Postglacial colluvium in western Norway: depositional processes, facies and palaeoclimatic record

Postglacial colluvium in western Norway:depositional processes, facies and palaeoclimatic record

LARS H. BLIKRA* and W. NEMEC *Geological Survey of Norway, PO Box 3006 Lade, N-7002 Trondheim, Norway(E-mail: [email protected]) Geological Institute, University of Bergen, N-5007 Bergen, Norway (E-mail: [email protected])

ABSTRACT

The postglacial Quaternary colluvial systems in western Norway are arrays of steep

fans, often coalescing into aprons, developed along the slopes of valley sides and fjord

margins. The coarse debris, derived from weathered gneissic bedrock and its glacial-

till mantle, varies from highly immature to mature. The depositional processes are

mainly avalanches, ranging from rockfalls and debris¯ows to snow¯ows, but include

also water¯ow and debris creep. The mechanics and sedimentary products of these

processes are discussed, with special emphasis on snow avalanches, whose role as an

agent of debris transport is little-known to sedimentologists. The subsequent analysis

of sedimentary successions is focused on colluvial-fan deltas, which are very speci®c,

yet little-studied, coastal depositional systems. The stratigraphic variation and

depositional architecture of the colluvial facies assemblages, constrained by abundant

radiometric dates, are used to decipher the signal of regional climatic changes from

the sedimentary record. The stratigraphic data from two dozen local colluvial

successions are compiled and further compared with other types of regional

palaeoclimatic proxy record. The analysis suggests that the colluvial systems,

although dependent upon local geomorphic conditions, have acted as highly sensitive

recorders of regional climatic changes. The study as a whole demonstrates that

colluvial depositional systems are an interesting and important frontier of clastic

sedimentology.

INTRODUCTION

This paper discusses the sedimentology of thepostglacial Quaternary colluvium in westernNorway, based on case studies from two dozenlocalities in the Mùre-Romsdal region (Fig. 1).The colluvium comprises arrays of steep, gravellyfans developed along the valley sides and fjordmargins. The depositional processes include awhole range of avalanches, and the paper putsspecial emphasis on snow¯ows, whose mechan-ics and products are little-known to sedimentolo-gists. The paper further focuses on the faciesanatomy, chronostratigraphy and depositionalhistory of the colluvium, with emphasis oncoastal colluvial systems, which represent a

speci®c type of fan delta. In the ®nal part, thepalaeoclimatic record of the colluvial successionsis discussed and compared with other types ofregional palaeoclimatic data.

TERMINOLOGY

Colluvium is a general term for clastic slope-wastematerial, typically coarse grained and immature,deposited in the lower part and foot zone of amountain slope or other topographic escarpment,and brought there chie¯y by sediment-gravityprocesses (Holmes, 1965; Bates & Jackson, 1987;Blikra & Nemec, 1993a). In geomorphologicalliterature, colluvium is also referred to as `talus',

Sedimentology (1998) 45, 909±959

Ó 1998 International Association of Sedimentologists 909

`scree', `debris slope', `slope-waste deposits' and`hillslope (or hillside) deposits'; bedded col-luvium is often called `strati®ed debris slope'.Colluvial depositional systems have typically theform of relatively steep and short fans, or `cones',often coalescing into aprons. Modern colluvialfans are readily distinguishable from alluvial fans(Fig. 2), also due to the characteristic products ofavalanche processes.

Avalanches are rapid gravitational movementsof a wet or dry rock debris, snow, or their mixture,occurring on steep slopes. These mass move-ments are also called `catastrophic', because rapidoften means faster than an escaping human canrun and the speed of large avalanches can be ashigh as 50±80 m s±1 (Voight, 1978). Colluvialavalanches include rockfalls, debrisfalls, debris-¯ows, snow¯ows, and possibly also some slidesof rock or debris. The individual processes are

de®ned and discussed further in the text. (Thispaper suggests that the terms `debrisfall' and`debris¯ow' be written as single words, forconvenience and in semantic analogy to suchterms as rockfall, mud¯ow and snow¯ow.) Adebris¯ow or a snow¯ow rapidly descending asteep slope is said to be avalanching, because themass¯ow behaviour in such conditions tends tobe different than on a gentle slope.

The term `grain¯ow' refers to a cohesionlessdebris¯ow, possibly dry, whose movement ischaracterized by pervasive shear and ubiquitousgrain collisions (Bagnold, 1954; Lowe, 1976). Notevery cohesionless debris¯ow must necessarilybe a grain¯ow (cf. Lowe, 1982), for the phenom-enon of grain collisions ± the de®ning feature of agrain¯ow ± may be of minor volumetricimportance in many such frictional mass¯ows(Nemec, 1990b). This notion is particularly true

Fig. 1. Locality map of the Mùre-Romsdal region, western Norway, showing the main sites of detailed case studiesand the position of the Scandinavian ice-sheet margin in the Younger Dryas time. The localities are: (1) Gardvik,érstafjorden; (2) FlaÊskjñr; (3) Nordre Vartdal; (4) Tverrelva, Barstadalen; (5) Litlekoppen, Ytre Standal; (6) Try-ttesvora and Seljesvora, Hjùrundfjorden; (7) Vellesñterdalen; (8) Sunnylvsmoldskreddalen; (9) Korsbrekke, Hellesylt;(10) Fiksdal; (11) Fiksdalstrand; (12) Tomrefjorden; (13) Skorgedalen; (14) Midsund and Stormyr, Oterùya; (15)Eikesdalsvatnet; (16) Sunndalsùra; and (17) Gravem, Sunndalen.

910 L. H. Blikra and W. Nemec

Ó 1998 International Association of Sedimentologists, Sedimentology, 45, 909±959

for relatively thick, arenitic sand¯ows, whichoften carry `¯oating' cobbles or boulders. Aspointed out by Middleton & Southard (1978,chapt. 8), `the idea that grain¯ow alone isresponsible for the emplacement of thick, massivebeds of sand ¼ can only have been based on amisunderstanding of Bagnold's theory'.

Some authors have considered `avalanche' to bea speci®c type of mass movement, different fromrockfall, debris¯ow or slide (e.g. Selby, 1982,1994; Cas & Wright, 1987). However, the de®ni-tions offered are quite odd, suggesting a mecha-nism analogous to Bagnoldian grain¯ow anddepositional features similar to those of rockfallsand many cohesive debris¯ows. Likewise, someauthors have oddly restricted the term `avalanche'to snow¯ows only. The reader should also beaware that many colluvial avalanches, includingrockfalls and debris¯ows, have been describedunder the misleading label `landslide' ± which ingeomorphology means little more than a massmovement.

The descriptive terminology for gravel charac-teristics, including clast fabric notation, used inthe present study is after Harms et al. (1975) andCollinson & Thompson (1982). The fabric nota-tion uses symbols a and b to denote the long andintermediate axes of the clast, indices (t) and (p)

refer to the axes orientation transverse or parallelto ¯ow direction, respectively, and index (i)indicates axis imbrication.

DEPOSITIONAL SETTING

The Mùre-Romsdal region (Fig. 1) is representa-tive of the postglacial development and geomor-phic conditions in western Norway. The regionhas a rugged topography, with mountain peaks ofup to 1800 m, rolling plateaux, glacial cirques,large valleys and deep fjords. The bedrock, chie¯yPrecambrian gneisses, is strongly fractured andweathered, but has been eroded by glaciers andits primary weathering cover is thus thin andirregular. The NE topographic trend of the mainvalleys and fjords is related to the Mùre-Trùnde-lag Fault Zone, an early Mesozoic offshore linea-ment that impinges onto the Scandinavian land-mass in the region (Gabrielsen et al., 1984). Otherfjords and valleys, trending NW or NNW, arerelated to an older fault system that has beenreactivated in Mesozoic times and parallels theJan Mayen Fracture Zone (Aanstad et al., 1981).

