 Lower Hutt, NZ, pers. comm. 2009). A rock avalanche
onto the Bualtar Glacier, Karakoram, in 1986, caused substantial
relative ice thickening under rock avalanche debris covering only 15%
of the ablation zone, but the positive impact on mass balance was
equivalent to a 20% increase in annual accumulation (Hewitt, 2009).
As a result, the Bualtar Glacier (which had been retreating for several
decades) surged during first 2 years after the rock avalanche and
continued to advance during the following 12 years.
The mass added to the glacier both immediately (rock avalanche
debris) and gradually (relative ice thickening) will tend to increase
the glacier flow velocity. Firstly, the rock avalanche material itself is
equivalent to about double the weight of the same volume of ice
(Shulmeister et al., 2009). Secondly, the relative thickening of ice
under the deposit is added to the normal (ice) mass of the glacier. For
example, the area of glacier covered by the Mt. Beatrice rock
avalanche 5 years after emplacement has the equivalent of about
2 × (1.4–2 × 10 5) m 3 additional ice weight from the rock avalanche
deposit plus about 13 × 10 5 m 3 of ice (if the average additional ice
thickness is 27 m), which is increasing annually. Totally it is
equivalent of an extra 17 × 10 5 m 3 of ice on an area of about 5 ha, or
34 m of ice depth. This additional mass must alter the glacier
dynamics by increasing basal sliding, and this will occur in proportion
to the percentage of the ablation zone affected and in inverse
proportion to the original ice depth.
The sliding velocity of a warm-based valley glacier is proportional
to the square of the bed shear stress, which depends on the ice-surface
slope and the mass of the ice per unit area of bed (Paterson, 1994).
Thus 10% additional ice depth should increase flow velocity by 20% at
a given location. In addition, extra sediment finding its way to the
glacier base via crevasses and moulins may affect the subglacial
topography and drainage and hence the basal water pressure, which
in turn affects sliding (Turnbull and Davies, 2002; Davies and Smart,
2007); however, the magnitude of this effect is difficult to determine.
This effect was first observed after the St. Elias earthquake by Tarr and
Martin (1914), who proposed the “earthquake advance theory” when
seven glaciers were surging and advancing 10 years after the event.
The advances were attributed to added mass from rock, snow, and ice
avalanches, and also to the uplift of the St. Elias Mountains by 14 m
during the earthquake and the corresponding increase in glacier
gradients (McSaveney, 1975).
The addition of sufficient rock avalanche debris to a glacier
evidently can have significant effects on both mass balance and glacier
motion (Shulmeister et al., 2009). An equilibrium glacier may
advance; a retreating glacier may retreat more slowly, or come into
equilibrium, or advance; and an already advancing glacier may
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advance more rapidly. A surge-type glacier may surge prematurely
(Hewitt, 1988).
In case of the Mt. Beatrice rock avalanche onto the Hooker Glacier,
no changes in glacier behaviour have been observed, because the area
covered by the deposit was small proportionally to the ablation area.
Shulmeister et al. (2009) estimated that glacier behaviour would be
altered by a rock avalanche deposit that covered more than about 10%
of the ablation zone area. With smaller deposits, the glacier may
experience local effects from the additional supraglacial load, but the
overall mass balance will not be significantly affected. The Mt. Beatrice
rock avalanche covers about 0.05 km 2 of the roughly 7 km 2 of the
ablation zone for the Hooker Glacier, which is proportionally b 1% of
the area (Fig. 6A). The Aoraki/Mt. Cook avalanche deposit covered
about 1.7 km 2, b4% of the ablation zone of the Tasman Glacier (about
45 km 2; Kirkbride, 1989). The deposit has since been redistributed
and compressed and now covers an even smaller area (Fig. 6B).
Therefore, these two deposits are not expected to noticeably affect the
overall behaviour of their glaciers.
4.3. Modelling
The effects of rock avalanche debris on glacier behaviour can be
predicted quantitatively. Alexander et al. (2011) developed a simple
steady-state mass balance model, calibrated on the modern (1999–
2010) quasisteady-state conditions of the Franz Josef Glacier and
successfully proven against its Little Ice Age terminal moraine
positions and trim lines. This was used to test the hypothesis of
Tovar et al. (2008) that a catastrophic rock avalanche could have
caused the advance of the Franz Josef Glacier (western Southern Alps,
New Zealand; Fig. 9) to the location of the Late Pleistocene/Early
Holocene Waiho Loop moraine; and thus have formed the moraine.
