Columbia Mountain Landslide Late-glacial Emplacement And

.Geomorphology 41 2001 309322www.elsevier.comrlocatergeomorph

Columbia Mountain landslide: late-glacial emplacement andindications of future failure, Northwestern Montana, USA

Larry N. Smith)Montana Bureau of Mines and Geology, Montana Tech of The Uniersity of Montana, Butte, MT 59701-8997, USA

Received 7 November 2000; received in revised form 2 March 2001; accepted 9 March 2001

Abstract

The well preserved and undissected Columbia Mountain landslide, which is undergoing suburban development, wasstudied to estimate the timing and processes of emplacement. The landslide moved westward from a bedrock interfluve ofthe northern Swan Range in Montana, USA onto the deglaciated floor of the Flathead Valley. The landslide covers an areaof about 2 km2, has a toe-to-crown height of 1100 m, a total length of 3430 m, a thickness of between 3 and 75 m, and anapproximate volume of 40 million m3. Deposits and landforms define three portions of the landslide; from the toe to the

.head they are: i clast-rich diamictons made up of gravel-sized angular rock fragments with arcuate transverse ridges at the . .surface; ii silty and sandy deposits resting on diamictons in an internally drained depression behind the ridges; and iii

.diamictons containing angular and subangular pebble-to block-sized clasts some of which are glacially striated in an areaof lumpy topography between the depression and the head of the landslide. Drilling data suggest the diamictons coverblock-to-slab-sized bedrock clasts that resulted from an initial stage of the failure.

The landslide moved along a surface that developed at a high angle to the NE-dipping, thinly bedded metasediments ofthe Proterozoic Belt Supergroup. The exposed slope of the main scarp dips 30378W. A hypothetical initial rotational failureof the lower part of a bedrock interfluve may have transported bedrock clasts into the valley. The morphology and deposits

.at the surface of the landslide indicate deposition by a rock avalanche sturzstrom derived from a second stage of failurealong the upper part of the scarp.

The toe of the Columbia Mountain landslide is convex-west in planview, except where it was deflected around areas nowoccupied by glacial kettles on the north and south margins. Landsliding, therefore, occurred during deglaciation of the valleywhile ice still filled the present-day kettles. Available chronostratigraphy suggests that the ;1-km thick glacier in the regionmelted before 12,000 14C years BPwithin 3000 years of the last glacial maximum. Deglaciation and hillslope failure arelikely causally linked. Failure of the faceted interfluve was likely due tensile fracturing of bedrock along a bedding-normaljoint set shortly after glacial retreat from the hillslope.

Open surficial tension fractures and grabens in the Swan Range are limited to an area above the crown of the landslide. .Movement across these features suggests that extensional flow of bedrock sackung is occurring in what remains of the

ridge that failed in the Columbia Mountain landslide. The fractures and grabens likely were initiated during failure, but theirmorphologies suggest active extension across some grabens. Continued movement of bedrock above the crown may result infuture mass movements from above the previous landslide scarp. Landslides sourced from bedrock above the scarp of the

) Tel.: q1-406-496-4379; fax: q1-406-496-4343. .E-mail address: [email protected] L.N. Smith .

0169-555Xr01r$ - see front matter q 2001 Elsevier Science B.V. All rights reserved. .PII: S0169-555X 01 00062-9

( )L.N. SmithrGeomorphology 41 2001 309322310

late-glacial Columbia Mountain landslide, which could potentially be triggered by earthquakes, are geologic hazards in theregion. q 2001 Elsevier Science B.V. All rights reserved.

Keywords: Deglaciation; Landslide; Sackung; Geologic hazards; Northern rocky mountains

1. Introduction

Mass movement of bedrock and surficial sedi-ments immediately after retreat of valley glaciers is arecognized paraglacial landscape process that occurs

.in response to deglaciation Shroder, 1998 . Studiesof landslides along glaciated valley walls have docu-

Fig. 1. Location of study area in NW USA. Enlarged map showssouthern limit of late-glacial ice modified from USDA Forest

.Service, 1998 , location of the Rocky Mountain trench, part of thecontinental divide, and major modern rivers. Pleistocene glacialice is shown in a stipple pattern; flow directions of glaciers areshown by double arrowheads. SRsSwan Range, WRsWhitefishRange, TMsTeakettle Mountain. Range outlines are not meantto show extent of glaciation in those areas.

mented a range of forms and processes of failuresranging from deep-seated, bedrock failures Augus-

.tinus, 1995a to surficial slumping and falls of sedi-ment Owen and Sharma, 1998; Watanabe et al.,

.1998 . Understanding of timing and processes ofvalley wall failures after deglaciation, and the possi-bility of landslide reactivation, is important for as-sessing local and regional landslide hazards indeglaciated mountain regions. Examples of bedrock-involved landslides where the timing of the eventcan be shown to follow deglaciation directly haverarely been reported. Except for cases in a few

well-studied areas e.g., Panizza et al., 1996; Berris-ford and Matthews, 1997; Matthews et al., 1997;

.Soldati, 1999 , most examples lack direct evidencefor timing; usually, the dating of emplacement is

poorly constrained Whalley et al., 1983; Dawson et.al., 1986; Owen et al., 1995; Hewitt, 1998, 1999 .

