Stability of Natural Slopes - Hungr

8/12/2019 Stability of Natural Slopes - Hungr

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STABILITY AND FAILURE BEHAVIOUR OF NATURAL SLOPES

Oldrich Hungr

Department of Earth and Ocean Sciences, University of British Columbia,

6339 Stores Road, Vancouver, B.C., V6T 1Z4, Canada

INTRODUCTION

Considerable advances in knowledge related to the stability of natural slopes have been made inrecent decades (Morgenstern, 1992). But engineering concerns with natural slopes do not endwith this issue. In case of natural slopes, the question of whether the slope could fail or not isoften secondary to that concerning the probable character and consequences of a failure(Morgenstern, 1978). What methods are available for predicting the consequences of slopefailure?

The term failure requires a definition. As reviewed by Skempton and Hutchinson (1969), thehistory of a mass movement comprises pre-failure deformations, failure itself and post-failuredisplacements. Many slow-moving landslides exhibit a large number of movement episodes,separated by long or short periods of relative quiescence. The following definition is proposedfor the purposes of this paper:

Failure is the single most significant movement episode in the known or anticipated history ofa landslide and also one which involves the formation of a continuous rupture surface as adisplacement or strain discontinuity.

Landslide engineering lacks a unified scale of destructiveness, such as exists in the Mercalliscale of earthquake intensity (Morgenstern, 1985). Intensity is a spatial function, defining thedestructive power of an event in terms of certain parameters. Hungr (1981, p.23) argued thatvelocity is the most important intensity measure. Other intensity parameters may include depthof movement, thickness of deposits or strain (Hungr, 1987). Velocity can be related to thetypical human response to landslide occurrence, as shown in Table 1 (see also Cruden andVarnes, 1996). What is most important is whether failure will lead to very rapid or extremevelocities (about in excess of 1 m/s), which pose direct risk to human life. It must be noted thatslower landslides may also cause substantial damage, or threaten human life through secondaryeffects such as flooding or structural damage. However, in this discussion we concentrate onprimary impact.

The second landslide dimension relevant to destructiveness is the mass of the slide, which isroughly proportional to volume (magnitude) and the area affected (e.g. Hungr, 1990). Theinitial volume of a landslide is the volume contained between the rupture surface and the pre-failure ground surface. This is referred to here as the source volume. The areal regionaffected by a landslide consists of the source area i.e. the plan area of the rupture surface, thepath along which the slide masses move after leaving the source and the deposition area


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Table 1Landslide velocity scale (IUGS,WGL , 1995, Cruden and Varnes, 1996, see also Hungr, 1981,).

Class Description Velocity(m/s)

Typicalvelocity

Human response

7 Extremely Rapid Escape impossible---------------------- 5 5 m/sec6 Very Rapid Escape possible, but uncertain

---------------------- 5x10-2 3 m/min5 Rapid Escape, but limited evacuation

---------------------- 5x10-4 1.8 m/hr4 Moderate Evacuation

---------------------- 5x10-6 13 m/month3 Slow Removal or maintenance

---------------------- 5x10-8 1.6 m/year2 Very Slow Maintenance of structures

---------------------- 5x10-10

16 mm/year1 Extremely Slow No serious effects

where they come to rest. (This terminology was first suggested to the writer in an unpublishedletter by J.King of the GEO, Hong Kong).

In case of mobile landslides, there may be a considerable distance between the source and thedeposit. Further, the slide volume, as well as the character of its material, may change during itsmotion due to comminution, entrainment of solid material and incorporation or drainage ofwater. Many mobile landslides are strongly enlarging. Others behave retrogressively (e.g.

Mitchell and Markell, 1974). Landslides which exhibit extremely rapid motion and significantimpacts in areas outside the source area could be called catastrophic landslides, to recognizetheir high hazard potential.

This brief discussion indicates that slope stability analysis is only the first step towards hazardassessment of catastrophic landslides. The second step, given that a failure is possible, is topredict whether of not catastrophic motion could develop and what volume of material willultimately be involved. The third step, not covered here, is the process of mapping the value anddistribution of intensity parameters over the extent of the hazard area (runout analysis). In asummary, slope hazard assessment for potentially catastrophic landslides involves the followingsteps:

1) Slope stability analysis2) Assessment of failure behaviour3) Intensity mapping (runout analysis)

This article deals with Step 2, assessment of failure behaviour. It reviews several differentscenaria, in which unique site characteristics determined the failure behaviour of large landslides


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in soil and rock. The review is not exhaustive, as undoubtedly many other complex mechanismscapable of producing catastrophic failures exist.

