Catastrophic landslides, glacier behaviour and moraine formation – A view from an active plate margin

Quaternary Science Reviews 28 (2009) 1085–1096

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Catastrophic landslides, glacier behaviour and moraine formation – A view from an active plate margin James Shulmeister a, *, Tim R. Davies a, David J.A. Evans b, Olivia M. Hyatt a, Daniel S. Tovar a a b

Department of Geological Sciences, University of Canterbury, Private Bag 4800, Christchurch, New Zealand Department of Geography, University of Durham, South Road, Durham DH1 3LE, UK

a r t i c l e i n f o

a b s t r a c t

Article history: Received 5 March 2008 Received in revised form 24 October 2008 Accepted 20 November 2008

The influence of large bedrock landslides (‘‘rock avalanches’’) on the behaviour of glaciers is incompletely recognised. Here we present an example from an active tectonic margin in South Island, New Zealand where large earthquakes leave a significant imprint on glacial records. We demonstrate that terminal moraines on the western side of the Southern Alps record both ‘ordinary’ (i.e. climate-driven) and landslide-initiated glacial advances. Following consideration of the processes involved in rock avalancheinitiated moraine construction we suggest ways of determining the nature of the advance that built the terminal moraine. The implications of these observations are important in breaking the conventional linkage of individual terminal moraines with climate forcing. Ó 2008 Elsevier Ltd. All rights reserved.

1. Introduction The fluctuations of glaciers, especially valley glaciers, are widely regarded as good proxy evidence for climate change, and midlatitude valley glaciers in particular are regarded as weather vanes for global warming (e.g. Oerlemans, 2005). Terminal moraines, the constructional debris ridges produced at the limit of the advance of the ice fronts of these glaciers, are the most widely used evidence for old glacial limits and, by inference, paleoclimate. Ice limits as defined by terminal moraines are routinely interpreted to result from climatically induced glacial advance and still-stand, and thus are used to quantify cooling (e.g. Anderson and Mackintosh, 2006) and/or precipitation changes (e.g. Lemkuhl and Owen, 2005; Shulmeister et al., 2005). A classic example of the climatic interpretation of moraines is the Younger Dryas as represented by its terminal moraine ridges. The Younger Dryas (YD) is the name given to a major cooling event in the North Atlantic region at the end of the last ice age (ca. 11,500– 13,000 cal yr BP). It is recorded in many parts of Europe by glacial moraines marking a significant re-advance (e.g. the Salpausselka¨ moraines of Finland; Tschudi et al., 2000). From non-glacial evidence, including changes in foraminiferal assemblages in the North Atlantic (e.g. Ruddiman and McIntyre, 1981) and the reappearance of alpine and sub-arctic pollen types in temperate NW European pollen records (e.g. Walker et al., 1994), it is clear that the

* Corresponding author. Tel.: þ64 3 364 2762; fax: þ64 3 364 2769. E-mail address: [email protected] (J. Shulmeister). 0277-3791/$ – see front matter Ó 2008 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2008.11.015

YD marked an abrupt cold phase in this region and glacial moraines of YD age are interpreted to reflect this cooling. Evidence for YD cooling outside the North Atlantic area is more controversial but glacial advances in South America (Osborn et al., 1995), and western North America (e.g. Friele and Clague, 2002) have been linked to this event and attributed to abrupt cooling. There have been direct comparisons between the inferred YD Misery Moraines in the Otira River Valley in New Zealand and the YD Egesen Moraine complex in Switzerland (Ivy-Ochs et al., 1999). Much research has also focussed on the climatic implications of the Waiho Loop moraine in Westland, South Island, New Zealand (e.g. Denton and Hendy, 1994; Mabin, 1995; Barrows et al., 2007). In fact the Waiho Loop has been used as prima facie evidence for interhemispheric climate connectivity (e.g. Broecker, 2000) While there is little doubt that the first order controls on the scale of glacier advance are climatic, Hewitt (1999), Porter (2000) and Larsen et al. (2005) have recently brought the contribution of landslides to moraine formation into distinct focus. Hewitt (1999) re-interpreted a large number of reported moraines in the Karakoram as deposits of large rock avalanches; Porter (2000) suggested the same process in the Southern Alps of New Zealand, based on mid-Holocene dates from a moraine in the Mt Cook region; and Larsen et al. (2005) demonstrated significant correspondence between dated terminal moraines in the Southern Alps and the known dates of major earthquakes on the Alpine Fault that bounds the range to the west. This suggests that coseismic landsliding contributed to – and in fact may have dominated – moraine formation. The implication is that a terminal moraine may be created by the occurrence of a rock landslide depositing a large

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volume of debris onto the glacier surface, whereupon it is carried to the terminus and deposited as terminal moraine. Lateral moraines can form in the same way. A connection has also been established between the occurrence of large supraglacial rock avalanches and the occurrence of glacier surges (Gardner and Hewitt, 1990; Hewitt, 2009). The mechanisms involved are discussed in detail later, but it appears that a thick layer of rock debris can both severely reduce ablation, altering the mass balance; and also affect basal friction, leading in both cases to glacier advance and perhaps to initiation of a surge. Hence landslides cannot only contribute significant volumes of material to glacial deposits, they can also significantly and rapidly alter the behaviour of glaciers. In this paper we explore the extent to which glacial deposits on the West Coast of South Island, New Zealand, may reflect the influence of large landslides, and examine the processes involved in order to develop diagnostic tools for distinguishing climatically initiated moraines from landslide-initiated ones.