Documented modern-day risks associated with con-tinued movement or reactivation of landslides thatoccurred during late Pleistocene deglaciation have

been described from only a few areas Pellegrini and.Surian, 1996 .

One such paraglacial landslide has been mappedalong the western front of the northern Swan Range,

northwestern Montana, USA Fig. 1; Harrison et al.,.1992; Smith, 2000a . The landslide is readily recog-

nized on aerial photographs, topographic maps with . 20-ft 6.1-m contours, and in the field Figs. 2 and

.3 . Documentation of the timing of emplacement, thepossible triggering mechanisms, and the indicationsfor continued movement of bedrock masses in thearea are the main purposes of this paper.

2. Geologic setting

The study area is in the Swan Range along theeastern side of the nearly flat-floored Flathead Val-ley, north of Flathead Lake in the northern Rocky

.Mountain region of North America Fig. 1 . Therange trends northsouth for a distance of 12 km in

( )L.N. SmithrGeomorphology 41 2001 309322 311

.Fig. 2. Stereo vertical aerial photographs of the landslide and tension fractures flight GS-CJ, 11 September 1946 ; crest of Swan Range is inupper part of photographs. Note that north is to left.

the area, forming the east side of the Rocky Moun-tain trench, a linear extensional basin that continuesNW into Canada. The west-facing flank of the rangeexhibits 1250 m of relief over a horizontal distanceof 2.5 km from the valley to the crest of the range.

The Swan Range is an east-dipping fault block ofmiddle Proterozoic greenschist-facies metasedimen-tary rocks of the Belt Supergroup. The Grinnell and

.Empire Formations sensu Winston, 1986 crop out .immediately upslope of the landslide Fig. 3 . Domi-

nant lithologies in both formations are argillite, silt-stone, and a few beds of quartzite and limestone .Johns, 1970; Harrison et al., 1992 . Most beddingranges in thickness from millimeter-scale lamina-tions of argillite to centimeter and decimeter beds ofquartzite and limestone. Locally, the Grinnell Forma-tion contains gray and red, siliceous and calcareousargillite with a few beds of feldspathic quartzite,whereas the Empire Formation is more greenish grayand contains a few thin limestone beds and anincreasing amount of calcareous argillite upsection.

The contact between the units is transitional Johns,.1970 . The crest of the range is underlain by the

.Helena Formation Fig. 3 , which contains morelimestone and dolomite than the subjacent units buthas similar bed thicknesses. Compressional tecton-ism created locally intense penetrative cleavage and

.small-scale folds and thrusts Johns, 1970 . The thinbedding and crosscutting fractures lead to disaggre-gation of the bedrock units into talus clasts mostly ofa few centimeters to a few decimeters in size. Bed-ding strikes N 30408W and dips 35578NE, mak-ing the west-facing flank of the range nearly perpen-dicular to bedding.

The Rocky Mountain trench and intermontanevalleys north of Flathead Lake were nearly filled bythe Flathead Lobe of the Cordilleran ice sheet during

late Pleistocene glaciation Richmond, 1986; Locke,.1995 . Glacial ice flowed SW out of mountains north

.of the Swan Range over Teakettle Mountain Fig. 1and merged with the Flathead Lobe that flowed SEdown the trench from central British Columbia and


Fig. 3. Simplified geologic map showing locations of cross-sec- X X X . tions AA, BB , and CC of Fig. 4 and photographs A, B,.and C of Fig. 9. Traces of tension fractures are shown as solid

.lines where sparsely vegetated group upslope of the crown and .as dashed lines where vegetated group south and west of crown .

The locations of water wells that are greater than 35 m deep areshown by circles. Solid circles denote wells that penetrated bedrockor large bedrock clasts in the lower portions of the well. An areawhere AbedrockB appears to be at anomalously shallow depths is

.outlined see text for discussion . Qlss landslide deposits; Qtstill; Qglsglaciolacustrine and outwash deposits; Qafsalluvial

.fan deposits; base is from Columbia Falls South 20-ft contours .and Doris Mountain 40-ft contours 7 1r2-min topographic

maps. Note that north is to left.

.with ice from the Whitefish Range Fig. 1 . Theresulting ice mass impinged against the northern partof the Swan Range and flowed southward. Thenorthern Swan Range contains glacial aretes, horns,

and cirques on its east and west flanks. Glacierheadwalls extend to the narrow crest of the range.Valley glaciers were smaller on the steeper andnarrower west-facing flank, with most of the rangesubjected to south-directed glacial flow by the Flat-head Lobe. The crest is typically less than 100 m inwidth except in a few low relief areas like that at

.Columbia Mountain Fig. 3 , suggesting that glacialerosion extended nearly to the range crest. Exceptalong glacier-margin stream channels, erosion onsteep slopes and weathering of the thinly beddedbedrock have obscured glacial trimlines, striations,and erratics in the immediate area of the landslide.