The region was covered by the Scandinavianice-sheet during the last-glacial (Late Weichsel-ian) maximum, when the ¯oors and slopes of the

Fig. 2. A comparison of the distinctive features of colluvial fans and alluvial fans.

Postglacial colluvium in Norway 911


mountain valleys were mantled with glacial till. Itis uncertain whether the highest peaks protrudedabove the ice-sheet, as nunataks (Sollid & Sùrbel,1979; Nesje et al., 1987; Follestad, 1990; Larsenet al., 1995). The region was deglaciated, from theouter coast progressively to the south-east, be-tween c. 13 000 and 10 000 years BP (Mangerudet al., 1979; Mangerud, 1980; Fareth, 1987; Larsenet al., 1991). Since then, the glacial till and localkame terraces, perched on the valley sides, havebeen subject to erosion and intense resedimenta-

tion by gravitational processess, resulting inarrays of colluvial fans, solitary (Fig. 3) or co-alescing into aprons (Figs 4 and 5). Many of thesehave prograded into deep-water fjords, formingcolluvial-fan deltas (Figs 3 and 5D).

Studies of the Holocene glaciers in westernNorway have indicated that the postglacial cli-mate ¯uctuated considerably (Nesje & Kvamme,1991; Matthews & KarleÂn, 1992; Nesje & Dahl,1993; Dahl & Nesje, 1994; Nesje et al., 1995). Forexample, a major climatic deterioration occurrednear the end of the deglaciation phase, in theYounger Dryas time (c. 11 000±10 000 years BP),when the shrinking ice-sheet re-advanced (Fig. 1)and local alpine glaciers developed (Reite, 1967;Larsen et al., 1984). The Younger Dryas climatewas cold-oceanic (Larsen et al., 1984) and causeda marked increase in avalanche processes (Blikra& Nemec, 1993a; Nesje et al., 1994a; Blikra &Nesje, 1997).

The present-day annual temperatures at sealevel range between 12 and 14°C (July) and � 1°C(February), averaging 6±7°C. On higher mountainslopes, the mean annual temperatures are lower,ranging between 2°C and ± 4°C, respectively. Themean annual precipitation is between 1500 and2200 mm at sea level and even greater at higheraltitudes. The snow accumulation potential isgreatest on mountain slopes facing the SE, E andNE, due to the prevalent winds in these direc-tions. Modern avalanches abound, many reachingthe shoreline and causing considerable damage tothe local vegetation and roads (Lied, 1989), withoccasional loss of human life.

The region has a well-established curve ofpostglacial relative sea-level changes (Svendsen& Mangerud, 1987, 1990), which is used as apalaeogeographic reference framework in thepresent study. The postglacial isostatic uplifthas considerably exceeded the eustatic sea-levelrise, resulting in a relative sea-level fall by 50±80 m and gradual emergence of the colluvial-fandeltas (Figs 3 and 5D). Modern human activity,with numerous roadcuts and gravel pits, hascreated good outcrops of the colluvium.

COLLUVIAL PROCESSES AND FACIES

The colluvium is associated with bedrock moun-tain slopes of 35° to 50°. The slopes have variedmorphometric pro®les, with local inclinationsfrom 25±30° to as much as 70±90°. The colluvialfans and aprons that abut these slopes are steep,although their surface inclination depends upon

Fig. 3. Map showing the geomorphic setting and mo-rphometry of two colluvial-fan deltas at Gardvik (lo-cality 1 in Fig. 1). Note that the successive lobes (I-III)of the now-emerged fan deltas were built at progres-sively lower relative sea level. The fan surfaces shownumerous debris¯ow lobes and leveÂes, and some freshsnow¯ow deposits (asterisks). Letters A-E denote out-crop sections in the gravel pits.



Fig. 4. Map showing the geomor-phic setting and morphometry of asubaerial colluvial apron in NordreVartdal (locality 3 in Fig. 1). Notethat this valley-side apron is an ar-ray of coalesced colluvial fans.

Fig. 5. Examples of colluvial systems in the study area: (A) coalescent colluvial fans dominated by rockfalls, inOterùya; (B) a valley-side colluvial apron formed by debris¯ows, rockfalls and subordinate snow¯ows, at Litlekop-pen; (C) coastal colluvial apron dominated by snow¯ows, in fjord-head area at Korsbrekke; and (D) coalescentcolluvial-fan deltas dominated by snow¯ows, at Eikesdalsvatnet. Photographs from late spring 1992; for localities,see Fig. 1.



the depositional processes. Fans dominated byrockfalls are generally much steeper and shorterthan those formed by debris¯ows, whereas fansdominated by snow¯ows have intermediate gra-dients. Fan slope depends also upon the preva-lent avalanche size and headwall characteristics(Blikra, 1994). The steepest fans have slopes of43±45° (apex) to 20±30° (toe) and the least steepfans, 30±36° to 5±15°, respectively. Fans domi-nated by water¯ow processes, or long-runoutsnow avalanches, are no steeper than 13±15°and morphometrically transitional to commonalluvial fans.

The main depositional processes responsiblefor the development of the colluvial systems aresummarized in Fig. 6 and discussed in thefollowing sections. The review combines generalknowledge with our ®eld observations and givesthe genetic facies criteria used in the presentstudy.

Rockfall avalanches

Rockfall is a gravitational movement of a mass offragmented bedrock liberated abruptly from a cliff

or steep headwall, whose components tumblefreely downslope by rolling, bouncing and slid-ing; or it may be an analogous movement of asolitary rock fragment. The rock debris is charac-teristically angular, very immature. Large blocksoften disintegrate upon impact, spawning smallerclasts. Rockfalls have been studied extensively(Kent, 1965; Bjerrum & Jùrstad, 1968; Carson &Kirkby, 1972; Whalley, 1984; Statham & Francis,1986; Selby, 1994).

In a rockfall, clasts move through a series ofimpacts on the colluvial slope, which may retardthe clast, accelerate it, or stop it instantly. Factorsthat control the runout of a falling clast includeclast size (weight) and shape and the gradient androughness of the slope surface (Parsons & Abra-hams, 1987). Slope roughness is considered interms of the ratio of the diameter of the fallingclast to the diameter of the clasts resting on theslope. Clasts roll down easier on relativelysmooth, low-roughness slopes.

The mobility of a rockfall avalanche dependsalso upon the amount of the debris involved.Colluvial fans formed by larger rockfalls arelonger and less steep (Statham & Francis, 1986).

Fig. 6. Summary of the main depositional processes and facies of colluvial fans/aprons, with special reference to thepostglacial colluvium in western Norway.



In a voluminous rockfall, the interacting claststend to exchange and share their momentum, in amanner similar to that prevailing in a grain¯ow(Bagnold, 1954; Lowe, 1976; Campbell, 1990;Campbell & Brennen, 1983, 1985). However, theconcentration and collision frequency of clasts ina rockfall are lower than in a grain¯ow, whichrenders the larger clasts relatively free to movedownslope and segregate according to sizes(Campbell & Gong, 1986; Savage & Hutter, 1991).The heavier the clast, the greater its momentum,whereby large clasts tend to move faster andoutpace the smaller ones. The tongues of rockfallgravel thus show a marked downslope coarseningand a corresponding increase in surface rough-ness (Figs 6 and 7). The downslope part of arockfall deposit is typically openwork, composedof boulders and large cobbles, often with someblocks scattered further downslope as `out-run-ners'. The upslope part is thinner, comprised of®ner gravel and often in®ltrated with granulesand.