1.1. The geological and climatic setting of the Western Southern Alps, New Zealand Though this paper deals with the general issue of glacier responses and moraine formation as a result of catastrophic rock avalanches, our specific examples are derived from the West Coast of South Island, New Zealand (see Fig. 1) and this context underpins our observations. The South Island of New Zealand is characterised by a spine of mountains, the Southern Alps, that run SSW to NNE along the southern two-thirds of the island. This range lies close to the west coast for most of its length, rises to nearly 3800 m at Mt Cook/Aorangi, and has many peaks over 2500 m. The range front is defined by the Alpine Fault, a major transcurrent fault with an offset of more than 500 km since the mid-Miocene that is still very active (e.g. Kamp and Tippett, 1993). This fault marks the plate boundary between the Pacific and Australian plates, and both the

physiography and geology of the west coast region of the South Island are controlled by this boundary. Geologically the area can be divided into two zones. Immediately east of and parallel to the Alpine Fault are successive bands of metasediments. Close to the fault, uplift and erosion are pronounced (>5 mm yr1), and marine sediments are being dragged up from depths of 30 km, forming highly metamorphosed schists (Grapes and Watanabe, 1992). The metamorphic grade decreases steadily eastward away from the fault and virtually unaltered Torlesse (Mesozoic) greywackes crop out within 15 km of the fault. West of the Alpine Fault quartz rich Paleozoic sediments and granites of the Australian plate form the country rocks (Cox and Barrell, 2007). Actual outcrop is rare, and is limited to isolated hills and the cores of some hill ranges. The hill ranges are surfaced by complex systems of Otiran and pre-Otiran moraines occupying the interfluves between the major river catchments. Outwash plains associated with the rivers and the most recent glacial troughs dissect the moraine hills and extend to the coast. Extensive mountain glaciation still persists in the Southern Alps. The two largest valley glaciers that flow west of the divide (Fox and Franz Josef glaciers) are sourced from ice-fields in the highest part of the Southern Alps. These systems have been extensively investigated from hydrological and climatological standpoints (Sara, 1968; Fitzharris et al., 1992; Anderson and Mackintosh, 2006). These glaciers have some of the highest turnover rates on Earth with annual precipitation of up to w14 m (Henderson and Thompson, 1999), ice velocities of about 1 m day1 (Purdie et al., 2008), and response times of the order of 7–20 years (Fitzharris et al., 1992; Oerlemans, 1997). Ablation in the terminal regions is in the order of 20–30 m a1, and mean annual temperatures at the termini are about 9  C (Anderson et al., 2008; Purdie et al., 2008). The warm climate combined with the abundant sediments from the tectonically active ranges results in active sub- and pro-glacial river and fan systems (Davies et al., 2003; Davies and Smart, 2007).

Fig. 1. Location map showing the central part of the West Coast of South Island, New Zealand. Much of the specific work discussed in the paper comes from the Waiho Valley (inset box). Gillespies Beach is also marked. The Alpine Fault is marked as a black line with the relative sene of motion on either side of the fault indicated by arrows.

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2. Rock avalanches Large rock avalanches are major contributors to denudation processes in actively uplifting mountains (e.g. McSaveney, 2002; Hewitt, 2009). Magnitude-frequency relationships for rock avalanches in a number of regions show well-defined distributions (Malamud et al., 2004; Korup, 2006) that also apply to rock avalanches triggered by earthquakes. Malamud et al. (2004) concluded that ‘‘. bedrock landslides and massive rock falls are dominant contributors to the long-term erosion in seismically active tectonic zones’’. Hence we expect the sediment deposits from rivers and glaciers in active mountains to originate mainly from large mass movements, which are often, but not always, associated with seismic activity. Since massive rock falls are infrequent in any given valley (Whitehouse and Griffiths, 1983), we expect the resulting sediment deliveries to form deposits whose characteristics reflect occasional, massive sediment inputs, between which the system recovers to lower input rates until the next massive delivery. Davies and Korup (2007) have confirmed this hypothesis in the case of Holocene rangefront alluvial fans in Westland, showing that the long-term fan-head morphology is dominated by sediment pulses from these rare, massive events. Hewitt (2009) found that although a catastrophic rock avalanche lasts only a minute or two ‘‘.its legacy can persist as a morphogenetic influence for millennia or tens of millennia through disturbance of other processes.’’ Herein we offer an equivalent suggestion for glacial moraines, again demonstrating that the morphology of some moraines on the west side of the Southern Alps reflects the occurrence of large rock avalanches onto glaciers. This is not in principle a novel concept – it has been known for many years that a proportion of glacial sediment originates from rock avalanches (e.g. Marangunic and Bull, 1969; Porter and Orombelli, 1980; Evans and Clague, 1988; McSaveney, 2002) – but we now suggest that from time to time supraglacial rock avalanche debris can dominate glacial sedimentation and hence glacier behaviour, leading to formation of moraines that can be distinguished from their more conventionally formed equivalents. Large rock avalanches occur predominantly in mountain ranges with large relief (Korup et al., 2007), which are often actively uplifting, and because these ranges are also where glaciers are often found, it is reasonable to expect that large rock avalanches will interact with glaciers from time to time (Evans and Clague, 1998; Hewitt, 1999). We should therefore expect to find evidence of rock avalanche activity in the composition and morphology of moraines. Rock avalanche debris is sourced from higher elevations than the glacier surface, so it must initially be emplaced supraglacially. However some debris may quickly penetrate to en-and subglacial regions via crevasses and moulins, and with the passage of time increasing amounts of debris from a given rock avalanche will arrive in these regions. What distinguishes the effect of large rock avalanches from that of smaller rockfalls, however, is that a large proportion of the much thicker debris from the former remains on the glacier surface, whereas much of the latter becomes cycled through the subglacial drainage system. We note that snow, and the motion of the glacier, tend to obscure and distort the supraglacial deposit characteristics. Many glacier surfaces are rough, so a thin debris layer, up to a few metres thick, and thus capable of significantly affecting glacier behaviour, may not be obvious once covered with snow. Furthermore, glacier motion, particularly in steep reaches, causes crevassing that can absorb large volumes of debris. Many glaciers have centimetersthick blankets of supraglacial debris, so additional debris may not show a textural contrast. Downvalley glacier motion can distort the deposit so that it soon becomes quite different in appearance to