During the regions last glacial maximum at14 .12,70015,300 C years BP Carrara et al., 1996 ,

ice thicknesses exceeded 1 km in some valleys.Deglaciation of a cirque near the continental divideeast of the Swan Range had occurred before about

14 .12,200 C years BP Carrara, 1995 . Glaciolacus-trine sediments overlie till throughout the valley. Bythe time the Glacier Peak tephra was deposited in thevalley at 11,200 14C years BP, the Flathead Valleywas ice-free, the glacial lake had receded, and eolian

.dune fields were active Smith, 2000a .Although the Swan Range front is nearly linear

and has steep interfluves with faceted spurs suggest-ing recent faulting, neotectonism is enigmatic. Glacialerosion and deposition along the west-facing flankobscure evidence of range-front faulting. Few seis-mic events above magnitude 4 have been recordednear the northern Swan Range in about 20 years ofmonitoring, and no Holocene movement on the

range-front fault has been documented Stickney,.2000 . Like other areas in the intermountain seismic

belt of Montana, range-front faulting of the SwanRange appears to have a long recurrence interval.

Footslopes locally exceed 508 on longitudinal to-pographic profiles of interfluves along the west flankof the Swan Range near the Columbia Mountain

.landslide Fig. 4 . Most of the interfluves have steeplower segments and convex-up profiles. The longitu-dinal profiles are similar to glaciated slopes on thinlybedded and jointed bedrock in other areas .Augustinus, 1995b . The landslide itself has a pro-file that is generally concave-up or planar and attainsa slope of 378. The steep facets have been attributed

.to range-front faulting Konizeski et al., 1968 . How-ever, a lack of evidence for offset Pleistocene or


X . X . X .Fig. 4. Hillslope profiles of landslide BB and drainage interfluves north AA and south CC of the landslide. Vertical scale appliesto profile B; locations and bedrock contacts shown on Fig. 3.

Holocene deposits and glacial striations on thebedrock surfaces show the features were either

.formed Davis, 1920 or overridden by glacial ice.

3. Landslide form and deposits

3.1. Description

The Columbia Mountain landslide covers an areaof about 2 km2, has a height from toe to crown of1100 m, and a total length of 3430 m. The landslidesurface is entirely covered by crops, grassland, Dou-glas fir forest, or housing developments. Descrip-tions of deposit lithologies are derived from a fewroadcuts and 20 descriptive logs of water wells thatpenetrated the landslide deposits and subjacent sedi-ments. Thicknesses of the landslide interpreted fromwell logs are mostly between 3 and 30 m. However,the thickness exceeds 75 m in an area where block-to

slab-sized clasts terminology of Blair and McPher-.son, 1999 of Belt Supergroup bedrock are inter-

preted to rest on glacial deposits, as described below .Fig. 5 . Volumetric calculations using thicknessesfrom well logs and relief from 1:24,000-scale topo-graphic maps suggest an approximate volume of 40million m3 for the landslide, however, the thick-nesses and lateral extent of large bedrock clasts arepoorly constrained.

The landslide scarp has an upper portion shapedlike an inverted AvB with inward-facing slopes alongits flanks that meet at the crown near the contactbetween the Grinnell and Empire Formations. Thelower part of the scarp broadens and steepens at

.mid-slope on the mountainside Figs. 2 and 3 .Landforms and near-surface deposits define three

subunits in the depositional portion of the landslide . .Fig. 5 . From the toe to the head, the units are i asemicircular area where diamictons containing cob-ble and boulder-sized angular rock fragments have aseries of convex-west transverse ridges at the sur-

.face, ii a flat area behind the transverse ridgeswhere up to 10 m of sandy silt accumulated abovediamictons in an internally drained depression, and


.Fig. 5. A Vertically exaggerated profile of landslide showing lithologies and thicknesses of sediments interpreted from 15 descriptive logs .of water wells. The base of the bedrock clasts is not known and is shown by a question mark. B Cross-section at true scale.

.iii an area of lumpy topography underlain by di-amictons between the depression and the head of thelandslide. The toe of the landslide is convex-west inplanview, except at indentations on the north andsouth lateral sides, around areas now occupied byponds. The arcuate toe forms a distinct distal rim thatis 36 m above nearby land. The series of transverseridges that parallel the toe of the landslide partlywrap around the westernmost indentations into the

.toe of the landslide Figs. 2 and 3 . Topographybehind the westernmost transverse ridges slopesdownward toward the mountain front and the flat,internally drained depression. Relief between thetransverse ridges and the depression is about 5 m.

The area of lumpy topography upslope of thedepression is made up of small hills and depressionswith 520 m of local relief. Deposits near the sur-face are diamictons composed of fine cobble-throughvery coarse boulder-sized, angular to subroundedclasts of dolomitic argillite and minor quartzite in amatrix of sandy silt. A few of the observed clastswere glacially striated. Well-log data show diamictonthicknesses of at least 30 m. Three water wells in anarea about 300 m on a side were completed belowthe diamictons in more than 10 m of fractured

.bedrock at depths ranging from 27 to 65 m Fig. 3 .These depths are anomalously shallow for depth toBelt Supergroup bedrock on the hanging wall of the


range-bounding fault that separates the valley fromthe northern Swan Range. For instance, bedrock wasnot penetrated in wells 200 m north or 800 m southof this location. An interpretation more consistentwith regional thickness of valley-fill sediment is thatthe three wells partially penetrated one or morebedrock clasts that were transported along a scoop-shaped landslide failure plane. The distribution andsize of these inferred bedrock clasts beneath thesurficial diamicton are poorly understood due tosparse drill-hole data. The bedrock clasts are inter-preted to have been buried by landslide diamicton .Fig. 5a . Additional drilling or geophysical data areneeded to confirm this interpretation.