In a rockfall avalanche, the faster-moving largeclasts usually come to rest ®rst, and the trailingmass of ®ner debris then overrides this frontaldeposit, resulting in an apparent normal grading(Fig. 6). When a series of rockfall avalanchesdescend the slope along the same track, one afterthe other, the effective runout of the successiveavalanches may be shortened, resulting in anoverall upward coarsening (Statham & Francis,1986; Parsons & Abrahams, 1987; Nemec, 1990b,®g. 11).

The role of clast interactions in a rockfallavalanche increases with the volumetric concen-tration. The collisional stress component be-comes dominant when the concentration of clasts

exceeds c. 18 vol.% (Campbell, 1989a), at whichpoint the rockfall regime turns into the shear-¯owregime of an avalanching grain¯ow (Nemec,1990b, ®g. 9). This type of dynamic transforma-tion is common in the ®ner-grained `tails' ofrockfall avalanches (Bolt et al., 1975; HsuÈ , 1975;Wasson, 1979; Cruden & Hungr, 1986; Grimstad &Nesdal, 1991; Nemec & Kazanc�, 1999).

Rockfall debris is often subject to transientaccumulation in a ravine or other slope recess,and subsequently tumbles down as a loose massonce the threshold of frictional yield has beenexceeded. Likewise, rockfall gravel that has ac-cumulated in the apical part of a colluvial fanoften loses stability and avalanches further down-slope as a secondary rockfall, or a debris¯ow. Thelatter may be a grain¯ow, or a cohesive ¯ow if thefan-head gravel has been illuviated with mud(Nemec & Kazanc�, 1999).

Deposits. Colluvial rockfall deposits range fromscattered or randomly clustered boulders andcobbles, to distinct, tongue-shaped beds of im-mature coarse gravel characterized by markedupslope ®ning, common normal grading andmainly openwork texture (Figs 6 and 7). Theopenwork texture may be well-preserved (Nemec& Kazanc�, 1999), but the gravel in the presentcase is more often illuviated with sand and mud(Fig. 8). This secondary in®ll includes soil mate-rial rich in plant detritus, derived by contempo-raneous slopewash processes. Where emplacedsubaqueously, the rockfall gravel is commonly®lled with a wave-derived silt or ®ne sand,possibly bearing fauna shells or their crumbs.

The clast fabric of rockfall deposits is varied.Boulders and cobbles may show a `rolling' fabric,a(t) or a(t)b(i), when emplaced solitarily or as the

Fig. 7. An openwork, boulderygravel lobe formed by severalmodern rockfall avalanches, withnumerous `out-runner' blocks, up to4±5 m long, in the forefront; notethe slush¯ow tongue to the left.Valley-side colluvial apron inGlomsdalen. Photograph by P.A.Hole, from summer 1985.



frontal elements of an avalanche on a low-roughness slope. However, rockfall clasts gener-ally tend to adjust their resting positions to thestatic slope debris, are often subject to minorsecondary sliding and are vulnerable to re-orien-tation by subsequent avalanches. Many claststhus show an a(p) or random orientation, and theoverall fabric is disorderly. The fabric in somecases may show weak downslope alignment(Statham, 1973; PeÂrez, 1989; Bertran et al.,1997), and an aligned fabric a(p) or a(p)a(i) iscommon in the rockfall `tails', where the colli-sional stresses often prevail over the free-rollingmotion of the individual clasts.

A typical example of a local colluvial succes-sion relatively rich in rockfall deposits is shownin Fig. 9.

Debrisfall avalanches

Debrisfall (Holmes, 1965) is a process mechani-cally analogous to rockfall, but involving anolder, resedimented gravel, rather than freshlyfragmented bedrock, which means debris that isrelatively mature, subrounded to well-rounded(Fig. 6). Examples include ¯uvioglacial gravelfalling from a perched kame terrace down thevalley-side slope, or a river-derived gravel fallingdown the steep subaqueous slope of a Gilbert-type delta (Nemec, 1990b). The term `rockfall' insuch cases would be inappropriate and mislead-ing. Although debrisfall is not a separate processcategory, its distinction from rockfall may becrucial to a sedimentologist, because the differ-ence in maturity has important implications as to

the debris provenance, mobility and runoutpotential.

In the present case, debrisfall processes arethought to have played a role in the resedimen-tation of mature glacigenic gravel from steepmountain slopes. However, the importance ofdebrisfall deposits in most of the colluvial fansand aprons is minor, compared to that of eitherrockfall or debris¯ow deposits. The relativescarcity of debrisfalls is attributed to the matrix-rich texture of the glacial till, preventing even thecoarsest gravel from rolling freely down the slope.The few debrisfall deposits that have been recog-nized are thought to indicate the local slopeconditions where the glacial mantle included akame terrace or has been strongly eluviated byslopewash. Apart from the roundness and high

Fig. 8. Detail of a rockfall deposit in vertical outcropsection. The large basal interstices of this immature,openwork gravel have been ®lled with a strati®edmuddy sand derived by sheetwash. The lens cap(centre) is 5 cm. Colluvial fan at Midsund (locality 14in Fig. 1).

Fig. 9. Portion of outcrop section and detailed log of acolluvial fan dominated by rockfall deposits, at Stor-myr (locality 14 in Fig. 1). The log is normal to thebedding. Letter code in the log: CS � clast-supportedgravel texture; OW � openwork gravel texture.



apparent maturity of debris, the characteristics ofthese sporadic debrisfall deposits are similar tothose of rockfall deposits (see previous sectionand Fig. 6).

Debris¯ow avalanches

Debris¯ows are a common and important agent ofcolluvial sedimentation in a wide range of cli-matic zones (Rapp, 1960, 1963, 1987; Voight,1978; Wasson, 1979; Rapp & Nyberg, 1981, 1988;Keefer & Johnson, 1983; Eisbacher & Clague, 1984;Williams, 1984; Nyberg, 1985; Cruden & Hungr,1986; Van Steijn, 1988, 1996; Blikra et al., 1989;Brabb & Harrod, 1989; Jibson, 1989; Kotarba,1991; Nicoletti & Sorriso-Valvo, 1991; Luckman,1992; Rickenmann & Zimmermann, 1993; Nemec& Kazanc�, 1999). In the geomorphological liter-ature, debris¯ows are referred to also as `debristorrents', `debris slides', `debris streams', `debrisavalanches', `grain¯ows', `earth¯ows', `¯owingdebris masses', `lahars', `landslides', `mud¯ows',`mudspates', `rock ¯ows', `rocky mud¯ows',`sand¯ows', `sand runs' and possibly othernames. Both subaerial and subaqueous debris-¯ows have long been a subject of laboratory and®eld studies, and the existing knowledge isconsiderable (Johnson, 1970; Hampton, 1975,1979; Lowe, 1976, 1982; Brunsden, 1979; Lawson,1982; McTigue, 1982; Haff, 1983; Savage,1983; Johnson & Rodine, 1984; Nemec & Steel,1984; Campbell & Brennen, 1985; Campbell,1986, 1990; Pierson, 1986; Costa & Wieczorek,1987; Van Steijn & Filippo, 1987; Van Steijn &Coutard, 1989; Savage & Hutter, 1989, 1991;Nemec, 1990b; Takahashi, 1991).

Debris¯ow is a type of sediment-gravity ¯ow,de®ned as a gravitational movement of a shearing,highly concentrated, yet relatively mobile, mix-ture of debris and water. In some climatic zones,debris¯ows may involve an admixture of snow orslush, as is common on the mountain slopes inNorway. Dry debris¯ows are rare in the presentcase, but common in other climatic settings(Melton, 1965; Oberlander, 1989; Nemec & Ka-zanc�, 1999). A debris¯ow may be turbulent,although a subaqueous turbulent mass¯ow is tobe classi®ed as a turbidity current (Lowe, 1982).