The modelling shows that a steady-state terminus position at the
Fig. 9. An example of the modelled rock avalanche debris cover on the Franz Josef Glacier, NZ, showing percentage of ablation reduction, which is staggered below the ELA to a
terminus position at the Waiho Loop, assuming temperatures 2 °C less than the present-day, with an ELA of 1480 m asl (satellite image of 2001, LandSat). The staggered levels of
ablation reflect the accumulation of large quantities of debris toward the terminus, while in the upper-ablation zone, ablation rates will return to those of clean ice once the debris
has moved downvalley. Location of the Franz Josef névé is also shown in Fig. 6.
N.V. Reznichenko et al. / Geomorphology 132 (2011) 327–338
Waiho Loop requires mean temperatures about 4–5 °C cooler than
present-day. Anderson and Mackintosh (2006), using an ice flow/mass
balance model, found that an advance to the Waiho Loop required
3–4 °C cooling from present-day. The results from both studies
contradict proxy data, which indicate temperatures no more than
~ 2.5 °C cooler than present-day during the last glacial-interglacial
transition (LGIT) about 12,000–13,000 YBP (Carter et al., 2003; Turney
et al., 2003; Barrows et al., 2007). Direct dating of the Waiho Loop
moraine by Barrows et al. (2007) yielded an age of about 10,500 YBP,
corresponding to significantly warmer temperatures than those that
occurred during the LGIT and making a climatically driven advance
even less likely.
To explain the advance of the Franz Josef Glacier from an
equilibrium position close to Canavan's Knob to the Waiho Loop
position during a time without any cooling (but ~2° cooler than
present), Tovar et al. (2008) suggested, based on the sedimentology of
the Loop moraine, that the ablation zone of the glacier was covered
with debris by a rock avalanche. Alexander et al. (2011) showed that
65% reduction in total ablation (or a gradual reduction from 20 to 80%
along the glacier) is needed to cause the required 3 km of advance
from near Canavan's Knob to an equilibrium position at the Waiho
Loop moraine (Fig. 9), assuming constant climate; based on the above
data this appears reasonable. A coseismic rock avalanche of the
required volume in the Waiho valley is certainly possible; Tovar et al.
(2008) provided evidence of sufficiently large deep-seated, sourcearea scars in the correct location to match the dominant Loop
lithology. Similar events have been reported elsewhere (e.g., Post,
1967; McSaveney, 1975, 2002; Gordon et al., 1978; Evans and Clague,
1988; Orombelli and Porter, 1988; Gardner and Hewitt, 1990; Jibson
et al., 2004, 2006; Lipovsky et al., 2008; Deline, 2009; Hewitt, 2009).
Vacco et al. (2010) recently presented a one-dimensional
numerical model of the Franz Josef Glacier. Though approximate, it
confirmed that a debris driven advance could reach the Waiho Loop
under certain circumstances and, significantly, that the resulting
deposition would have a prominent, high end moraine with an
extensive blanket of ablation moraine to its rear, as appears to be the
case at the Waiho Loop based on available evidence (Shulmeister
et al., 2010).
In summary, there is substantial evidence that a rock avalanche
deposit covering a sufficient proportion of an ablation zone can cause
glaciers mass balance fluctuations and its nonclimatically driven
advance.
5. Consequences
5.1. Terminal moraine formation
Here we describe theoretically how a rock avalanche may cause
glacial advance and frontal moraine formation without any hindrances. A rock–avalanche-driven glacier advance will be short-lived;
perhaps in the order of decades, compared to centuries to a
millennium in the case of an advance driven by climate such as the
Younger Dryas or the Little Ice Age. As the rear of the debris deposit is
carried downvalley by ice flow, ice-surface ablation will increase and
mass balance will gradually become less positive. It will return to
normal conditions, or the mass balance as before the event, when all
the debris has been carried past the original terminus. Then the
advance of the detached debris-covered ice will eventually slow and
stop, during which time substantial quantities of rock avalanche
debris will reach the new terminus and form a moraine. The sequence
of events is illustrated conceptually in Fig. 10; ice covered with rock
avalanche deposit reduces in elevation less slowly because of
suppressed melting than the previous clean or thinly-covered ice, so
it flows more rapidly. In addition, ice at the terminus has compressive
flow which is less rapid than that farther upvalley, so the length of
debris-covered ice reduces and its thickness increases accordingly.
335
Fig. 10. Response of a glacier to rock avalanche debris covering the ablation zone:
(A) glacier in equilibrium prior to rock avalanche; (B) rock avalanche emplaced; total
ablation reduced by ~ 70%; (C) lack of melting increases ice depth and sliding velocity;
rock-covered reach reduces in length caused by relative velocity and ice thickens;
(D) glacier now debris-free to original terminus; debris-covered ice continues to flow
and slowly melt, forming terminal moraine by dumping RA deposit; and (E) all ice
melted, terminal moraine remains with ablation moraine.