The landslide deposits rest on sandy silt, silt andclay, and silty diamicton deposits associated withproglacial lacustrine deposits and till and other sub-

.glacial deposits Smith, 2000b . The distal rockavalanche portion of the landslide essentially appearsto overlie glaciolacustrine silt and clay conformably .Fig. 5 .

3.2. Interpretation

3.2.1. ProcessesThe form and deposits of the Columbia Mountain

landslide suggest a two-stage failure of a bedrockinterfluve in the western flank of the Swan Range,shortly after deglaciation of the Flathead Valley. Thefirst stage of failure involved a rotational emplace-ment of the block-to slab-sized clasts of bedrock inthe subsurface, and the second stage was depositionof a rock avalanche, or sturzstrom, onto the valleyfloor.

Comparison of longitudinal profiles of bedrockspurs near to the landslide suggest possible dimen-sions of the spur that failed. Because most inter-fluves on the west flank of the northern Swan Rangehave steep lower slopes and smooth upper slopes,approximating the form of the failed interfluve byconnecting topographic contours across the scarp is

.reasonable Fig. 6 .The landslide was derived from a failure that cut

at a high angle across bedding in the Grinnell Forma-tion of the Belt Supergroup. The failure plane islikely to have developed along a joint set that strikesNW and dips perpendicularly to bedding, at 45

Fig. 6. Hypothetical reconstructed contours of bedrock interfluve .that failed; 200-ft 61-m contour interval for reconstructed hill-

slope.

.558SW Fig. 7 . The crown of the landslide is nearthe contact between the Grinnell and Empire Forma-

.tions Figs. 3 and 4 . A greater percentage of argilliteand less carbonate in the Grinnell may cause the unitto be weaker and more thinly bedded than the Em-pire, which could have contributed to the crowns

.position at the contact Fig. 4 .If the interpretation of the allochthonous nature of

bedrock clasts at the basal of part of the landslide iscorrect, the clasts represent the first stage of move-ment where bedrock rotated along a failure surface


Fig. 7. Lower-hemisphere, equal-area, stereonet projections ofstructural features in the immediate area of the landslide. Beddingorientations are shown as great circles, fracture orientations are

.shown by poles to the planes. A Data from bedrock exposures .north of scarp, B data from fractures in grabens above of scarp.

that propagated into valley-fill deposits. The crownof this rotational rupture surface is suspected to havebeen near or below where the scarp steepens at an

. .elevation of about 4900 ft 1490 m Figs. 3 and 8 .Failure of the lower part of the interfluve and em-placement of the bedrock clasts would have causedthe landslide to entrain valley-fill sediments withbedrock and till material derived from upslope. Thispossible initial failure of the lower part of the inter-fluve would have removed support for the rock masswithin the upper portion of the scarp.

The second stage of movement is represented bythe diamicton that extends about 1.5 km into the

.valley Fig. 3 . The abrupt convex toe, series oftransverse ridges, internally drained center, and di-mensions of the landslide suggest extremely rapid

. deposition by a rock avalanche sturzstrom Hsu,1975; Nicoletti and Sorriso-Valvo, 1991; Cruden and

.Varnes, 1996 . The form of the toe, transverse ridges,and the landslides topographically low center areevidence for outward flow during deposition of the

.landslide cf. Hsu, 1975 . The valley surface is es-sentially horizontal, suggesting that flow took placeover a short time immediately following failure, asthe diamicton deformed plastically before it came to

.rest cf. Schuster et al., 1995 . The dimensions of thelandslide are similar to the less mobile sturzstroms of

.Hsu 1975 . Whether or not the first stage rotationalfailure of the hypothesized lower slope occurred,failure along the upper scarp between about 4900 ft .1490 m and the crown was likely responsible forthe rock avalanche portion of the landslide. Thedetached rock mass and any overlying till possiblybecame airborne where the scarp steepens at mid-slope, resulting in a 5001000 m slide and fall ontothe valley surface. The second stage of the failuremost likely occurred immediately after the first stage.Avalanching rock debris is interpreted to have en-trained some glaciolacustrine sediments, previouslyexcavated valley-fill deposits, and bedrock clastsassociated with the first stage of failure.

Silty sand deposited in the internally drained de-pression on top of the landslide diamicton is inter-preted to be lacustrine based on its similarity tolaminated sand, silt, and clay throughout the upper

.Flathead Valley Smith, 2000a . The glaciolacustrinesilt below the landslide shows that the landslide fellinto a proglacial lake basin. The lacustrine silt in thedepression on top of the diamicton can be explained

either by deposition in the proglacial lake the land-.slide was, thus, partly subaqueous or that the


X . X .Fig. 8. Hypothetical reconstructed topographic profile of failed interfluve along BB Fig. 6 , modern profile BB in Fig. 4 , and possible .position of initial failure surface dashed line . The lower portion of the failed interfluve is interpreted to first have moved along a curved

plane into the valley; the upper portion failed second, producing the rock avalanche.

proglacial lake had receded and a separate lakeformed on the landslide after emplacement. Distribu-tions of glaciolacustrine sediments throughout the

.valley Smith, 2000a suggest a proglacial lake mayhave been 530-m-deep near the landslide.