In rheological terms, a debris¯ow is a simple- topure-shear plastic ¯ow, which means that thegranular material tends to spread, behaves like asingle-phase ¯uid and has a ®nite yield strength.The shear strength derives from a combination ofcohesive forces, due to the electrostatic bondsbetween clay particles, and frictional forces ± due

to particle interlocking and the frictional resis-tance against interparticle slip (cohesion of clay-free silt and ®ne sand is very low, highlydependent upon water content and consideredto be negligible; KeÂzdi, 1979.) Because one of thetwo forces usually predominates, debris¯owshave been categorized as cohesive and cohesion-less (frictional), respectively (Nemec & Steel,1984), with the classic mud¯ow and grain¯owas end-member models (Lowe, 1982). This rheo-logical distinction corresponds with the conven-tional engineering classi®cation of `soils', ornatural clastic materials, into two analogouscategories (KeÂzdi, 1979; Keedwell, 1984).

When the shear stress, or the downslopecomponent of the material's weight, exceeds theyield strength, the material begins to shear and¯ow, behaving like a dense, viscous ¯uid. Theapparent viscosity of a debris¯ow is the `bulk'dynamic viscosity of the shearing mass, whichmeans the dynamic resistance to shear derivingfrom the interparticle ¯uid and the sedimentparticles themselves. For example, a grain¯owbehaves like a high-viscosity ¯uid even when thegranular material is dry (Lowe, 1976; McTigue,1982; Melosh, 1983). In mechanical terms, theapparent viscosity is a dynamic coef®cient trans-lating the applied shear stress into the shear-strain rate, and thus controlling the mobility andinternal shear regime of a debris¯ow, includingthe possible development of turbulence.

The dynamic viscosity of a debris¯ow is nec-essarily much higher than that of water, but mayvary greatly from one ¯ow to another; it may be ashigh as c. 103 Pa s (Sharp & Nobles, 1953) or aslow as c. 10 Pa s (Pierson, 1986). The apparentviscosity of both cohesive and cohesionless de-bris¯ows varies, depending upon the sedimenttexture and the content of water, or possiblysnow. Further, the viscosity in some debris¯owsis roughly constant and independent of the shear-strain rate, like in Newtonian ¯uids, whereas inothers it varies with the strain rate like in non-Newtonian ¯uids. The former case represents theBingham plastic category and applies reasonablywell to cohesive debris¯ows (Johnson, 1970;Hampton, 1975, 1979; Brunsden, 1979; Johnson& Rodine, 1984), whereas the latter case repre-sents non-Bingham plastics and applies to cohe-sionless debris¯ows (McTigue, 1982; Haff, 1983;Campbell, 1990), as well as to those waterydebris¯ows in which the collisional stresses ingravel fraction predominate despite the presenceof a ¯uidal muddy matrix (Bagnold, 1954; Taka-hashi, 1991).



As the frictional debris¯ow accelerates, itdilates and its apparent viscosity increases. Theviscosity of such a `shear-thickening' debris¯owincreases typically as a quadratic function of theshear-strain rate (McTigue, 1982; Campbell &Brennen, 1983, 1985; Haff, 1983; Savage, 1983;Takahashi, 1991). A sand¯ow or sand-rich de-bris¯ow, particularly subaqueous, may com-mence its movement as a `shear-thinning' lique-®ed ¯ow, in which the shear strain makes thesand expel its pore water and becomes lubricatedby it, and the apparent viscosity is thus lower athigher strain rate. However, the liquefaction ofloose sand is relatively rapid, and the subaqueousmass¯ow may then: (1) cease to move after a fewseconds or so; (2) accelerate and turn into a high-density turbidity current (Middleton & Southard,1984); or (3) turn into a ¯owslide ± a debris¯owthat glides on a thin, shearing basal layer (Camp-bell, 1989b; Nemec, 1990b; see also Bolt et al.,1975; Lang & Dent, 1983; Lang et al., 1989). Theundrained conditions required for sand liquefac-tion are rare on subaerial mountain slopes, butmany of the cohesive debris¯ows in our studyarea were probably triggered by a spontaneousliquefaction of the glacial till.

The conceptual rheological framework outlinedabove has been the basis for the interpretation ofdebris¯ow deposits in the present study. Ourobservations indicate that the colluvial debris-¯ows, although derived mainly from glacial till,have shown considerable variation. The glaci-genic mantle was likely heterogeneous, and itswater content probably varied greatly on both alocal and temporal basis. The sur®cial part(c. 1 m) of the mantle was affected by vegetationand weathering, and the debris¯ow sources alsoincluded the weathered bedrock and upper-slopecolluvium, and possibly local snowpacks.

The debris¯ow deposits are pebbly to boulderygravel beds ranging from matrix- to clast-support-ed. Matrix varies from a muddy, poorly sorted sandto nearly arenitic sand or sandy granule gravel. Therelationship between the bed thicknesses andmaximum clast sizes (Fig. 10) indicates that thematrix-rich debris¯ows can generally be classi®edas cohesive and the matrix-poor ones as cohesion-less (cf. Nemec & Steel, 1984). The high apparentcompetence of the colluvial debris¯ows (see theb-values in Fig. 10) can be attributed to two factors.Firstly, the debris¯ows, particularly those poorerin matrix, are likely to have borne an admixture ofsnow. A dense snow would increase the ¯owcompetence, and as the snow melted, the gravelbed thickness would decrease, resulting in an even

higher maximum clast size/bed thickness (MCS/BTh) ratio (the data in Fig. 10A are, in fact, fromdebris¯ow deposits of the snowy Younger Dryasperiod.) Secondly, the colluvial debris¯ows de-scended very steep slopes and were rapid. Thedownslope component of clast weight on a steepslope would be higher than the slope-normalcomponent, which might effectively delay theclast settling and allow the ¯ow to maintainrelatively large clasts over the short descent time.

Furthermore, the geometry of the depositsindicates marked variation in the spreading modeof the debris¯ows, even on the same colluvial fanor similar fan surfaces. Some debris¯ows havespread as relatively broad lobes, whereas others

Fig. 10. Relationship between the maximum clast size(MCS) and bed thickness (BTh) of debris¯ow deposits,based on data from three colluvial fans. Letter symbols:n � number of data; r � coef®cient of linear correla-tion; b � regression coef®cient (gradient of the least-squares regression line).



have formed highly elongate tongues, passingupslope into furrows with leveÂes (Figs 11, 12 and13). This variation is apparently shown by bothcohesive and cohesionless debris¯ows, and isthought to represent ¯ows of high and lowviscosity, respectively (Fig. 6). The low-viscositybehaviour is attributed to a relatively high contentof water or lubricating slush, whereas the high-viscosity behaviour is attributed to a low watercontent or an admixture of dense drier snow.

The rheological behaviour of a debris¯ow isknown to be highly sensitive to relatively minorchanges in water content (Pierson, 1980, 1981,1986; Johnson & Rodine, 1984). When a debris-¯ow is relatively rich in water and/or slush, itseffective shear strength is considerably reduced.For example, Johnson & Rodine (1984; p. 286)noted that the strength of a debris¯ow comprising

silty mud and pebble gravel had changed fromc. 320 Pa at a water content of 14 wt.% to c. 40 Paat a water content of 16á5 wt.%. The shearstrength decreased by an order of magnitude withan increase of water content by merely 2á5 wt.%.The density of the mixture had changed from2200 to 2130 kg m±3, respectively, which is achange of c. 3% only. The driving shear stress ofthe debris¯ow on a given slope would thus havechanged insigni®cantly, while the apparent vis-cosity, mobility and internal regime of the ¯owwould be affected quite drastically. On the otherhand, a signi®cant admixture of a dry or dampsnow, further densi®ed by the load and motionsof the host debris, would render the apparentviscosity of a debris¯ow very high and result inhigh apparent cohesive strength (Tusima, 1973;Fukue, 1979; Salm, 1982).