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normal deposits of such events (Eyles and Rogerson, 1978a,b; Spedding and Evans, 2002; Swift et al., 2006). All of these factors can make it difficult to recognise a rock avalanche deposit a short time after its emplacement onto a glacier, hence it is likely that many such deposits have remained unrecognized in the past. Supraglacial deposits of large rock avalanches have characteristic textures and morphologies which allow them to be distinguished from conventional moraines, whose sediment has been transported englacially or subglacially (Spedding and Evans, 2002; Hewitt, 2009). First, the clasts of a rock avalanche are invariably angular because they are sourced directly from bedrock and because during runout they are intensively fragmented (McSaveney and Davies, 2007). No process occurs during the rock avalanche that can reduce clast angularity. Angularity can be subsequently reduced somewhat by jostling during transport downvalley by the glacier (Hewitt, 2009). Second, the grain-size distribution of rock avalanche debris is invariably fractal, with a fractal dimension close to 2.58 (Crosta et al., 2007; McSaveney and Davies, 2007); the deposit is initially characterised by the presence of very fine grains down to nanometre size. During post-emplacement travel downglacier, the fines content can be depleted by rainfall, runoff and wind (Hewitt, 2009). Third, the surface layer of a large rock avalanche deposit is coarser than the lower layers, forming a blocky carapace (Dunning, 2004). Fourth, when a rock avalanche is emplaced onto ice the deposit morphology is often, but not always, spectacularly digitate or lobate (McSaveney, 1978; Evans and Clague, 1998) with the lobes often carrying a large individual distal clast. Finally, the source area of even a very large rock avalanche is usually monolithological, and therefore the deposit is too; by contrast, sediment that has been transported beneath a glacier may have become mixed with rock sourced from other parts of the glacier catchment. 2.1. Rock avalanche frequency Large rock avalanche deposits (w107 m3 or more in volume) are rare in any given valley in the Southern Alps. Using the West Coast of New Zealand’s South Island as an example, large earthquakes on the 400-km long range-bounding Alpine Fault can be expected a few times per millennium (Rhoades and Van Dissen, 2003), but whether these M w8 events always cause large rock avalanches, and how many, is unknown. Keefer (1984) found empirically that a M ¼ 8 earthquake generates w109–1010 m3 of avalanche debris, and Davies and McSaveney (in press) found that of the coseismic landslides generated by the M ¼ 7.2 Fiordland earthquake in 2003 the largest 1% represent 52% of the total debris volume, and the largest 10% represent 96% of the total debris volume. If these proportions apply to a M ¼ 8 Alpine fault earthquake, we would expect several rock avalanches of  108 m3 to occur in the western Southern Alps. The largest (presumed coseismic) rock avalanche deposit known in New Zealand is the 26 km3 Green Lake event in Fiordland (Hancox and Perrin, 1994). There are two known large rock avalanche deposits at the western front of the Southern Alps. Both have deep-seated sources characteristic of coseismic rock avalanches, but neither appears to be the result of either of the last two Alpine Fault earthquakes in ca. 1620 or 1717 (Wright, 1998; Korup, 2004; Chevalier, 2008). Of course other large rock avalanches may have fallen into valleys or onto glaciers, and all trace of them subsequently been removed by the high rates of denudation characteristic of the area (Hovius et al., 1997). It is also known that rock avalanches can occur in the absence of earthquakes – there have been six such events greater than 107 m3 in volume in the Southern Alps since 1991 (M. McSaveney, personal communication April 2008). Thus we suggest that in any given glacierized valley, we might expect a rock avalanche

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sufficiently large to significantly affect glacier motion about once per millennium on average. This frequency seems low; however a typical glaciation has a duration of tens of millennia, so we can expect tens of large rock avalanches to fall onto any given valley glacier during a glaciation. As noted above, the debris volume involved in these events will probably dominate other new debris sources contributing to moraine formation; but it will not dominate the exhumation of earlier valley fills by glacier advance. Hence we expect that large rock avalanchesourced debris should be common in moraine deposits.