3.2.2. TimingPonds immediately north and south of the rock

.avalanche Figs. 2 and 3 are two in a series ofdepressions interpreted as glacial kettles, or possiblypingo-remnants. The fact that the rock avalanche didnot fill the depressions suggests that these areas wereice filled and topographically higher than their sur-roundings at the time of landslide emplacement. Theorientations of the transverse ridges show that thelandslide flowed around the previously higher areason either side of the deposit. Ice blocks in the kettleareas could have been either grounded in the bottomof a glacial lake or stood higher than outwash in thesurrounding area. The possibility that the landslidefell entirely onto glacial ice and was then loweredonto the valley during downwasting is inconsistentwith the interpretation of a rotational emplacement inthe first stage of failure and the excellent preserva-tion of the toe and the transverse ridges on the rockavalanche. The close association of the ColumbiaMountain landslide with isolated ice blocks shows

that the hillslope failed shortly following deglacia-tion of the immediate area.

4. Possible triggering mechanisms

Steep faceted, glacially modified footslopes char-acterize each of the convex-up bedrock spurs on the

.western flank of the northern Swan Range Fig. 4 .Possible factors that contributed to instability of thefailed spur include glacial undercutting, reduction invalley-wall support and dilation of fractures parallel

.to the hillside upon deglaciation Panizza, 1973 , andseismicity associated with tectonism or isostatic ad-

.justments in the area cf. Ballantyne, 1997 . Catas-trophic failure of bedrock hillslopes immediately af-ter deglaciation of mountainous areas has been

proposed or proven for other areas Whalley et al.,1983; Dawson et al., 1986; Owen et al., 1995;Matthews et al., 1997; Panizza et al., 1996, 1997;

.Soldati, 1999 , but conclusive evidence for the actualtiming of the events in many areas is uncommon c.f.

.Hewitt, 1999 . The Columbia Mountain landslide isone such example where the triggering of the land-slide can be related to deglaciation. The lack ofevidence of surface faulting of glacial sediments thatare, in turn, overlain by the landslide suggests that


reduction of lateral support of an oversteepened hill-side was the most likely triggering mechanism. Stud-ies of slope stability along bedrock faces that wererecently deglaciated show that whatever the actualtriggering mechanism, slopes equilibrate quickly to a

.changed stress environment Augustinus, 1995a,b .

5. Indications of continued movement of bedrock

5.1. Description

Upslope from the crown of the latest PleistoceneColumbia Mountain landslide, several grabens, ten-

. .Fig. 9. Photographs of sparsely vegetated grabens. A Looking NE at a relatively large graben with colluvial fill located at AAB in Fig. 3 ; . .vertical bedrock walls trend N 65758E; dogs at arrows for scale are about 50 cm tall. B Looking east in the direction of bedrock dip

. .from ABB in Fig. 3 ; bedding is offset in a left-lateral sense across vertical fractures trending N 378E. C Looking east at a small grabenwith little colluvial fill; soil horizons were in vertical exposures beneath grass on right; orientation of graben is shown by line with doublearrowheads; bedding is offset in a left-oblique direction with 61 cm of horizontal separation along fractures that appear to sole into a joint

surface that dips toward the viewer and comes to the surface near the trees in the distance approximate location shown by strike and dip.symbol on a projected plane .


sion fractures, and sparsely vegetated scarps occuralong the upper part of the remnant of the failed

.bedrock spur Figs. 2 and 3 . Additionally, moresubtle, vegetated grabens and scarps occur fromsouth of the crown to 1 km down the SE flank of thespur. Most of the larger grabens with sparsely vege-tated scarps are NESW-trending, left-stepping, anden echelon in pattern; and they range from 100 to400 m in length and between 3 and 15 m in depth .Figs. 2 and 9 . Although the depth and width ofmost grabens are difficult to measure because ofcolluviation, where bedding surfaces can be matchedacross some smaller grabens, horizontal separations

.of 6080 cm have been measured Figs. 9b,c . Thegroup of grabens trends parallel to an interfluvialridge from the crown of the Columbia Mountain

.landslide to the crest of the Swan Range Fig. 2 .

Some smaller grabens at or within ;10 m of therange crest appear to sole into down-slope dipping

.fractures that are perpendicular to bedding Fig. 9c .The grabens are defined by steeply NW-and SE-dipping fractures that generally are normal to bed-ding and parallel to the remains of the interfluvial

ridge that failed during the late-glacial landslide Fig..7b . The fractures and grabens are similar to exten-

.sional structures related to rock flow sackung andlandsliding in other mountainous areas Savage and

Varnes, 1987; Varnes et al., 1989; Bovis, 1990;Dramis and Sorriso-Valvo, 1994, McCalpin and

.Irvine, 1995; Bisci et al., 1996 . However, in thenorthern Swan Range, these tension fractures andgrabens have only been recognized along an inter-fluve in the area near the crown of the ColumbiaMountain landslide.

.Fig. 10. A Map showing inferred distribution of actively extending rock mass shown by dotted line at AAB above the previous scarp. .Attitudes of prominant joints that strike nearly parallel to bedding but dip in the opposite direction are shown. B NWSE oriented

cross-section showing inferred movement direction of rock down the 308 slope.