Eventually the original glacier ablation zone is debris-free because all
the debris has been carried beyond the original terminus. The debriscovered ice downvalley is now independent of the glacier and
continues to flow slowly under its own surface gradient while slowly
melting, forming a terminal moraine; it eventually becomes stagnant
and slowly melts to form a sheet of ablation moraine to the rear of the
terminal moraine. This form is similar to that generated by the model
of Vacco et al. (2010), and is strikingly similar to that displayed by the
small Classen terminal moraine in New Zealand shown in Fig. 11.
Supraglacial rock avalanche debris is normally between 1 and
10 m thick; this allows estimation of the size of a rock avalanche
required to cover a specified proportion of the ablation zone of a given
glacier and suppress ablation accordingly. For example, the Franz Josef
Glacier ablation zone area is about 6 km 2, so a rock avalanche of the
order of 10 6 m 3 in volume would be needed to cover 10% of the
ablation zone to a depth of N1 m and cause a significant change in
glacier behaviour. Such a volume of rock avalanche deposit arriving at
the approximately 0.4 km wide terminus could form a moraine of
cross-sectional area ~10 3–10 4 m 2; this would certainly be a prominent feature in the landscape.
How moraines are formed by climatic glacier advances is also
incompletely understood. Bulldozing is able to generate small “push”
moraines of the order of several metres high (Molnia, 2008; Winkler
and Matthews, 2010), while sluicing of debris by subglacial water
flow generates characteristic stratified “outwash” moraines (Hyatt,
2010); however, “dumping” of supraglacial debris, which can build
very high and steep lateral moraines (Benn and Evans, 2010) seems to
be the most likely potential cause of very large, steep-fronted terminal
moraines such as the Waiho Loop (Vacco et al., 2010), and the
requirement for large volumes of supraglacial debris to do this
suggests that a rock avalanche may be the most likely source of such
moraines.
336
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5.3. Effects on society
Fig. 11. Classen moraines, Godley Valley, NZ, 2009; the upper terminal moraine has the
form indicated by Vacco et al. (2010) and by Fig. 10 (Photo by S. Winkler).
Given the potential for rock avalanches to generate terminal
moraines, improved understanding of moraine formation processes
by rock–avalanche-driven advances is an urgent requirement.
5.2. Paleoclimatology
Terminal moraines are commonly used as paleoclimatic indicators.
This method is based on the assumption that all moraines owe their
existence to climatically driven advance-retreat sequences, but our
work above clearly shows this assumption to be potentially in error.
Any purely rock–avalanche-generated moraine whose date is
assigned paleoclimatic significance generates an error in the inferred
paleoclimate record. Of course, a rock avalanche can fall onto a glacier
that is in the course of a climatically driven advance-retreat sequence,
in which case the moraine age will have paleoclimatic relevance.
Terminal moraines form when a glacier reaches an advanced
position; when it retreats again the moraine records the previous
maximum extent (Benn and Evans, 2010). If a glacier terminus
undergoes an advance-retreat sequence and is supplied with
sediment by any process, a terminal moraine is likely to form that
will remain after retreat has begun. The greater the volume of
sediment able to be delivered to the moraine, the larger it will be.
Thus, a rock–avalanche-driven advance will be capable of generating a
substantial terminal moraine because it involves a very large volume
of debris. Such a moraine will have no necessary association with any
climatic change (Larsen et al., 2005; Tovar et al., 2008; Shulmeister et
al., 2009). Even if a rock avalanche deposit does not cause a glacier to
advance, the eventual arrival of a large volume of supraglacial debris
at the terminus can form a large moraine in a short time, whereas to
form the same size of moraine without rock avalanche debris would
require the terminus to remain in one position for a relatively long
time.
In the same way, moraines left by surge-type and tidewater
glaciers may record nonclimatic ice-front fluctuations (Yde and
Paasche, 2010). Therefore the terminal moraines that are used as
paleoclimatic indicators must be known to be generated by climatic
variation rather than by nonclimatic processes. To date, however, no
such investigations of this topic have been reported, and all terminal
moraines are assumed de facto to result from climatic variations.
Given the crucial contemporary importance of paleoclimatology in
understanding climate change, this is potentially a serious deficiency.
The various difficulties outlined herein can only be resolved when the
processes of formation of both climatically driven and rock–
avalanche-driven moraines are much better understood, so that
criteria can be developed for distinguishing them in the field.
The potential for large rock avalanches to affect some present-day
glaciers may resulted in significant impacts on society and business,
because of the increasing level of tourism-related human activity in
the vicinity of accessible glaciers (which are often the larger ones that
extend to low altitudes). A good example is the Franz Josef Glacier
because of its location in a tectonically active setting that frequently
experiences intense seismicity (Shulmeister et al., 2009) and because
of its popularity as a tourist destination owing to its easily accessible
terminus (at ~ 300 m asl).