3. Effect on glacier behaviour A large supraglacial landslide deposit has the ability to dramatically alter the behaviour of a glacier by three independent mechanisms: 1. The thick debris layer can reduce ice-surface ablation to close to zero, altering the glacier mass balance significantly and causing ice depth to increase, which in turn increases the downvalley driving force due to gravity (Nakawo and Rana, 1999): 2. The additional weight of the rock debris corresponds to about twice the same depth of additional ice. This also increases the downvalley force, but usually by a smaller amount than ice thickening. 3. The arrival of large additional quantities of rock debris in the subglacial drainage system, or at the basal contact between ice and substrate, can alter the basal resistance to sliding (Davies and Smart, 2007). Whether or not these effects are significant depends largely on the mass of rock deposited by the rock avalanche, and in particular the proportion of the ablation zone covered by the debris to sufficient thickness to severely reduce ablation. If the proportion is small, the effect on the glacier may also be small. Supraglacial rock avalanche deposits are generally of metre-scale thickness (e.g. Jibson et al., 2006), and an average depth of 1 m is quite adequate to suppress ablation significantly (see Figs. 2 and 3). Hewitt (2009) reported that about 20% of Bualtar Glacier was covered by the 1988 rock avalanche, leading to a dramatic thickening and surge advance. If w10% of the ablation area needs to be covered to cause a significant response, the minimum rock avalanche volume required, Qmin, relates to the ablation zone area Ab as follows:

Qmin w0:1 Ab

(1)

3.1. Reduced ablation Nakawo and Rana (1999) summarise previous research on the effect of supraglacial debris on ablation rate. A layer of rock debris less than a few centimetres thick increases ablation due to heat absorption by the low-albedo debris and conduction to the icesurface. A thicker layer sharply reduces ablation due to the low conductivity of the very porous debris (Table 1). A 90% difference between ablation rates of clean ice and debriscovered ice was recorded by Purdie and Fitzharris (1999) on the Tasman Glacier in New Zealand based on direct measurements of clean ice ablation and estimates of conduction rates under a 1.1 m debris cover using the methodology of McSaveney (1975) and Kirkbride (1989). Purdie and Fitzharris (1999) examined the effect of advected heat by percolating rainfall through the debris mantle and concluded that it accounted for only w1% of the total ablation. The key factors they identified were the temperature of the water when it contacts the ice, and the rainfall intensity. The Tasman Glacier terminus is at a slightly higher elevation (700 m) than those of the larger West Coast glaciers (300 m), and has somewhat (about 40%) lower rainfall values but it is clear that this effect is minor. Heat advection by percolating rainfall will be even less with the debris of a large rock avalanche, because it has a high fines content and fractal grain-size distribution (McSaveney and Davies, 2007) and will therefore be much less permeable to water than normal supraglacial debris. Ablation due to subglacial drainage will continue to occur with a thick debris mantle but although meltwater channels have been a focus for understanding glacial flow regimes (e.g. Harbor et al., 1997) their direct contribution to ablation is unquantified. When pro-glacial lakes form at the front of down-wasting debriscovered ice tongues, the rate of ablation increases by orders of magnitude due to a combination of thermal erosion of the ice sole and calving (Purdie and Fitzharris, 1999; Ro¨hl, 2006). Rock avalanche debris on glaciers is usually metre-scale in thickness (the Sherman Glacier avalanche of 1964, one of the thinnest reported, had an average depth of 1.2 m; McSaveney, 1978), so its effect is to reduce surface ablation significantly. Our discussion is confined to those sections of glaciers below the equilibrium line altitude (ELA). Glacial reaches above the ELA will retain snow cover, which will rapidly bury landslide debris in ice and return ablation

Fig. 2. Rock avalanche debris on McGinnis Glacier, Alaska caused by the 2002 Denali earthquake. Foreground is supraglacial rockfall debris, background is rock avalanche deposit. Figure (circled) indicates depth of rock avalanche deposit. From Jibson et al. (2006).

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Fig. 3. Rock avalanche onto Black Rapids glacier caused by the 2002 Denali earthquake. Debris volume w12  106 m3: average debris depth is w3 m. From Jibson et al. (2006). Note that as of August 2008 this glacier appears not to have responded to these rock avalanches (J.J. Clague, personal communication, July 2008).

rates to those of clean ice. Below the equilibrium line, rock avalanche debris reduces surface ablation, so the mass balance alters from strongly negative to less negative, and the glacier will respond by increasing its flow velocity, which will in turn tends to cause an advance. The debris loading and ablation rates alter very rapidly, but it may take a considerable time, perhaps months to years, for the accumulating ice to noticeably affect glacier behaviour. Debris cover is common on alpine glaciers (e.g. Nicholson and Benn, 2006), so the significance of landslide debris may be questioned. In New Zealand, many if not most modern glaciers have debris cover in their lower reaches, including Whymper, Burton, Victoria, Balfour and La Perouse glaciers on the west side of the Southern Alps and Meuller, Hooker, and Tasman Glaciers on the east. The average debris thickness in these cases is typically very much less than the metres associated with large rock avalanches emplaced onto glaciers (e.g. Hewitt, 2009). Further, most New Zealand glaciers that are currently debris-covered in their lower reaches are either thinning in situ without retreating or are calving into pro-glacial lakes, behaviour which follows the thinning phase (Purdie and Fitzharris, 1999). Though this process is producing extensive debris cover, it is associated with virtually stagnant ice, and is not analogous to a catastrophic landslide deposit. It is not considered further here, though it is noted that this exhumed debris can also strongly reduce net ablation. 3.2. Ice thickening as a result of thick debris cover Reduced melting under a thick debris cover causes the debriscovered ice to become thicker than adjacent clean ice. Hence the effect of the debris blanket is to alter the mass balance. If an area Ab