5.2. Interpretation

The distribution and orientation of the tensionfractures and grabens suggest NW-directed gravita-tional spreading of bedrock in a discrete area NE ofthe crown of the late-glacial Columbia Mountainlandslide. The open network of brecciated bouldersfilling the bases of some grabens and vertical expo-sures of bedrock, colluvium, and rooted soil horizonson the sides of the grabens suggests that extension isactive across the features between the crown of the

late-glacial landslide and the mountain crest Figs. 3.and 9b,c . The vegetated grabens that extend downs-

lope below the crown show lower relief due to .significant colluvial filling Fig. 3 , and may be less

active than the upper grabens.Like bedrock extensional features in other land-

slide areas Chigira and Kiho, 1994; Dramis and.Sorriso-Valvo, 1994 , these features likely were initi-

ated by unloading of the hillside during or sometimeafter movement of the Columbia Mountain landslide.Unloading of the NW portion of the bedrock spur ispresumed to have allowed WNW extension of theremnant to occur.

If all of the fractures above the crown of theprevious landslide sole into a single zone of detach-ment or failure plane, like those described by Bovis . .1990 , Chigira and Kiho 1994 , Bovis and Evans . .1995 , and Sorriso-Valvo et al. 1999 , the planewould extend from the crest of the Swan Rangedown unnamed drainages north and south of the

.late-glacial landslide Fig. 10 . Evidence for an ac-tive failure plane emerging on the hillside NW of thefractures and grabens would be water seepage, topo-

graphic bulging, or production of debris Varnes et.al., 1989; Bovis, 1990; McCalpin and Irvine, 1995 .

Local steepening of the hillslope NW of the fracturesand grabens may be due to topographic bulging and

.flow downslope Fig. 10b but the data are notconclusive. Assuming a NW-dipping, concave singlefailure surface as much as 110 m below land surface,the rock volume is about 33 million m3. A morelikely scenario for future failure is that multiplefailure planes follow the NWSE striking, SW-dipping joint surfaces that are normal to bedding.These surfaces may define smaller bodies of rockthat could fail independently. Multiple active talusslopes and toppling bedrock clasts suggest intermit-

tent, small-scale failures are occurring near the top ofthe mountain.

Exposed soil profiles and openwork accumula-tions of colluvium in grabens, and left-stepping enechelon grabens with left separation are consistentwith extremely slow to very slow NW-directedmovement of the rock mass above the crown of the

.late-glacial Columbia Mountain landslide Fig. 10 .Deformation of the rock mass could possibly changeto an accelerated phase resulting in catastrophic col-

lapse cf. Evans, 1987; Chigira and Kiho, 1994;Dramis and Sorriso-Valvo, 1994; Bovis and Evans,

.1995; Bisci et al., 1996 . Strong shaking, such as thatassociated with a seismic event on or near the SwanRange-bounding fault, could trigger such a collapse.Where ground-rupturing earthquakes in NW Mon-tana have occurred they have recurrence intervals of

.roughly 48 ka Ostenaa et al., 1995 . The sparsedata on paleoseismicity and lack of evidence forHolocene displacement on the Swan Range front areconsistent with the long time elapsed since cata-strophic failure in the Columbia Mountain landslidearea. Further investigation of current movementacross, and the colluvial fill of, grabens above thelandslide is needed to assess rates of movement, andthe dimensions and probability of future catastrophicfailure.

6. Conclusions and implications

The Columbia Mountain landslide is the onlyrecognized major landslide in the upper FlatheadValley. The landslide contains deposits related to aninitial rotational bedrock failure of the lower part ofthe bedrock spur that led to a reduction of support ofthe upper part of the spur. Rock that fell or slid fromthe upper part of the scarp traveled a vertical dis-tance of 5001100 m, producing a rock avalanchethat reached 3450 m from the crown. Morphologicevidence shows that the rock avalanche traveledaround isolated blocks of glacial ice that occupiedpresent-day kettles. The landslide is presumed tohave been initiated by glacial undercutting of abedrock spur and dilation of bedding-normal jointsby unloading of the valley wall during glacial with-drawal. Available dates on deglaciation suggest thatice had withdrawn from the area between 12,000 and


15,000 14C years BP; the landslide apparently lastmoved at or before 12,000 14C years BP.

Extension of the remnant bedrock interfluve bygravitational spreading produced fractures andgrabens along the SE-facing flank of a remainingpart of the bedrock spur that failed in the landslide.Morphologic evidence suggests bedrock above thecrown of the Columbia Mountain landslide is mov-ing downslope toward the NW along joint surfacesthat act as failure planes. The deepest and leastvegetated grabens in the area trend NE from near thecrown of the previous landslide to the crest ofColumbia Mountain, defining the SE boundary of anactively extending bedrock mass. This potentialsource for a future landslide is about 1250 m verti-cally above and 2000 m horizontally away from anarea of residential development.

Reactivation of landslides initiated during land-scape adjustment contemporaneous with or closely

following deglaciation such as the late-glacial.Columbia Mountain landslide clearly poses less risk

to populations than more recent mass movements .Pellegrini and Surian, 1996 . However, where ex-tensional features can be spatially related to previouslandslides and shown to be active, further investiga-tion of the geologic hazards is warranted.