Fig. 11. Coalescing debris¯ow lobeson the surface of a valley-side col-luvial fan in Grasdalen. Note thatsome of the debris¯ows have spreadas relatively broad lobes, whereasothers have formed elongate ton-gues. The slope catchment at thepresent time is prone to snow ava-lanches, to which the scattered largeblocks on the fan surface are as-cribed. The short dimension of thephotograph is c. 350 m and thegeographical north is downwards.Aerial photograph by FjellangerWiderùe A/S.



Deposits of high-viscosity debris¯ows

The deposits attributed to high-viscosity debris-¯ows (Fig. 6) are beds of matrix-rich to clast-supported gravel, 0á4±1á7 m thick, which arerelatively extensive and tabular on a gravel-pitoutcrop scale. They contain `¯oating' cobbles orboulders and commonly show inverse grading.The grading in matrix-supported beds is of`coarse-tail' type, limited to the coarse gravel,and can be attributed to the presence of a `rigidplug' above the shearing lower part of thedebris¯ow (Johnson, 1970; Naylor, 1980). Thegrading in clast-supported beds is more of a`distribution' type, involving a wider range ofclast sizes, and is attributed to clast collisions,with the upward displacement of larger clasts dueto dispersive pressure (Bagnold, 1954; Lowe,1976; Walton, 1983) combined with kinematicsieving (Middleton, 1970; Scott & Bridgwater,1975; Savage & Lun, 1988; Jullien et al., 1992).The beds of subaqueous debris¯ows often showalso a weak normal grading at the top (Fig. 14A),probably due to the shear at the ¯ow/waterinterface (Hampton, 1972). The debris¯ow bedsin Gilbert-type delta foresets tend to be thinnerand have more distinct boundaries (Fig. 15). TheMCS/BTh data (Fig. 16) indicate that the compe-tence of subaqueously emplaced debris¯ows wasgenerally lower, probably because these moremobile ¯ows were more watery or incorporatedwater when plunging into the sea (cf. Nemec &Steel, 1984 ®g. 23).

The lobate geometry of the debris¯ow depositsof this category (Fig. 6) is most apparent on thefresh surfaces of many colluvial fans (Fig. 11).Some of the debris¯ow lobes can be traced nearlyto the fan apex, where they are thinner, muchnarrower and often have bouldery leveÂes. Otherlobes are more `detached' from the apex, pinchingout upslope in a mid-fan zone. The latter debris-¯ows have apparently bypassed the steepest fanslope with little or no deposition. Their bypasstracks are shallow scours, with or without leveÂes.Another trace of a bypassing debris¯ow are`debris horns': cusps of debris¯ow material accre-ted onto the upslope sides of large blocks pro-truding above the fan surface. The high-reliefobstacles apparently cause local plastic `freezing'of a bypassing debris¯ow, if not too watery.

The clast fabric in the debris¯ow beds of thiscategory is mainly a(p) or a(p)a(i), at least for thelarger clasts (Figs 14 and 15). The steep fronts ofmodern debris¯ow lobes commonly show an a(t)fabric, with large boulders that have apparently

Fig. 12. Fresh deposits of low-viscosity, slush-rich de-bris¯ows on vegetated colluvial aprons in Norangs-dalen (near locality 8 in Fig. 1). (A) Note the `digitated'shape of this elongate lobe, due to the lateral splays; theslope height is c. 100 m. (B) Note the elongate, tongue-like geometry of this lobe, with a relatively thick,coarser `head' and thinner upslope `tail'; the latterpasses upslope into an erosional furrow with leveÂes;the slope height is c. 50 m.



rolled down the front or been `bulldozed' by it.These frontal features, however, are seldomrecognizable in the outcrop sections of olderdebris¯ow deposits. Likewise, the individualbeds show little or no downslope change in clastsize within the limits of a large gravel-pit outcrop,although the deposits of apparently similar deb-ris¯ows on modern fan surfaces often showdownslope thinning and ®ning (Mears, 1980;Statham & Francis, 1986). It is possible that thelatter trend occurs on a lateral scale larger thanthat of a gravel pit, for the debris¯ow beds are, infact, often thinner and ®ner grained (in terms ofMCS) in the downfan outcrops. However, thisapparent trend may re¯ect the greater spreading(smaller thickness) and lesser competence of themore mobile debris¯ows.

Deposits of low-viscosity debris¯ows

The deposits attributed to low-viscosity debris-¯ows (Fig. 6) typically include an erosive, fur-row-like track with leveÂes and a highly elongate,

coarse-fronted gravel lobe in the downslopedepositional zone (Figs 12, 13, 14B, 17 and 18).The watery or slushy debris¯ows are more sensi-tive to the slope topography, and their tracksoften show bends. The tongue-shaped gravellobes are characterized by a relatively thick,clast-supported, bouldery to cobbly `head' thatpasses upslope into a thinner, ®ne cobbly topebbly, clast- to matrix-supported `tail'. There islittle visible difference between the subaerial andsubaqueous deposits of this category. In the ¯ow-parallel outcrop sections, these debris¯ow lobesare recognizable as long, but highly asymmetricallenses, either solitary or multiple. The multiplelenses are often stacked upon one another in an`imbricate' fashion, dipping upslope, or show amore complex style of stacking, with the succes-sive debris¯ow lobes having overridden andoverstepped one another to a variable extent(Figs 6 and 14B). The head of the lobe typicallyshows a tractional a(t) fabric (Figs 12B and 18),whereas the fabric in the tail is more varied; a`shear' fabric a(p) or a(p)a(i) is common, but

Fig. 13. Fresh deposits of slush-richdebris¯ows and snow¯ows on veg-etated colluvial aprons in Norangs-dalen. (A) Note the `digitated' lobeof a slushy debris¯ow to the left; theleveÂed narrow tracks and small,`digitated' debris lobes of slush¯owsto the right; and the scattered and`patchy' debris deposited by snow-¯ows in the lower middle; this col-luvial apron's height is c. 50 m. (B)Frontal part of a lateral `spill-over'lobe of a watery debris¯ow, showingmainly `rolling' clast fabric andcommon imbrication (¯ow towardsthe viewer); the main lobe hasspread to the right (see in the back-ground).



`rolling' fabric occurs as well, particularly ingravel beds showing normal grading.