Catastrophic landslides have previously been caused by earthquakes on the Alpine fault, a 700-km-long plate boundary fault
running along the western rangefront of the Southern Alps of New
Zealand (e.g., Chevalier et al., 2009; Dufresne et al., 2010). This fault
generates earthquakes of M ~ 8 several times per millennium on
average, and the last one was in 1717 according to paleoseismic
information (Rhoades and Van Dissen, 2003). The large number of
deep-seated landslide source-area depressions are observed in the
Southern Alps, whose morphology indicates that their origin was
most likely coseismic (Tovar et al., 2008). Shulmeister et al. (2009)
suggested that a large rock avalanche could be expected in the order
of once per millennium on average in the Franz Josef Glacier
catchment. However, given that since 1999 six rock avalanches have
occurred in the Southern Alps, all of the order of 10 7 m 3, and none
were triggered by either earthquake or rain (Aoraki/Mt. Cook and Mt.
Fletcher I and II; McSaveney, 2002; Mt. Adams; Hancox et al., 2005;
Young River; unreferenced; Joe River; unreferenced), the probability
of any kind of rock avalanche in any specific mountain valley is clearly
nontrivial.
The ever-present and immediate hazard from a rock avalanche
onto the Franz Josef Glacier (and perhaps running out along the
proglacial upper Waiho valley) is to the lives of people on the glacier
and in the valley (several hundred visitors at any one time on a fine
day), since warning of such an event may only be a minute or so for
people at the terminus and less for people on the glacier. This
therefore represents a not insignificant threat to peoples’ safety, given
the ~ 250,000 tourists that visit the valley, terminus, and ice surface
annually (Corbett, 2001).
To estimate the potential effect of a rock avalanche on the presentday glacier, Alexander et al. (2011) assumed that debris from a rock
avalanche in the mid-upper catchment is voluminous enough to cover
the entire present-day ablation zone (area about 6 km 2) to metrescale depth . This would require a volume of about 10 7 m 3. If total
ablation were reduced by 80% as a result, the modelling indicates that
the Franz Josef Glacier would advance almost as far as the township
(about 6 km; Fig. 9) if the present climate conditions and accumulation rates prevailed. The speed of a rock–avalanche-driven terminus
advance can be estimated from the known rate of ice accumulation in
the catchment, and appears to be of the order of 1 km/y (Tovar et al.,
2008). Apparently, the glacier could possibly advance to the vicinity of
Franz Josef township in perhaps a decade; human life would not be
threatened by such an advance, but the main highway along the west
coast of South Island would be. The consequent effects on the
behaviour of the already troublesome Waiho and Callery rivers
(McSaveney and Davies, 1998; Davies and McSaveney, 2001) would
also need to be investigated—when the Franz Josef Glacier advanced
about 1.5 km between 1982 and 1999, the proglacial area aggraded of
the order of tens of metres. Clearly, there is the potential for major
disruption of tourism and other business in this iconic tourist
destination.
This scenario has not to date been part of any assessment carried
out for the area, and, therefore, no management strategies are
developed. A similar threat could exist for the locality of the Fox
Glacier, 20 km south of the Franz Josef Glacier, whose glacial and
geomorphological characteristics are similar. Evidently, if a rock–
N.V. Reznichenko et al. / Geomorphology 132 (2011) 327–338
avalanche-generated glacier advance were to approach Franz Josef
township, the carefully planned, ordered, and structured relocation of
the settlement and of the highway are required. No active mitigation
methods appear to be practicable. The requirement for further
investigation of this scenario is clear.
6. Conclusions
(i) Large rock avalanches falling onto glaciers can affect glacier
behaviour significantly by reducing the total ablation beneath
the deposit by up to 90%, causing substantial changes in glacier
behaviour such as advances and surging.
(ii) Rock–avalanche-driven glacier advances can form prominent
terminal moraines.
(iii) Terminal moraines generated by rock–avalanche-driven advances have no necessary climatic association; therefore, the
validity of paleoclimatic inferences from moraines depends on
identifying the origin of the moraine.
(iv) Further investigation is required into the processes of terminal
moraine formation caused by rock–avalanche-driven and
climatically driven advances and into criteria for distinguishing
moraines formed by these processes.
(v) Rock–avalanche-driven glacier advances have the potential to
cause serious disruption to society and business over decadal
timescales.
Acknowledgements
Stefan Winkler's comments on an early draft, together with those
of three anonymous reviewers, led to substantial improvements of the
manuscript, for which we are grateful.
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