[m2] of the ablation zone, whose total area is Aa, is blanketed with debris so that ablation beneath it is reduced by P%, and if the rate of production of ice in the accumulation zone is Qi [m3 a1], then the total ablation rate reduces by PQi(Ab/Aa) [m3 a1]; this is the extra ice accumulating beneath the debris (volume per unit time). The ice beneath the debris will thus thicken by PQi(Ab/Aa)/Ab ¼ PQi/Aa [m a1] as long as the debris is in place. Using Franz Josef Glacier as an example, its annual ice production Qi is of the order of 2  108 [m3 a1] (Tovar et al., 2008 supplementary data), and its ablation zone area Aa is presently about 6  106 [m2]. Hence the rate of ice thickening under a debris blanket will be P  2  108/ 6  106 ¼ 33P [m a1]. We can test this estimation procedure by comparison with measured thickening of Bualtar Glacier in the Karakoram Himalaya as a result of a 107 [m3] rock avalanche in 1988 (Hewitt, 2009): Qi w 6  107 [m3 a1] (Hewitt, 2009) and Aa is 15  106 [m2]. The thickening rate should be 6  107 P/15  106 ¼ 4P [m a1]. Hewitt (2009) noted that it was reasonable to assume that 4.5 m of ablation had been prevented in the debriscovered area on an annual basis. Our estimate is of the correct order of magnitude, provided that the value of P is high (close to 100%), thus it appears that Franz Josef Glacier would thicken by w30 [m a1] beneath the debris of a large rock avalanche. This conclusion is supported by recent measurements of ablation on the nearby and similarly sized Fox Glacier (Purdie et al., 2008); these give an average rate of about 80 mm/day, or 30 [m a1], and earlier data from Franz Josef Glacier are of the same order of magnitude (Owens et al., 1992).

3.3. Increased ice velocity Table 1 Effect of debris depth on ablation (after Nakawo and Rana, 1999).

At a given section the sliding velocity Vs of a valley glacier is proportional to the square of the bed shear stress (Patterson, 1994):

Depth (m)

% reduction

0.00 0.05 0.10 0.20 0.40

0 42 64 79 87

Vs ¼ k1 s2

(2)

where s ¼ (ice load per unit area þ debris load per unit area)sin q, q is the ice-surface gradient and k1 is a constant. In addition to sliding, glacier ice can also deform internally; the surface velocity Vp

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resulting from this mechanism (relative to the basal ice velocity) is proportional to the 4th power of bed shear stress (Patterson, 1994):

Vp ¼ k2 s4

(3)

where k2 is a constant. Hence the total surface velocity Vt is given by

Vt ¼ k1 s2 þ k2 s4

(4)

In warm, wet climates such as that of the West coast of New Zealand’s South Island (annual precipitation w10 m water equivalent), sliding is likely to dominate glacier motion; the average daily velocity of Franz Josef Glacier is w1 m/day (McSaveney and Gage, 1968; Purdie et al., 2008). We can estimate the increase of sliding velocity due to debris and ice loading in the case of Franz Josef Glacier: the ice is 100 m thick (Anderson et al., 2008), and 5 m of debris (with a bulk density about twice that of ice) is deposited on it, causing an increase in ice thickness of 30 m, then the sliding velocity increases by a factor of [{100 þ 30 þ (2  5)}/100]2 ¼ 2.0 from Equation (2). The debris load is responsible for 16% of this increase; this proportion decreases if the initial ice thickness is greater and vice versa. In drier, colder glaciers internal deformation may dominate the glacier motion; a similar event on Bualtar Glacier would increase an initial ice thickness of 100 m by say 10 m and the assumed internal-deformation-dominated sliding velocity by a factor of [{100 þ 10 þ (2  5)]/50]4 ¼ 2.1 from Equation (3). In fact, Hewitt (2009) reports that Bualtar Glacier surged as a result of the rock avalanche deposit, its velocity increasing by an order of magnitude from the pre-event rate of 0.8 m/day; motion clearly entered another phase, perhaps changing from internal deformation to sliding. These estimates are averages of conditions and responses over the whole of the ablation zone. In reality the clean-ice ablation rate increases from zero at the ELA to a maximum at the terminus, and the thickening and velocity increase will be distributed accordingly. The additional detail resulting from consideration of these distributions can be evaluated through numerical simulations, but these do not yet seem to be sufficiently reliable to provide significantly better data than averaging in our opinion. Our intent is to demonstrate that our order-of-magnitude estimates of thickening and velocity increase are significant in the context of terminal moraine formation, and our averaged data certainly suffice for this.