EXAMPLES

Rock structure and mobility of rock slides and slumps

In the absence of favourably oriented discontinuities, rock fails by rotational shearing (slumping)as if it were a cohesive soil. However, only the weakest rock masses can fail by shearingthrough intact rock material: theoretically, a 100 m high vertical cliff would fail only if theaverage cohesion of the rock mass was less than about 0.7 MPa. This is a very low value forintact rock. Strength reduction due to random discontinuities, fissures and other defects mustalso play a role.

Rock slump occurs in weak rocks such as shale, marl or tuff, which are isotropic, or structured sothat the weak direction is horizontal or dips opposite to the slope direction. The failure

mechanism is inherently self-stabilizing, as the rotation of the sliding body brings about areduction on the driving forces. This is especially so if the failure surface passes beneath theslope toe or if the slope angle decreases in the toe region

Slump failures tend to be slow-moving, although they may involve large displacements. Anexcellent example is the slumping of several millions m3 of Cretaceous calcareous shales(Valanginian), capped by massive limestone (Urgonian) in the Massif de Plat near Chamonix inthe Savoy Alps, France (Figure 1). The moderately folded Alpine Foreland sedimentarysequence is roughly horizontal at this location. An 18th. century eyewitness description of thefailure scenario was quoted by Eisbacher and Clague (1984, p. 142):

I found myself in the face of a mountain completely enveloped in smoke, from which brokecontinuously, day and night, masses of rock, with an astounding noise stronger than thunder orthe battery of a cannon.

This activity continued for over a month (Eisbacher and Clague, 1984). Only localized slide-head toppling of blocks in the Urgonian limestone cap was of a catastrophic nature, while theoverall failure was slow and gradual. The overall displacement of the large rock slide barely

Figure 1

A profile of a typical rotational slumpin weak calcareous shales overlain bymassive limestone, similar to thefailure at the Massif de Platdescribed in the text. (based on anunpublished sketch by Dr. A.Malatrait, B.R.G.M., Lyon, France).

100 m


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reached the toe of the steep valley slope.No landslide dam formed in the channel ofthe Arve River and no effects were felt onthe opposite bank (Figure 2).

An entirely different type of failureoccurred in 1281 in the same rock sequencesome 50 km away, at Mt. Granier.According to an interpretation due toCruden and Antoine (1984), sliding on marlbedding planes in the Valanginian, inclinedat only 17 towards the valley, took placeunder a mass up to 600 m thick (Figure 3).The resulting block of several hundreds ofmillions m3 moved towards the valley,disintegrated and covered an area of more

than 15 km

2

by rock avalanche debris in amatter of minutes. A provincial town wasobliterated, with a loss of possibly as manyas 5,000 persons. This catastrophic slideoccurred in a sequence of Cretaceous strata

identical to that which produced the moderately rapid Massif de Plat slump, mentioned earlier,and in a slope of similar initial geometry. The sole factor responsible for a dramatically differentbehaviour was the bedding orientation, with a modest dip out of the slope in the former case andinto the slope in the latter.

What is the reason for such diverse failure behaviour in two nearly identical slopes? The answermust lie in the way weak rock masses fail in shear. It appears that once weak, randomly jointedor fissured rock mass becomes overstressed, it deforms plastically, without a sudden drop ofstrength. Triaxial testing of rock samples at high confining stresses (relative to the compressivestrength) shows similar ductile behaviour (e.g. Byerlee, 1968). It should theoretically bepossible to define an upper limit of rock strength, below which ductile rock mass failure canoccur in a given slope. The strength should be normalized to the stress level using the StabilityNumber:

N H

s

c

=

2

Eqn. [1]

Here, ? is the unit weight of the rock material, H is the height of the slope, expressed as theelevation difference between the crest and toe of the rupture surface and s c is the uniaxialcompressive strength of the rock. For a typical large catastrophic rock slide involving strongrock and a 300 m high slope, Nsequals 2x26x300/100000, or 0.15. For a small extremely rapidcollapse of a 40 m high portion of a cliff in soft Chalk, with a s c of 5 MPa, Ns equals2x22x40/5000, or 0.35. For a large slow-moving rock slump in very weak shale or marl, with anH of 200 m and a s c of 2 MPa, it is 2x26x200/2000, or 5.2. It appears that Ns should be

Figure 2

A geomorphological map of the Masiff dePlat slope ( after Goguel and Pachoud,1981). The width of the map is 3.2 km.