Residential development is preferentially occur-ring on the Columbia Mountain landslide depositsinstead of neighboring hillsides because the lower-angle slopes on the landslide provide flatter building

.sites Figs. 3 and 4 . The lack of evidence forcatastrophic movements in the area during the last12,000 years suggests that the hazard of a futurelandslide may only be slight. If a significant seismicevent should occur in the area, however, failure ofthe bedrock mass near the crest of the Swan Range ispossible.

Acknowledgements

Support of this work was provided by the Mon-tana Bureau of Mines and Geology and the MontanaGround-Water Assessment Program. Reviews by Drs.Mauro Soldati, Richard A. Marston, and an anony-mous reviewer helped to correct terminology andhighlight omissions, and were greatly appreciated.Additional reviews by R.N. Bergantino, E.G. Deal,

M.C. Stickney, W.A. Van Voast, and P.A. Hargraveof various versions of the manuscript improved itsclarity and content.

References

Augustinus, P., 1995a. Glacial valley cross-profile development:the influence of in situ rock stress and rock mass strength,with examples from the Southern Alps, New Zealand. Geo-morphology 14, 8797.

Augustinus, P., 1995b. Rock mass strength and the stability ofsome glacial valley slopes. Zeitschrift f ur Geomorphologie 39 .1 , 5568.

Ballantyne, C.K., 1997. Holocene rock-slope failures in the Scot-tish Highlands. PalaoklimaforschungrPalaeoclimate Research19, 197205.

Berrisford, M.S., Matthews, J.A., 1997. Phases of enhanced rapidmass movement and climatic variation during the Holocene: asynthesis. PalaoklimaforschungrPalaeoclimate Research 19,409440.

Bisci, C., Dramis, F., Sorriso-Valvo, M., 1996. Rock flow .sackung . In: Dikau, R., Brunsden, D., Schrott, L., Ibsen, M. .Eds. , Landslide Recognition. Wiley, Chichester, pp. 150160.

Blair, T.C., McPherson, J.G., 1999. Grain-size and textural classi-fication of coarse sedimentary particles. Journal of Sedimen-

.tary Research 69 1 , 619.Bovis, M.J., 1990. Rock-slope deformation at Affliction Creek,

southern Coast Mountains, British Columbia. Canadian Jour-nal of Earth Sciences 27, 243254.

Bovis, M.J., Evans, S.G., 1995. Rock slope movements along theMount Currie Afault scarpB, southern Coast Mountains, BritishColumbia. Canadian Journal of Earth Sciences 32, 20152020.

Carrara, P.E., 1995. A 12000 year radiocarbon date of deglacia-tion from the continental divide of northwestern Montana.Canadian Journal of Earth Sciences 32, 13031307.

Carrara, P.E., Kiever, E.P., Stradling, D.E., 1996. The southernlimit of Cordilleran ice in the Colville and Pend OreilleValleys of northeastern Washington during the Late Wisconsinglaciation. Canadian Journal of Earth Sciences 33, 769778.

Chigira, M., Kiho, K., 1994. Deep-seated rockslide-avalanchespreceded by mass rock creep of sedimentary rocks in AkaishiMountains, Central Japan. Engineering Geology 38, 221A230A.

Cruden, D.M., Varnes, D.J., 1996. Landslide types and processes. .In: Turner, A.K., Schuster, R.L. Eds. , Landslides Investiga-

tion and Mitigation. National Academy Press, Washington,D.C, pp. 3675.

Davis, W.M., 1920. Features of glacial origin in Montana andIdaho. Association of American Geographers, Annals 10, 75147.

Dawson, A.G., Matthews, J.A., Shakesby, R.A., 1986. A catas- .trophic landslide sturzstrom in Verkilsdalen, Rondane Na-

.tional Park, southern Norway. Geografiska Annaler 68A 12 ,7787.

Dramis, F., Sorriso-Valvo, M., 1994. Deep-seated gravitational


slope deformations, related landslides and tectonics. Engineer-ing Geology 38, 231243.

Evans, S.G., 1987. Surface Displacement and Massive Topplingon the Northeast Ridge of Mt. Currie, British Columbia.Geological Survey of Canada Current Research, Paper 87-1A,pp. 181-189.

Harrison, J.E., Cressman, E.R., Whipple, J.W., 1992. Geologicand structure maps of the Kalispell 1=2-degree quadrangle,Montana, and Alberta and British Columbia. U.S. GeologicalSurvey Miscellaneous Investigations Series Map I-2267, scale1:250,000.

Hewitt, K., 1998. Catastrophic landslides and their effects on theUpper Indus streams, Karakoram Himalaya, northern Pakistan.Geomorphology 26, 4780.

Hewitt, K., 1999. Quaternary moraines vs. catastrophic rockavalanches in the Karakoram Himalaya, northern Pakistan.Quaternary Research 51, 220237.

.Hsu, K.J., 1975. Catastrophic debris streams Sturzstroms gener-ated by rockfalls. Geological Society of America Bulletin 86,129140.

Johns, W.M., 1970. Geology and Mineral Deposits of Lincoln andFlathead Counties, Montana. Montana Bureau of Mines andGeology Bulletin 79, 182 pp.

Konizeski, R.L., Brietkrietz, A., McMurtrey, R.G., 1968. Geologyand Groundwater Resources of the Kalispell Valley, North-western Montana. Montana Bureau of Mines and GeologyBulletin 68, 42 pp., scale 1:63,360.