3.4. Increase of subglacial sediment In addition to increasing the downvalley driving stress, a rock avalanche deposit on a glacier surface likely increases the input of debris to the subglacial drainage system. Davies and Smart (2007) recently carried out the first study of the effect of input of coarse material to subglacial drainage systems, in particular, the propensity of closed-conduit-flows to block when transporting bedload, might have on the pressure regime at the base of the glacier and hence the resistance to sliding. They found that the basal water pressures of the glacier could be significantly affected by restriction or blocking of subglacial drainage by sediment, with corresponding effects on sliding velocity. Hence it is possible that an increase in coarse sediment entering the subglacial drainage system could cause the sliding velocity to increase to a greater extent than Equations (1)–(3) suggest. We note however that the converse may also be true: if debris is added to the basal layers of a glacier without significantly affecting the basal water pressure, it is likely to increase resistance to sliding and reduce flow velocities in the ice itself. The critical factor is whether the additional debris interferes with subglacial drainage. If it does, than the addition of debris can

reduce effective stress and frictional resistance by increasing pore water pressure at the ice–sediment or ice–bedrock interface. As pointed out by Hewitt (2009), the ice velocity of a glacier may increase by an order of magnitude due to increase in ice- and debris- loading, causing a surge, which is often associated with a significant and rapid advance of the terminus. Turnbull and Davies (2002) showed how sediment buildup can occlude substrate depressions and allow glacier ice to slide over them, whereas ice previously flowed into the depressions and glacier motion was only possible by internal deformation or pressure melting and refreezing of the ice. Thus the arrival of large quantities of debris could cause a glacier to surge by increasing the ability of the glacier to move by basal sliding. Presumably the glacier in this situation would eventually have surged in response to the gradual accumulation of debris in the depressions by slower processes. It may be relevant, in this context, that Bualtar Glacier is known to have surged prior to the 1982 rock avalanches (K. Hewitt, personal communication, 2009). Moraines resulting from surges may be different from those produced by slower advances, but ‘‘No single landform or diagnostic criterion has yet been found with which to identify surging glaciers’’ (Evans and Rea, 1999), so any differences are probably slight. In summary, emplacement of rock avalanche debris onto a glacier can cause substantially increased ice thickness and flow velocity. The increase in flow velocity in particular indicates that a period of advance, or slowing of retreat, is likely to occur. In addition, there is reason to expect that glaciers capable of surging may surge in response to, and soon after, significant rock avalanche debris input. A major difference between glacier motion affected by arrival of a large quantity of rock debris, and glacier motion affected by arrival of different quantities of ice due to climatic variation, is that the former occurs much more rapidly than the latter. Individual high snowfall/low ablation years will cause small scale advances but this process must occur over a number of years or decades to cause a noteworthy advance. Alteration of glacier behaviour due to a large rock avalanche thus takes place much more quickly than an equivalent climate-driven alteration; we show later that a landslide-driven advance probably lasts for decades to centuries. 4. Formation of terminal moraines Terminal moraines comprise a mix of material dumped from supraglacial transport and materials transported sub-, en- and proglacially. In the perhumid western Southern Alps, terminal moraines are fan-head deposits (Tovar et al., 2008) most of which consist of sediment generated by subglacial fluvial recycling of sediment. This facies association is capped by a small volume of dumped supraglacial material. An example of this stratigraphy is visible in the so-called ‘‘1750’’ moraine in the Waiho Valley, constructed by Franz Josef Glacier (see Fig. 4 for location and Fig. 5 for stratigraphic detail). The moraine consists of decimeter-scale sheets of rounded, crudely stratified and moderately well sorted gravels and coarse sands, capped by a few large boulders. The upper beds of gravel and sand are slightly deformed on the up-ice side of the exposure. In colder and less humid settings, subglacial and proglacial deposits may be dominated by bulldozed and thrust materials. A glacier with a thick layer of rock avalanche debris on its surface will transport that debris to its terminus whether or not the debris load is sufficient to trigger an advance. En route to the terminus some of the debris will be incorporated into the subglacial system via crevasses and moulins. The thickness of the debris cover and the extent of the original rock avalanche run-out, along with

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Fig 4. Lithology and location map for the Waiho Valley. Major rock types, the Alpine fault and major geomorphological features mentioned in the text are marked. The histogram and pie chart information refers to pebble counts (n ¼ 50) from sample sites used. The histograms show angularity data from A = angular to R ¼ rounded. Note the high values of sub-rounded (SR) material in the Waiho River in comparison to the relatively high values of angular and sub-angular (SA) material from the moraines. The pie charts show lithological composition. Note the dominance of greywacke in the Waiho Loop moraine. Data in charts comes from Tovar et al. (2008, supplementary data).

the topography of the glacier, will determine the proportion of material that remains in a supraglacial position. Once the debris reaches the terminus, this supraglacial material will be dumped and a moraine formed. While the glacier is advancing it reworks its own outwash. Substantial aggradation of the forefield by Franz Josef Glacier during its 1982–1999 advance required reworking of a large volume of outwash from beneath the glacier (Davies et al.,

2003). The size of the moraine reflects the magnitude of the rock avalanche that caused the advance and the length of time the terminus is stationary. The length of time that a terminus needs to be stationary to construct a moraine can be illustrated using the Waiho Loop moraine of Franz Josef Glacier (Tovar et al., 2008) The moraine contains a minimum of w108 m3 of debris. It is about 5 km long,

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Fig. 5. Typical fan-head moraine systems in the modern Waiho Valley. Sample locations are shown on Fig. 4. Top: the modern glacier margin with a recent moraine which has a core of outwash gravels and a cap of about 1 m of supraglacial dumped material. Bottom: the 1750 moraine; more than 10 m of stratified gravel are exposed in the core of this moraine whereas the supraglacial debris is limited to a few large boulders. The materials in both moraines are relatively well rounded in comparison to the Waiho Loop moraine (see Fig. 4). Summary angularity and lithology data for these moraines are presented in the middle of the figure (raw data presented in Tovar et al. (2008, supplementary data). Both samples are fan-head material near the base of the sections; clasts are rounded and dominantly schist.