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considerably above 0.5 to allow the occurrence ofa slow ductile rock slump. However, this is asubject for future research.

The brittle behaviour of the structurally-controlled

translational failure at Mt Granier, on the otherhand, can be explained by the undrainedbrittleness of rock joints, sheared through peak toresidual strength. Given the great thickness of thefailed mass, rate-softenning behaviour due tofrictional heating along the controlling beddingjoints is possible (Goguel and Pachoud, 1972).

Catastrophic, structurally-controlled translationalslide movements take place most frequently incarbonates and less often in other sediments. The

largest rock avalanches of the historical record(e.g. Goldau, 1806), as well as the largestcatastrophic non-volcanic landslide on Earth, theSeimareh Slide in Iran (Shoaei and Ghayoumian,1998) are of this type.

2.2 Rock structure and the failure behaviour of rock topples

Many slopes formed in steeply dipping strength-anisotropic rocks exhibit outward rotation andreverse shearing of the rock mass, described as flexural toppling (Goodman and Bray, 1976).The original rock fabric often dips into the slope, but steep cataclinal slopes have also beenaffected (e.g. Cruden, 1989). The rock type is predominantly schist, phyllite or slate, lesscommonly closely jointed gneiss or sedimentary rocks. A well known example of flexuraltoppling is the Clapire slide in Southern France, described by Follaci (1987) and shown inFigure 4.

Figure 5 shows flexural toppling produced analytically using the Universal Discrete ElementModel (Itasca, 1993, Nichol, 2000, Nichol et al., 2001). The model shows all the characteristicmorphological features typical of this type of slope movement: gradual forward rotation of thebeds, a hinge zone, multiple shallow tension cracks and antislope scarps. The major principalstress in the model is parallel with the slope surface, except in a relatively thin de-stressedsurficial zone which contains tension cracks. The model stabilized itself at a joint inclinationequal to that predicted by the stability criterion proposed by Goodman and Bray (1976). In orderto produce this behaviour in the model, the strength of the rock comprising the steeply-dippinglayers has to be low and the rock mass quality poor on account of close jointing. This confirmsthe general observation that flexural toppling is limited to weak rocks.

The flexural toppling mechanism is self-stabilizing, because the shear stresses on the weaksurfaces decrease as the beds tilt forward. This, in combination with the general low strength of

Figure 3

An isometric sketch of the Mt. Granierblock slide (based on an interpretation byCruden and Antoine, 1987).


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Figure 4

La Clapire slide. Note multiple reversescarps caused by flexural toppling.

Figure 6

A view of the source area of the MysteryCreek rock avalanche in the middle ground(Nichol et al., 2001).

Figure 5

An analytical model of a flexural topple.

The width of the pictured region is 800 m(Nichol et al., 2001).

Figure 7

An analytical model of a catastrophic blocktopple. The width of the pictured region is800m (Nichol et al., 2001).


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the rock mass involved, produces ductile behaviour similar to that of rock slump failures. Highlydeveloped flexural toppling can transform into a slump, as the strata break in a curved hingezone and begin to slide. This is apparent in the crest area of the Clapire slide. There appear tobe few examples of such a mechanism leading to catastrophic velocities, while examples ofductile behaviour are abundant (e.g. DeFreitas and Watters, 1973).

The second type of toppling mechanism defined by Goodman and Bray (1976) is block toppling.This is a rotational failure of steeply dipping beds in relatively massive rock, where stability isderived mostly from the base of the rock columns, rather than from friction along their sides.Simple catastrophic toppling of a large isolated limestone tower (the "Pulverhorndl") wasdescribed from the Austrian Alps by Terzaghi (1950). Such detachments are limited in volume.

Of greater concern is catastrophic block toppling of series of massive columns separated bysteeply dipping joints. An example of this kind was observed by the writer in a railway cut insteeply jointed massive quartzite, involving some 10,000 m3. The 1928 Motto d'Arbino failure inSwitzerland, involving 30 to 40 million m3 of gneiss and marble, may have had a similar

mechanism (Heim, 1932). This failure was a small part at the front of a large moving mass.Heim's description of the 1928 event and the limited runout of the deposits, indicate that thefailure occurred in a piecemeal manner, with a succession of smaller falls continuing over manyhours.