Our observations from active fans support thenotion that the watery debris¯ows are triggeredby a rapid melting of slope snowpack (Harris &Gustafson, 1988) or episodes of heavy rainfall(Caine, 1980; Lawson, 1982). The reported veloc-ities of such debris¯ows are of the order of a fewmetres per second. As the watery debris¯owdescends a steep slope, the ¯ow accelerates anddilates rapidly, the dispersion of its larger clastsincreases and its two component phases ± thecoarse gravel and the ¯uidal matrix ± tend tobehave almost independently. The dynamic re-gime of the avalanching debris¯ow thus ap-proaches that of a debrisfall (Campbell, 1989a;

Nemec, 1990b, ®g. 9), with the clast-size distri-bution and the mean free distance between thelarge clasts becoming important controlling pa-rameters. The large clasts tend to move down-slope according to their own momentum, pushingthemselves, or `streaming' in engineering par-lance, through the ¯uidal ®ner material to thefront of the ¯ow (Suwa, 1988; Nemec, 1990b). Abouldery to cobbly `head' develops, whose grow-ing thickness, increasing clast concentration andgreater internal friction render its speed lowerthan that of the following `tail'. The latter isusually turbulent, so long as it retains an abund-ant interstitial ¯uid; otherwise, the apparentviscosity rises and suppresses the turbulence.The faster-moving tail continually feeds the head

Fig. 14. Debris¯ow deposits in out-crop sections. (A) Portion of rela-tively thick, tabular bed of unsorted,matrix-supported gravel attributedto a high-viscosity debris¯ow em-placed on the subaqueous slope of aconical colluvial-fan delta; note thecoarse-tail inverse to normal grad-ing, with `¯oating' cobbles andboulders in the middle part. (B)`Imbricate', lenticular gravel bedsattributed to low-viscosity debris-¯ows emplaced on the subaqueousslope of a Gilbert-type fan delta;note that the debris¯ow tongueshave clast-supported, bouldery tocobbly `heads' and a thinner, peb-bly, clast-to matrix-supported `tails';note also the thicker, more tabulardeposits of high-viscosity debris-¯ows in the lower part. The whiteruler is 23 cm. Examples from thecolluvial fan-delta complex atTomrefjorden (locality 12 in Fig. 1).



with the lubricating ¯uid, which percolates intothe frontal part due to gravity (in conditionspromoting such debris¯ows, the colluvial slope isusually water-logged; otherwise, the ¯uid wouldescape into the porous substratum.) The percola-tion rate depends upon the size distribution ofdebris, especially its ®ner fractions. If the perco-lation is not too rapid and the head lubrication iseffective, the debris¯ow may arrive on the lowerslope without much change in the internalregime, and the runout may be considerable.When the head of the ¯ow spreads and comes to ahalt, its ®nal in®lling by the watery tail occurs,often with a thin sheet or multiple ®ngers ofpebbly sand deposited in the forefront by theescaping turbid water. The fronts of slushy

debris¯ows are more `digitated' due to thepostdepositional melting and secondary mobili-zation (Figs 12 and 13).

The onset of turbulence in a cohesive debris-¯ow, approximated as a Bingham plastic, occurswhen the Hampton number of the ¯ow exceeds1000, which may mean ¯ow velocity in excess ofc. 3 m s±1 (Middleton & Southard, 1984). Suchvelocities can be attained by a debris¯ow, even ifbrie¯y, on the steep colluvial slope. The morewatery debris¯ows have lower viscosities andbecome turbulent much easier, but they are alsogenerally thinner. Vigorous turbulent churningmay occur when the Froude number of the ¯ow

Fig. 15. Details of the subaqueous foreset of Gilbert-type fan delta, Tomrefjorden (locality 12 in Fig. 1). (A)Diamict, mud-rich bed with `¯oating' cobbles andboulders, attributed to cohesive, high-viscosity debris-¯ow. (B) Two beds of matrix-supported gravel,attributed to high-viscosity debris¯ows, underlain by aclast-supported gravel unit attributed to subaqueouslyemplaced, debris-rich snow¯ows; note the coarse-tailinverse grading in the lower (lighter-shade) debris¯owbed. Transport direction is to the left. The beds areinclined at c. 25°, and the photographs are parallel tobedding. The ruler is 23 cm.

Fig. 16. Relationship between the maximum clast size(MCS) and bed thickness (BTh) of the subaerial andsubaqueous debris¯ow deposits in a colluvial-fan deltaat Eikesdalsvatnet (locality 15 in Fig. 1). Letter sym-bols: n � number of data; r � coef®cient oflinear correlation; b � regression coef®cient. Note thatthe b-value in the lower diagram is c. 25% smaller andthe regression line intersects the MCS axis near the zeropoint, which suggests less competent and cohesionlessdebris¯ows (for methodological details, see Nemec &Steel, 1984).



signi®cantly exceeds 1 (Pierson, 1986), and suchsupercritical conditions can easily be reached ona steep slope. However, a subaerial debris¯ow isunlikely to become fully turbulent, like a New-

tonian ¯uid, with the large eddies breakingcontinually down into progressively smallervortices. Unlike powder snow¯ows and somesubaqueous debris¯ows, the subaerial debris-¯ows do not burst into fully turbulent ¯owsdownslope. A watery subaerial debris¯ow 1±1á5 m thick may show vigorous turbulence, butcan hardly suspend clasts larger than smallpebbles (Pierson, 1986), due to the limited scaleand highly dissipative character of the eddies.The turbulence may be mixing the debris con-tinuously, but the larger clasts are lifted less,settle faster and segregate from the small ones.The signatures of turbulence in a debris¯owdeposit are thus normal grading and possibly a`rolling' fabric, which are features commonlyrecognizable in the tails of the colluvial debris-¯ows. It should be noted that subaqueousdeposits with such features are to be classi®edas deposits of high-density turbidity current of(Lowe, 1982).

When the ¯uidal matrix of the debris¯owpercolates into the head and/or substratum toorapidly and the slope is very steep, the tail turnsinto a frictional debris¯ow, possibly still slightlyturbulent, where Bagnoldian dispersive stressesmay prevail. The energy and frequency of clastscollisions (or the `granular temperature' of the¯ow) may be high in a fast-moving, leveÂe-con-®ned ¯ow, whereby the coarser clasts will tend tobe pushed inwards and upwards, and be furthertransferred to the head of the ¯ow in a conveyor-belt fashion (Hirano & Iwamoto, 1981; Johnson &Rodine, 1984; Pierson, 1986; Takahashi, 1991).The tail gravel will be clast-supported, inverselygraded and have an a(p) or a(p)a(i) fabric. Thehead gravel will have an a(t) fabric in the frontaland upper part.

Fig. 17. Elongate, tongue-shapedgravel lobes passing upslope intoleveÂed erosive furrows, formed bywatery debris¯ows on a valley-sidecolluvial apron in Liadalsdalen(south of locality 2 in Fig. 1); theupslope extent of the avalanchetracks corresponds to the upperlimit of the mountain slope's glaci-genic mantle. The width of the pic-ture is c. 250 m. Photograph by O.Longva, from early summer 1991.

Fig. 18. Fresh deposit of a low-viscosity debris¯ow onthe surface of a colluvial fan in Glomsdalen (upslopeview). The tongue-shaped lobe has a bouldery frontwith transported tree logs, and passes upslope into anarrow, gravel-lined furrow with leveÂes. Photograph byP.A. Hole, from summer 1985.



Some debris¯ow tails have a matrix-supported,nearly bimodal texture, though otherwise showsimilar characteristics as mentioned above. Thematrix is a poorly sorted sand, often quite muddy.The deposit shows coarse-tail inverse grading andfrontal concentration of large clasts, which hasbeen attributed to a combination of dispersivepressure and the conveyor-belt mechanism (John-son & Rodine, 1984; Takahashi, 1991). However,Bagnold's (1954) theory predicts that the colli-sions of sand grains result in relatively smalldispersive pressure, and the latter also tends to besuppressed by cohesive forces. It is more likelythat the large clasts in a shearing sand move awayfrom the margins and lower boundary of thecon®ned debris¯ow due to the so-called Magnuseffect: a lift force acting normal to the shearboundary and related to the rotation of the largeclasts by the local shear-velocity gradient. Thiseffect is thought to occur, for example, in somesnow¯ow avalanches (Hop®nger, 1983).