Locke, W., 1995. Modeling of icecap glaciation of the northernRocky Mountains of Montana. Geomorphology 14, 123130.

Matthews, J.A., Brunsden, D., Frenzel, B., Glaser, B., Wei, .M.M. Eds. , 1997. Rapid mass movement as a source of

climatic evidence for the Holocene. PalaoklimaforschungrPalaeoclimate Research 19, 444 pp.

McCalpin, J.P., Irvine, J.R., 1995. Sackungen at Aspen HighlandsSki Area, Pitkin County, Colorado. Environmental Engineer-ing Geoscience 1, 277290.

Nicoletti, P.G., Sorriso-Valvo, M., 1991. Geomorphic controls ofthe shape and mobility of rock avalanches. Geological Societyof America Bulletin 103, 13651373.

Ostenaa, D.A., Levish, D.R., Klinger, R.E., 1995. Mission FaultStudy. Seismotectonic Report 94-8, U.S. Bureau of Reclama-tion, Denver, 111 pp.

Owen, L.A., Benn, D.I., Derbyshire, E., Evans, D.J.A., Mitchell,W.A., Thompson, D., Richardson, S., Lloyd, M., Holden, C.,1995. The geomorphology and landscape evolution of theLahul Himalaya, Northern India. Zeitschrift f ur Geomorpholo-gie 39, 145174.

Owen, L.A., Sharma, M.C., 1998. Rates and magnitudes ofparaglacial fan formation in the Garhwal Himalaya: Implica-tions for landscape evolution. Geomorphology 26, 171184.

Panizza, M., 1973. Glacio pressure implications in the productionof landslides in the Dolomitic area. Geologia Applicata e

.Idrogeologia 8 1 , 289297.Panizza, M., Pasuto, A., Silvano, S., Soldati, M., 1996. Temporal

occurrence and activity of landslides in the area of Cortina .dAmpezzo Dolomites, Italy . Geomorphology 15, 311326.

Panizza, M., Pasuto, A., Silvano, S., Soldati, M., 1997. Landslid-ing during the Holocene in the Cortina dAmpezzo region,Italian Dolomites. PalaoklimaforschungrPalaeoclimate Re-search 19, 1731.

Pellegrini, G.B., Surian, N., 1996. Geomorphological study of theFadalto landslide, Venetian Prealps, Italy. Geomorphology 15,337350.

Richmond, G.M., 1986. Tentative correlation of deposits of theCordilleran ice-sheet in the northern Rocky Mountains. Qua-ternary Science Reviews 5, 129144.

Savage, W.Z., Varnes, D.J., 1987. Mechanics of gravitational .spreading of steep-sided ridges AsackungB . International As-

sociation of Engineering Geology Bulletin 35, 3136.Schuster, R.L., Cruden, D.M., Smith, W.L., 1995. Landslide of 2

July 1992 on Chief Mountain, northern Montana, U.S.A.Landslide News 9, 3334.

Shroder, J.F., 1998. Mass movement in the Himalaya: an introduc-tion. Geomorphology 26, 911.

Smith, L.N., 2000a. Surficial geologic map of the upper Flathead .valley Kalispell valley area, Flathead County, northwestern

Montana. Montana Bureau of Mines and Geology Ground-Water Assessment Atlas 2, part B, map 6, scale 1:70,000.

Smith, L.N., 2000b. Thickness of the confining unit in the Kalispellvalley, Flathead County, Montana. Montana Bureau of Minesand Geology Ground-Water Assessment Atlas 2, part B, map9 scale 1:100,000.

Soldati, M., 1999. Landslide hazard investigations in the Dolomites .Italy : the case study of Cortina dAmpezzo. In: Casale, R.,

.Margottini, C. Eds. , Floods and Landslides: Integrated RiskAssessment. Springer-Verlag, Berlin, pp. 281294.

Sorriso-Valvo, M., Gulla, G., Antronico, L., Tansi, C., Amelio,`C., 1999. Mass-movement, geologic structure and morpho-

.logic evolution of the Pizzotto-Greci slope Calabria, Italy .Geomorphology 30, 147163.

Stickney, M.C., 2000. Quaternary faults and seismicity in westernMontana. Montana Bureau of Mines and Geology SpecialPublication 114, scale 1:500,000.

USDA Forest Service, 1998. Glacial Lake Missoula and theChanneled Scabland: A Digital Portrait of Landforms of theLast Ice Age. USDA Forest Service Region 1, Missoula, MT,scale 1:1,000,000.

Varnes, D.J., Radbruch-Hall, D.H., Savage, W.Z., 1989. Topo-graphic and Structural Conditions in Areas of GravitationalSpreading of Ridges in the Western United States. U.S. Geo-logical Survey Professional Paper 1496, 28 pp.

Watanabe, T., Dali, L., Shiraiwa, T., 1998. Slope denudation andthe supply of debris to cones in Langtang Himal, CentralNepal Himalaya. Geomorphology 26, 185197.

Whalley, W.B., Douglas, G.R., Jonsson, A., 1983. The magnitudeand frequency of large rockslides in Iceland in the postglacial.Geografiska Annaler 65A, 99110.

Winston, D., 1986. Belt Supergroup Stratigraphic CorrelationSections, Western Montana and Adjacent Areas. MontanaBureau of Mines and Geology Geologic Map 40, 15 pp.

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