Under certain circumstances, block toppling can produce large-scale catastrophic slope failures,followed by rock avalanches. The pre-historic Mystery Creek rock avalanche near Whistler,British Columbia, Canada, involved 40 million m3of quartz diorite, containing a well developedsub-vertical joint set with a modal spacing of 2 to 5m, as well as more widely spaced, but highlypersistent joints dipping at a shallow angle in the slope direction. The failed slope was traversedobliquely by a vertical cliff over 100m high and this probably provided a kinematic release forlarge-scale block toppling. The source scar exhibits a sloping base covered by partially rotatedlarge blocks and is surrounded by tension cracks and anti-slope scarps (Figure 6). The resultingrock avalanche travelled 2 km across the valley floor.

An analytical model of the Mystery Creek case is shown in Figure 7 (Nichol et al., 2001). Thismodel uses rock with high strength. Because this strength inhibits bending of the strata, a set ofdiscontinuities dipping down slope must exist, which act as a base for rotation. Failure begins bya forward rotation of three tiers of blocks, forming a kink band. The rotated blocks lining thebase of the rupture surface in Figure 6 probably represent the lower part of this kink band. Acatastrophic general failure and disintegration of the source mass follows. The major principalstress during the initial stages of failure is oriented parallel with the near-vertical joint set. Thisis a different stress state from that observed in the flexural toppling model (Figure 5) and atotally different failure behaviour results from it.

Thus, recognition of catastrophic failure potential in toppling slopes requires careful search forstructures dipping in the direction of movement. It is uncertain, however, whether strongearthquake shaking could generate rupture within the toppling hinge, intensive enough to triggerbrittle detachment. Chigira and Kiho (1994) describe five examples of toppling slopes inJapanese Cretaceous sediments, which developed into highly mobile rock avalanches. It is


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unclear whether downslope-orientedstructures existed in these cases.However, two of the landslides occurredas a result of earthquakes.

2.3 The role of internal strength in acompound slide in overconsolidated

clay

Slides involving overconsolidated, non-sensitive clay tend to be slow and ductile(e.g. Skempton and Hutchinson, 1969).Many such slides are found along rivervalleys of the Canadian Prairies,exploiting weak surfaces in heavilyoverconsolidated glacio-lacustrine clays,

or the underlying argillaceous bedrock(Thomson and Morgenstern, 1977).

An unusual case is the Attachie Slide,which occurred in May, 1973 on the PeaceRiver upstream of Fort St.John, north-eastern British Columbia. The south sideof the 200 m deep valley had beenunstable for at least 30 years, before thefirst airphotos of the area were taken.Sliding was taking place in a thick andcomplex sequence of glacio-lacustrineclays, clayey silts and silts, covered byabout 30 m of glacial till. The glacio-lacustrine sequence was underlain by basalgravels and Cretaceous bedrock, neither ofwhich participated in the instability.Displacements totaling several tens ofmetres had occurred at several levels inthe slope by 1973, creating several largescarps (Figure 8a). The instabilityoccurred in two stages, upper and lower,separated by an intact intermediate scarpwhich can be seen on both in Figures 8aand b and on the profile, Figure 9.

Precedent from similar locations in prairievalleys would indicate that a slow, ductilefailure was in progress. However, in May,

a

b

a

b

Figure 8

Vertical airphotos of the Attachie Slide.a) in 1970 (BC5529,75) b) in 1973following the rapid flow slide of May,1973 (BC7279,70). The area covered bythe photographs is approximately 1,800m wide.


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0 10050 150 250200 300

Unit 2 Pre-glacial laminated clay and silt

Unit 4 Till

Unit 5 Colluvium

Unit 1 Cobble/gravel

Unit 3 Sand

?