The low-viscosity colluvial debris¯ows arecomparable to the watery debris¯ows describedfrom mountain ravines and narrow gorges (Okudaet al., 1980; Suwa & Okuda, 1983; Johnson &Rodine, 1984; Pierson, 1980, 1986; Takahashi,1991). However, the elongate debris¯ow lobes inthe present case have been deposited on opencolluvial slopes of 30° to 5±15°, whereas theravine-con®ned debris¯ows have reportedly beenturbulent and erosive on slopes as gentle as 5±8°.This difference may have two reasons: (1) thecolluvial debris¯ows are due to slumping ofwater-logged sediment, rather than to sedimententrainment by ¯ood-water surges; and (2) themechanism of ¯ow-head lubrication by ¯uidalmatrix may be less effective, or shorter-lived, onvery steep, rough and porous colluvial slopes.

The debris¯ow tracks are erosive furrows(Figs 12, 13, 17 and 18), widening and shallow-ing-out in the downslope direction. Their widths,excluding leveÂes, are several times their depth.The depth decreases from 50 to 70 cm in theupper reaches to less than 5±10 cm in the apicalpart of the associated gravel lobe, which is mainlynon-erosive. The furrows are probably scoured byturbulence, but their lower reaches may be ruts ofnonturbulent mass¯ow. The leveÂes are due to theplastic `freezing' of the lateral margins of the ¯ow,where the shear rate is at a minimum (Johnson &Rodine, 1984). These `dead zones' include cob-bles and boulders that have been shoulderedaside by the debris¯ow front (Pierson, 1986).

The morphology of modern deposits shows thatthe turbulent tail of a low-viscosity debris¯ow

may locally breach the leveÂe and form a smaller,secondary lobe of gravel outside the main track,particularly at a local bend (Figs 6, 12 and 13).The debouching may deplete the parental debris-¯ow of its lubricating ¯uidal part and reduce therunout. Our observations further indicate thatwhen the head of a debris¯ow `freezes' and jamsthe steep descent path, the tail may movesideways, develop a new head and form asecondary lobe further downslope; or the tailmay override the halted head and form a new lobedirectly in front of it. These secondary lobes aregravelly and their heads may be nearly as coarseas the primary ones (Fig. 13B). When formed on alower slope, at a relatively late stage of debris¯owevolution, the secondary lobe is relatively thin,composed of sandy pebble gravel, and its head ispoorly developed. These characteristics may helpto distinguish between the primary and second-ary lobes of debris¯ows, although the distinctionin outcrop sections is not easy.

The subaqueous foresets of colluvial-fan deltasabound in deposits of both high- and low-viscos-ity debris¯ows, associated with the graded bedsof sand and pebbly sand attributed to high-density turbidity currents. Turbidity currentscan be generated by subaerial debris¯ow ava-lanches plunging into the water (Hampton, 1972;Weirich, 1989). However, the turbidites are muchmore abundant in the foresets of Gilbert-type fandeltas, whose formation involved stream¯owprocesses, which suggests that the latter wereprobably more important in generating turbiditycurrents, by hyperpycnal out¯ow or mouth-barcollapsing (Nemec, 1990b).

Some illustrative examples of local colluvialsuccessions relatively rich in debris¯ow depositsare shown in Figs 19 and 20.

Snow¯ow avalanches

Snow¯ows (Fig. 5C,D), also referred to as `snowavalanches', have been a subject of intenseresearch by glaciologists, geomorphologists andengineers, primarily because these avalanches area serious hazard in mountainous terrains in manyparts of the world. Surprisingly, snow¯ows havedrawn little sedimentological interest, althoughthey are known to carry often abundant rockdebris.

The literature on snow avalanches is extensive,but widely scattered in the glaciological andtechnical journals, special reports and symposiavolumes. There are several monographs (Mantis,1951; Kingery, 1963; Oura, 1967; Perla &



Martinelli, 1975; Voitkovskiy, 1977; Fukue, 1979;Washburn, 1979; Colbeck, 1980; Glen et al., 1980;Gray & Male, 1981; Ramsli, 1981); a number ofreviews (Luckman, 1971, 1977, 1978; LaChapelle,1977; Mellor, 1978; Perla, 1980; Schaerer, 1981;Salm, 1982; Hop®nger, 1983); and a wealth ofempirical studies (e.g. Dent & Lang, 1980, 1982;Lied & Bakkehùi, 1980; Mears, 1980; Narita, 1980;

Schaerer & Salway, 1980; Gubler, 1982; Lang &Dent, 1983; McClung & Schaerer, 1983, 1985;Lang et al., 1985; Salm & Gubler, 1985; Hutteret al., 1989; Hermann & Hutter, 1991; McClung &Tweedy, 1993) and various theoretical consider-ations (e.g. Perla et al., 1980; Bakkehùi et al.,1981, 1983; Dent & Lang, 1983; Norem et al.,1985, 1987, 1989; McClung & Lied, 1988; Gubler,1989; Gubler & Bader, 1989; Lackinger, 1989;Nishimura & Maeno, 1989; Nishimura et al.,1989). Some publications are focused speci®callyon slush¯ows (Washburn & Goldthwaite, 1958;

Fig. 19. Portion of an outcrop section and detailed logof the subaerial part of a fan delta rich in debris¯owdeposits, Tomrefjorden (locality 12 in Fig. 1). Thedownslope direction is to the left, and the log is normalto bedding.

Fig. 20. Portion of an outcrop section and detailed logof the subaerial part of a colluvial-fan delta showing analternation of debris¯ow and snow¯ow deposits,FlaÊskjñr (locality 2 in Fig. 1). The downslope directionis to the right, and the log is normal to bedding. Notethe radiocarbon dates in the log. Texture code CS andOW as in the caption to Fig. 9.



Nobles, 1966; Washburn, 1979; Hestnes, 1985;Nyberg, 1985; Conway & Raymond, 1993).

The literature on snow¯ows seems to exceedthat on debris¯ows, but bears less consensus. Thescienti®c language of the publications varies fromthe formal terminology of physics to the descrip-tive parlance of geomorphology, and there seemsto be relatively little exchange of results and ideasbetween the two respective groups of researchers.The following short review attempts to synthesizethe diverse views and classi®cations into areasonably simple and coherent conceptualframework, which has been the basis of our studyof colluvial snow¯ow deposits. The review refersto the classi®cation in Fig. 21 and focuses onthose aspects of snow¯ow processes that arerelevant to a sedimentologist.

Snow rheology

Snow is a peculiar material, with a highlycomplicated rheological behaviour, althoughsnow¯ows are generally akin to sediment-gravity¯ows. On a microscopic scale, snow consists of aframework of ice particles, which may be singlecrystals or their larger aggregates. The mechanicalproperties of snow vary with the particle charac-teristics (size, shape, packing), interparticlebonds (number and type of contacts, pore-wallarea) and the spatial arrangement of bondedparticles (fabric). Snowpacks are commonly lay-

ered (bedded) and texturally heterogeneous dueto their incremental accumulation under variableweather conditions. A fresh powdery snow maycontain c. 90 vol.% air, whereas an older, com-pacted and recrystallized snow may be nearly asdense as ice.

In broad terms, snow can be regarded as aplastic material, whose shear strength ranges frommainly frictional to mainly cohesive. However,snow has a high and irreversible compressibilityand an equally high thermodynamic instability.The shear deformation of snow is volumetric,rather than purely deviatoric (Salm, 1982), andthe content of free water depends on the ambienttemperature (Ambach & Howorka, 1966), whichrenders snow rheologically quite different fromall natural clastic materials.