which is the minimum length of the ice-front that formed it. If the ice velocity at the terminus was w1 m/day, then an area of about 1.5  106 m2 of ice was melting each year. If debris was 5 m thick (the average thickness estimated by Tovar et al., 2008) it would take w25 years to construct the moraine. This is clearly an order-of-magnitude estimate, but we conclude that construction of a large terminal moraine by this process may require only decades. When the rock avalanche debris has been fully or substantially transported off the ice, the glacier’s mass balance becomes more negative. The terminus region is most sensitive to this mass balance change and consequently the preservation potential of the rock avalanche-induced terminal moraine is high, at least in the short term, as a rapid retreat of the terminus is likely. 5. Sedimentary and geomorphic characteristics of rock avalanche-generated moraines Heim (reported in Hewitt, 1999) identified eight diagnostic features of rock avalanches, ranging from lithologically uniform bands, through angular clasts to a crushed powdered matrix. Some of these characteristics will be obscured if the material is

transported supraglacially for years to decades and then redeposited into a moraine. However, we expect that the following characteristics of rock avalanches will be preserved in moraines sourced from them. 1. The moraine will be significantly more monolithological than one produced by a climate-driven advance, because a rock avalanche source area generally has a single-lithology. If more than one lithology is involved in the rock avalanche, the separate rock types will remain un-mixed. 2. Moraine material will be significantly more angular, because a much smaller proportion of its constituent material will have been cycled through the englacial and subglacial regions. 3. The moraine may contain more fine material, because rock avalanche grain-size distributions are dominated by very fine dust (McSaveney and Davies, 2007). However, this depends on the glacier substrate; alpine lodgement tills derived from shale and mudstone substrates have a clayey matrix. To this we add that, although not of itself diagnostic, the presence of a dominant supraglacial component would be expected for rock-avalanche-derived moraines.

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Examination of sections through terminal and latero-terminal moraines along the West Coast of South Island reveals that many of these moraines show two specific lithofacies that differ significantly in their sedimentary architecture (Fig. 6). Parts of the moraines are composed of clinoform-type beds that dip at a moderate to steep angle away from the former glacier (Fig. 6; left side of sketch). These deposits are ice-contact latero-frontal fans and ramps (sensu Owen and Derbyshire, 1989; Benn and Evans, 1998) that formed at the glacier margin by an avalanching process where the glacier bulldozes and recycles subglacial and icemarginal materials. The resulting deposits are crudely to well stratified, dominated by sub-angular to sub-rounded material (of subglacial origin) and contain the full range of lithologies that are found in the lower reaches of the glacier. These relatively orderly deposits contrast with overlying massive, chaotic, clast-supported, bouldery gravels which are markedly more angular and typically mono-lithological (Fig. 6, right). The large clast size, low matrix content and high angularity indicate that these deposits are supraglacial in origin. The two different sedimentary components of these moraines are consistent with modified glacio-fluvially emplaced fan-head and supraglacial dump components of a terminal moraine. If the moraine is rock-avalanche derived the diagnostic differences will be evident primarily in the supraglacial component, but we would expect rounding and lithology characteristics to be transmitted in part to the fan-head system via crevasses and moulins.

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We also note that although a climatically induced advance may generate larger moraines due to its much greater duration, there is no aspect of climatic forcing that is likely to disproportionately increase the supraglacial component in a valley relative to that cycled through the subglacial system. Indeed, climatically driven thickening of the ice in a valley sequence will reduce the relief above the surface of the glacier and is thus likely to reduce the relative importance of the supraglacial debris.

5.1. Lithology We have recently demonstrated that the well-known Waiho Loop moraine of the Franz Josef Glacier is the product of a large (>108 m3) and probably co-seismic landslide onto the glacier (Tovar et al., 2008). The diagnostic criterion in this case is lithology. The regional geology is dominated by the Alpine Fault, and both the mountain ranges and lithological boundaries trend parallel to its strike (Fig. 4). Glacial valleys aligned at angles to the range front cross discrete bands of different rocks. Low grade meta-sedimentary (greywacke) rocks dominant in the Waiho Loop have their most northerly outcrop in the Waiho Valley 13 km up-valley of the Loop. Both the smaller Holocene moraines up-valley from the Loop and the larger Late Glacial Maximum moraines downvalley from the Loop (Figs. 4 and 5) are dominated by high-grade schists that crop out in the Waiho Valley north of the greywacke limit. Only

Fig 6. Gillespie’s Beach moraine section. Note the contrast between stratified (brown) poly-lithological ice-contact fans at the left and massive (grey) bouldery deposits on the right. The former were formed as part of a fan-head margin, whereas the latter are inferred to be supraglacially dumped debris of rock-avalanche origin based on the larger size and mono-lithological nature of the clasts. Sample B1 from the supraglacial material is 100% schist, whereas sample B2 contains a mix of lithologies (n ¼ 50).