550

600450350 400 500 550 750650 700

400

500

450

Post-failure surface

Pre-failure surface

650

600

meters

Unit 6 Post-glacial laminated clay and silt

Upper stage

Lower stage

Figure 9

A central cross-section of the Attachie Slide (Fletcher, 2000)

1973, following a relatively wet year, the lower stage of the slope became suddenly mobile. Atotal of 12.4 million m3of material moved on both levels. About half of this volume (6.4 millionm3) was sufficiently mobile to descend a 60 m scarp at the toe of the slope and move rapidlyacross the floodplain of the Peace River. This mobile portion corresponded largely with thelower stage of the instability, between the toe and the intermediate scarp. The upper stage alsodisplaced, but with much less mobility. The flow slide had sufficient momentum to raise a wave

and impact the opposite shore. Details of the case history can be found in Evans et al. (1996),Fletcher (2000) and Fletcher et al. (2001).

The glacio-lacustrine soil, forming the most mobile part of the displacement material, consistedof approximately 31% of low-plastic silt, 48% plastic clayey silt and 21% sand (Fletcher et al.,2001). All these materials are complexly interbedded and very stiff or dense, having beenoverridden by glacial ice on at least one occasion. There is some evidence that the low plasticitysilt units are cemented by calcium carbonate or gypsum. The flat basal part of the rupturesurface formed in thinly laminated highly plastic illite clay with a clay fraction of 60-70%. Theresidual friction angles ranged from 17 to 25, on account of varying silt content. Threeconsolidated undrained triaxial tests on this material indicated dilative behaviour with values of

the pore-pressure parameter, A at failure of 0,12 to 0.55, indicating low sensitivity. Liquidityindices generally below 0.5 also pointed to low sensitivity.

It is not difficult to explain the occurrence of instability at this site, given the existence of pre-sheared layers of plastic clay and the presence of confined ground water aquifers. However, it ismuch more difficult to explain the occurrence of extremely rapid, catastrophic failure asobserved in May, 1973.


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In order to quantify the strength loss evidenced by the 1973 displacement, a dynamic analysis ofthe lower stage source volume was carried out with the program DAN (Hungr, 1995), usingfrictional rheology. Used in a back-calculation, the model can give the strength loss required toobtain the geometry change observed during the failure. In this case, it was first assumed thatundrained pore-pressure increase was the sole source of brittleness. The analysis showed that an

average pore pressure ratio increase from 0.26 to 0.50 was required to deform the sliding bodyas observed in Figure 9. The corresponding change in the Factor of Safety is approximatelyfrom 1.0 to 0.5 (Fletcher, 2000, Fletcher et al., 2001).

It is difficult to explain such a dramatic, instant strength loss, given the character of soil materialsat this site. None of the soil descriptions give an indication of the possibility of contractivebehaviour, which could translate into high undrained brittleness. Further, the large ductile pre-1973 displacements of the slope are hard to reconcile with the behaviour of a sensitive clay.Alternative explanations of the catastrophic failure are therefore needed. After a detailedexamination of several alternatives, Fletcher (2000) concluded that two failure mechanisms arepossible:

The first hypothesis considered the internal strength of the sliding body. The profile of thelower stage of the Attachie slide is compound. The rupture surface consists of a long, flat-lyingsegment which follows the pre-sheared clay layer and a steep back scarp cutting across layers.In order to fail, such a geometry requires strong internal distortion on secondary shearscomplementary to the back scarp (Hutchinson, 1988). The high internal strength of the unfailedvery stiff, possibly cemented soil, adds considerably to the overall sliding resistance, but reducessharply, once the brittle material fails. The quantitative effect can be modeled using a two-blockstability analysis (Fletcher, 2000).

The second hypothesis concentrates on the occurrence of multiple tension cracks in the disturbedslope, as a result of pre-1973 movements. The cracks divide the silty soil into a network of stiffblocks separated by discontinuities. Many of the cracks fill over time by loose silty debris,which may become saturated by surface water. Once such mixture of intact blocks and loosematrix is forced to move, localized liquefaction may occur, driving the toe portion of the slopeforward. This effect can be referred to as macroscopic brittleness. Analyses by Fletcher(2000) show that most likely both of the above mechanisms participated in the spectacular failureof May, 1973. Neither mechanism is normally considered in hazard analyses for slopes inoverconsolidated clays and silts.

2.4 The failure mechanism of a submarine sand flow slide

The main channel of the Fraser River enters the Pacific Ocean at a location on the front of theFraser delta called Sandheads, 15 km south of the City of Vancouver. In the past, the riverchannel meandered over the surface of the delta, and its mouth migrated laterally across the deltafront (Figure 10). Wherever the channel crossed the delta crest, a submarine canyon formed. In1932, a permanent jetty was constructed, confining the river outlet permanently to a location justsouth of the Sandheads lighthouse. A major canyon developed in front of this location duringthe 7 decades following the completion of the jetty.


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At present, Public Works Canada (PWC) conducts annual dredging of the main channelimmediately to the south of the jetty, to maintain a minimum depth of 10 m, required for

navigation. Submarine landslides occur periodically at the western end of the navigationchannel, just south of the jetty (Mc Kenna et al., 1992). The most significant of the recordedlandslide events occurred on June 30, 1985. A PWC survey crew was engaged in sounding thedepth of the shipping channel. They found that the depth of the channel bottom changed from 10to 30 metres between morning and afternoon, over an area nearly 300 m long along the river axisand 250 m wide. Thus, more than 1 million m3of sediment vacated the river mouth in course ofseveral hours and flowed west over the face of the delta and into the submarine canyon (Figure11). Neither the jetty nor the Sandheads Lighthouse were affected by the adjacent landslide. No

Figure 10

Location plan of the Fraser delta front.The square frame is 4 km wide . Notesubmarine canyons and historic channels

of the Fraser River

Figure 11

Position of the 10 m depthcontour preceding (i) andfollowing (f) each of thedocumented submarinelandslide events (fromMcKenna et al., 1992). Thedashed line on the left andthe full line on the right side

of the picture represent the1985 landslide.


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0 1 10 100 1000 10000

Rate of sedimentation (cm/ year)

0

20

40

60

80

100

Consolidation

degree(%)

Cv(cm2/ sec)

0.001 (silt)

0.01 (sand)

0.1 (sand)

Sandheads

Figure 12

A nomogram derived from the theory of coupled consolidation and sedimentation, following theapproach of Morgenstern (1967).

unusual waves, or even turbidity were reported in the shipping channel on that day. The spaceemptied by the 1985 landslide contained fresh river sediment, which accumulated during riverfreshets since the last preceding failure (Figure 11). Since 1985, the same space infillled againby new sediment. Thus, the landslides appear to periodically empty the same approximatelocation, to be filled anew by river sedimentation.

This periodic cycle of landsliding andsedimentation suggests a possible explanation forthe occurrence of the large failures. The Fraserdelivers large quantities of fine sand to the deltafront during the freshet (late June-July). Based onsounding data, as much as 10 m of verticalaggradation can occur at the slope crest per year.Considering that much of this occurs during lessthan two months, the aggradation rate can be ashigh as 5 m per month. Morgenstern (1967)considered the effects of underconsolidation on thestability of rapidly aggrading delta slopes.Following his approach, Figure 12 shows anomogram of underconsolidation, derived using atheory of coupled sedimentation and consolidation.Given the high sedimentation rate of fine sand atSandheads, it is quite possible for the degree ofconsolidation to remain at only 50% during thefreshet period. This may create excess porepressures, capable of triggering local instability.

a

b

c

Figure 13

Schematic view of a stage in theformation of a retrogressive flow slide.

Retrogression, 4 m

Liquefaction


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A typical longitudinal profile along the centerline of the river channel at the crest of the deltaslope is shown in Figure 13a. A limit equilibrium analysis showed that small-scale slumpingfailure can easily occur at the crest, prompted by underconsolidation and possibly also byspontaneous undrained collapse of the loose sediment. Once an initial failure occurs, the

liquefied sand will flow down the slope as shown in Figure 13b. The wave of sand will over-rideloose material on lower slopes. An approximate undrained limit equilibrium analysis of the frontof the propagating sand wave proved that entrainment of a large quantity of sand from the pathof the sand wave is likely (cf. Sassa. 1988).

In this manner, a relatively shallow flow slide can sweep over the slope, removing a slice ofmaterial as shown in Figure 13c. The crest of the slope is left unsupported and may failretrogressively by the same mechanism. In this way, a more-less continuous flow of liquefiedsand may be set up on the slope, moving material steadily towards the submarine canyon andenlarging the landslide crater in the upstream direction. Such a process would explain how anenormous quantity of material can be removed in a few hours, without causing a noticeable

disturbance of the water surface. The material entrainment process may continue on lowerslopes of the delta, deepening the submarine canyon and possibly diluting itself to form aturbidity current.