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supraglacial transport could carry a large volume of greywacke debris in a relatively unmixed state through the schist-rich lower valley. The fact that the rock avalanche that created the Waiho Loop occurred sufficiently far up-valley to give an unequivocal lithologic signature is fortuitous, but many rock avalanches will also occur in the lower valleys where schists crop out. Glaciers will be thinner in the lower reaches of the valleys, and these valley walls are more likely to be exposed above the glacier surface and available for sediment delivery than lithologies farther up-valley. The fan-head deposits in Figs. 5 and 6 contain rocks of all the lithologies the glacier has passed through. In the case of smaller gravity-driven contributions to the supraglacial load, such as rockfalls, small rock avalanches and debris flows, the supraglacial component should reflect some admixture of all the rock types that have freeboard in the glacial valleys. We expect a higher proportion of this material to be cycled through the subglacial drainage system because it is less voluminous than rock avalanche debris. Although ice-flow rates on the modern west-coast glaciers are among the highest in the world (e.g. Oerlemans, 1997), it still takes one to several decades, depending on the position of the glacier, for ice from a ne´ve´ to reach its terminus (Larsen et al., 2005). There should thus be considerable admixture of lithologies, and this is what we observe. Examination of the upper unit of supraglacial component of Fig. 6 shows that 100% of the material recovered was schist. This, combined with the voluminous nature of the deposit and the very large and angular boulders incorporated, makes this feature a likely rock-avalanche sourced moraine. Being part of a LGM moraine sequence it was presumably part of a climatically driven advance.

5.2. Particle shape A perhumid environment like New Zealand’s west coast is an ideal one to distinguish supraglacial debris from subglacial material because the subglacial fluvial environment is conducive to rounding rocks. The glacio-fluvially emplaced fan-head components of modern terminal moraines in the Waiho Valley are difficult to differentiate from gravel bar deposits in the adjacent rivers on the basis of their particle shape (Figs. 4 and 5). Consequently, it is relatively easy to distinguish supraglacial deposits from subglacially transported ones, and we have highlighted examples of supraglacial material in Fig. 4 and subglacial material in Fig. 5. Clast shape does not of itself identify a rock avalanche origin but it is a useful filter.

5.4. Morphology We have previously noted that end moraines derived from rockavalanche deposits contain much more dumped material than more typical end moraines. Supraglacial dumping creates steepsided moraines. Large capping boulders are common. A large steepsided moraine is not of itself indicative of a rock avalanche source, but an anomalously large and prominent moraine in a sequence of low-angle subdued moraines signifies a large supraglacial load. The moraine shown in Fig. 6 was derived from a glacier that largely filled the lower valley reaches of the Fox River; the potential for rockslides would be lower than in the modern valley because freeboard was reduced. The simple existence of a moraine with voluminous supraglacial load in this setting is strong circumstantial support for a rock avalanche origin.

6. Summary We view large landslides as an important link between tectonics and glacier behaviour along the West Coast of the South Island of New Zealand. We have outcrop evidence at several locations indicating that large volumes of angular single-lithology sediment have been transported supraglacially to glacier termini, forming terminal moraines. These mono-lithological deposits contrast with the ‘normal’ terminal moraines which are better stratified and comprise debris that is more rounded and lithologically more diverse. In the case of the Waiho Loop, we have prima facie evidence that the materials were derived from a large landslide far up the Waiho Valley in Torlesse greywacke. Even where the dominant lithology is locally sourced schist, supraglacial transport of large volumes of mono-lithological material to the terminus is consistent with a landslide origin. We note that within a short period from the recognition of the first rock avalanche deposit in the Karakorum (Hewitt, 1988), the number of recognised rock avalanche deposits in that mountain range has increased exponentially (Hewitt, 1999, 2009). We expect that rock-avalanche sourced terminal moraines will likewise be increasingly recognised. The implication of the preservation of non-climatic moraines in this setting is important. It precludes straightforward inference of climate forcing (Anderson and Mackintosh, 2006; Rother and Shulmeister, 2006) from moraines and may help explain apparently anomalous ages from some of the moraines that have been dated (Barrows et al., 2007).

7. Conclusions 5.3. Grain-size Of possible diagnostic characteristics of rock avalanches, the most promising is perhaps grain-size. Debris of large rock avalanches is rich in fine matrix and is different from matrix-free debris of smaller rockfalls and from the coarser matrices of other colluvial materials; indeed most of the grains in rock avalanche debris are (sub)micron in size (McSaveney and Davies, 2007). These fines may be preserved in supraglacial transport deep within the deposit, but we note that infiltrating and percolating rainwater will slowly remove these fines from the debris apron (Hewitt, 2009). The absence or lower-than-expected content of these fines in the upper debris layers is not indicative of the absence of rock avalanche debris. Furthermore, the process of terminal moraine formation undoubtedly leads to the wholesale loss of very fine angular grains through wash-out and deflation. However, we think some may be present in rock avalanche-derived end moraines, and we are presently investigating this possibility.

1. Rock avalanches are the source of much of the debris that forms glacial moraines in active mountains. 2. Debris from large rock avalanches deposited on ablation zones can affect glacier mass balance significantly, with corresponding implications for glacier behaviour and moraine emplacement. 3. There is theoretical and empirical evidence that a sufficiently large rock avalanche can cause a glacier to advance substantially and rapidly. 4. An expected consequence of such an advance is a terminal moraine. Such moraines are expected display steeper-sided profiles, and to contain more angular and monolithologic debris, than climate-driven terminal moraines because the former have a larger supraglacial component. 5. Landslide-generated terminal moraines have no necessary paleoclimatic significance. This is the case with the Waiho Loop moraine in New Zealand, which has previously been cited as evidence for a Southern Hemisphere climate response to Northern Hemisphere forcing at the end of the last ice age.

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