3. CONCLUSION

The need for predicting landslide failure behaviour and consequences was expressed byN.R. Morgenstern in the following words (Morgenstern, 1978):

There is a class of problems that arises in practice that is not concerned with the evaluationof whether a slope will fail or not; but instead obliges us to assume movement and design againstthe consequences. Where it is not practical to eliminate movements, the following questionsarise:

1) How much material will move?2) What will be the time history of the movements in terms of velocities and accelerations?3) How are protective structures designed against moving masses?

Regrettably, we have not made a very major progress towards answering these quantitativequestions during the 23 years since they were posed. As shown in the four examples presentedhere, it is not an easy task and certainly not one that is amenable to standardized solutions.Prediction of failure behaviour requires understanding of both geology and mechanics, combinedwith a large measure of intuition and even imagination.


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REFERENCES

Byerlee, J.W., 1968. Brittle-ductile transition in rocks. Journal of Geophysical Research,73:4741-4750.

Chigira, M. and Kiho, K., 1994. Deep-seated rockslide-avalanches preceded by mass rock creepof sedimentary rocks in the Akaishi Mountains, central Japan. Engineering Geology 38:221-230.

Cruden, D.M., 1989. Limits to common toppling. Canadian Geotechnical Journal, 26, 737-742.

Cruden, D.M. and Antoine, P., 1984. The slide from Mt. Granier, Isre and Savoie, France onNov. 24, 1248. Procs., 4th. International Symposium on Landslides, Toronto, 1, 475-481.

Cruden, D.M. and Varnes, D.J., 1996. Landslide Types and Processes. In LandslidesInvestigation and Mitigation. Transportation Research Board, US National Research Council,Turner, A.K. and Schuster, R.L. (editors). Special Report 247, Washington, DC 1996, Chapter

3, 36-75.DeFreitas, M.H., Watters, R.J., 1973. Some examples of toppling failure. Geotechnique, 23,495-514.

Eisbacher, G.H. and Clague, J.J., 1984. Destructive mass movements in high mountains: hazardand management. Geological Survey of Canada, Paper 84-16, 230 pp.

Evans, S., Hu, X.Q., Enegren, E.G. 1996. The 1973 Attachie Slide, Peace River Valley, near FortSt. John, B.C., Canada: a landslide with a high-velocity flow slide component in Pleistocenesediments. In Proceedings, 7th International Symposium on Landslides, Trondheim, Norway,715-720

Fletcher, L., 2000. Failure behaviour of two large landslides in silt and clay. M.A.Sc. Thesis,Geological Engineering, University of British Columbia, 170p.

Fletcher, L., Hungr, O. and Evans, S.G., 2001. Contrasting failure behaviour of two largelandslides in clay and silt. Canadian Geotechnical Journal (in review).

Follacci, J.P., 1987. Les mouvements du versant de la Clapire a Saint tienne de Tine (AlpesMaritimes). Bull. Liaison Labo. Ponts et Chauses, Paris, 151, 39-54.

Goguel, J., Pachoud, A., 1972. Geology and dynamics of the rockfall of the Granier Rangewhich occurred in November 1248. Bulletin, Bureau de Rcherches Geologiques et Minires,Hydrogeologie, Lyon, 1, 29-38.

Goguel, J., Pachoud, A., 1981. Les mouvements de terrain du versant sud du Massif de Plat,Haute Savoie, France. Bull. Liaison Laboratoire des Pontes et Chauses, Spcial X:15-25.


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Goodman, R.E., Bray, J.W., 1976. Toppling of rock slopes. Rock Engineering, ASCE,Geotechnical Eng. Division Conference, Boulder, pp. 2, 201-237.

Heim, A. (1932): Landslides and human lives (Bergsturz und Menchenleben)., N. Skermer,Translator, Bi-Tech Publishers, Vancouver, B.C., 196 pp.

Hungr, O., 1981. Dynamics of rock avalanches and other types of slope movements. Ph.D.Thesis, University of Alberta, Edmonton.

Hungr, O., 1997. Some methods of landslide hazard intensity mapping. Invited paper, Procs.,Landslide Risk Workshop, R.Fell and D.M. Cruden, Eds., Balkema, Rotterdam , pp. 215-226.

Hungr, O., 1990. Mobility of rock avalanches. Reports of the National Research Institute forEarth Science and Disaster Prevention, Tsukuba, Japan, 46: 11-20.

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