Snow has a high capacity to dissipate stress, bydensi®cation. For example, the compaction ofsnow from an initial density of 300 kg m±3 to adensity of 900 kg m±3 absorbs three times asmuch energy as its compaction from 600 to900 kg m±3 (Brown, 1980a). Further, the contentof water is very important. The free water spreadsthrough the pores by capillary forces, but maybegin to percolate through the snow, due togravity, when the saturation exceeds c. 1 vol.%(Langham, 1981). At low, `pendular' saturations(< c. 7 vol.%), the interparticle water ®lm isdiscontinuous, the snow abounds in air bubblesand is relatively strong, as is common with early

Fig. 21. Tentative classi®cation ofthe relevant ¯owing media in termsof the relative proportion of rockdebris, snow and water involved.The terminology pertains to subae-rial ¯ows. For subaqueous condi-tions, the solid-laden, high-densitywater¯ow and the turbulent waterydebris¯ow would be classi®ed asturbidity currents. The classboundaries are based partly onBeverage & Culbertson (1964) andFukue (1979).



spring snowpacks (Male, 1980). At higher, `fun-icular' saturations (> c. 7±12 vol.%), the water®lm becomes continuous and the snow is rela-tively weak. When the water saturation exceedsc. 25 vol.%, the snow turns into slush (Fig. 21).Saturations of up to 35±50 vol.% have beenreported from watery slush¯ows (Nobles, 1966;Mellor, 1978). In undrained laboratory tests, slushreaches a maximum water saturation of 58 vol.%(Nyberg, 1985).

Frictional interparticle forces dominate in freshpowdery snow and dry granular snow, althoughcohesive forces are also active in the former:down to a temperature of ± 5°C in dry conditionsand down to ± 30°C in regelational `wet' condi-tions; below these temperatures, a granular snowdevelops in either case. Cohesive forces predom-inate in damp/wet snow, due to the adhesiveeffect of free water; in semibonded snow, due tothe interparticle water ®lms and crystallinebonds; and in bonded (sintered) snow, due tothe crystalline bonds alone. Sintering occurs attemperatures below the melting point, but thebonds form faster as the temperature nears thelatter (Fukue, 1979). The shear strength variesgreatly with the snow density, from c. 102 Pa for avery low-density snow (q £ 100 kg m±3) to c. 106

Pa for a high-density snow (q » 600 kg m±3). Thedensity of a slushy or densi®ed snow may reach c.700 kg m±3, and a strongly sintered snow ap-proaches the density of ice (917 kg m±3). How-ever, a high water content renders snow weaker(Salm, 1982), hence the strength of slush is verylow.

As the cohesive strength prevails in some snowvarieties and the frictional strength dominates inothers, it has been suggested that snow¯ows beclassi®ed into cohesive and cohesionless (orfrictional) (Nyberg, 1985). Notably, this rheolog-ical distinction is similar to the classi®cationsuggested for debris¯ows (Nemec & Steel, 1984).

The static angle of internal friction for snow ishigh, typically in the range of 35±46°, and also thedynamic angle of friction for a ¯owing loose snowis high, 25±32° (Fukue, 1979), comparable to theangle for classic grain¯ows (Bagnold, 1954).However, the dynamic angle depends stronglyupon the shear-strain rate (Fukue, 1979). Thedynamic viscosity increases almost exponentiallywith the density, much stronger for dry than fordamp or wet snow (Salm, 1982, ®g. 15). Forexample, laboratory tests show that the dynamicviscosity of a wet snow increases from c. 103 toc. 106 Pa s due to a change in density from 700 to900 kg m±3, whereas a dry snow shows a similar

change in viscosity due to a density increase ofless than 10 kg m±3. Rapid rebonding of particlesoccurs in the failing framework of dry snow,whereas extensive bond slips and particle growthoccur in wet snow. The rate of particle growth atan ambient temperature of + 20°C is nearly 50%higher than at 0°C, and the increasing size anddecreasing number of the particles and bondsrender wet snow considerably weaker (Salm,1982). Further, the effective stress at the particlecontacts in wet snow is signi®cantly reduced bythe repulsive electrostatic forces due to thedouble-charged layers at each solid±liquid inter-face of the water ®lm separating the particles.These repulsive forces are stronger at low watersalinities (Colbeck et al., 1978).

The constitutive relationship between the shearstress and shear-strain rate for snow is generallyconsidered to be nonlinear, with a rate-dependentapparent viscosity, but depends strongly uponthe strain rate itself (Mellor, 1978; Fukue, 1979;Salm & Gubler, 1985). The relationship is nearlylinear at very high rates (Fukue, 1979), particu-larly for powdery and granular snows, which isprobably why turbulent snow¯ows are akin toturbidity currents. A dry loose snow at high strainrate shows shear-softening and strong dilation,after the initial collapse of the particle framework(McClung, 1979; Narita, 1980). At lower strainrates, the relationship varies. At a very low rate,snow creeps due to the rupture of interparticlebonds and the localized melting and slip withinthe ice crystals and their aggregates. The creep isoften approximated as a linear viscous deforma-tion (Brown & Lang, 1973; Salm, 1982), althoughthe strain rate decreases with depth, roughlylinearly, as the viscosity increases linearly withthe normal stress (Desrues et al., 1980; McClung,1980; Watanabe, 1980). The constitutive relation-ship is nonlinear once the snow mass hasdetached itself and ¯ows at a low rate. Theviscosity then increases almost exponentiallywith the strain rate, much faster than the snow'sYoung modulus, whose increase is nearly linear(Kry, 1975). This is the state of `brittle' ¯owage(Fukue, 1979), with the strain breaking andthoroughly reworking the particle framework.Importantly, the shear strain does not dilate theloose snow, but rather increases its density. Thesnow thus densi®es, yet its rheological behaviouris similar to that of dilational grain¯ow (LaCha-pelle, 1977). The shear stress/rate relationshiphas been modelled as a third-order dilation,proportional to the cube of the rate (Brown &Lang, 1973).



The apparent viscosity increases with the strainrate until the microstructure of the snow has beenpervasively reworked. The viscosity then shifts toa lower level and remains approximately con-stant, almost like in a Newtonian ¯uid. This is thestate of `ductile' ¯owage (Fukue, 1979), whichcauses snow hardening in laboratory shear-boxtests. On a mountain slope, the ¯owing snowmass at this point accelerates and its furtherbehaviour may be quite different. For example,turbulence may be ignited in a dry snow¯ow,which will then rapidly dilate into suspension.Unfortunately, the available empirical data aresparse and often inconclusive, and the existingconstitutive models show large discrepancies(Voellmy, 1955; Mellor, 1978; Salm & Gubler,1985).

For example, Mellor (1978) reported that theshear rate in a purely deviatoric (nonvolumetric)snow¯ow increased rapidly with the shear stress,which would imply a decreasing viscosity andthus a contractional non-Newtonian behaviour.However, the velocity pro®les measured by Salm& Gubler (1985) and Nishimura & Maeno (1989)indicate that the apparent viscosity of relativelydense snow¯ows remains constant, as in a Bing-ham plastic ¯ow, but increases (proportionally tothe second power or so of the shear rate) in faster¯ows with greater particle dispersion. This dila-tional non-Newtonian behaviour would be akin tothat of a grain¯ow.

It would thus appear that the rheologicalbehaviour of snow varies greatly, depending onthe snow type and the shear stress level. There isno universal constitutive equation for snow, butrather equations for its behaviour under particularconditions (Salm, 1982). As a convenient simpli-®cation, the ¯ow of a relatively dense snow(> 300 kg m±3) is often considered to be contrac-tional (`shear-thinning'), and the ¯ow of a lower-density snow to be dilational (`shear-thickening'),although these behaviours may change or combinein an evolving snow¯ow (Brown, 1979, 1980b).

Snow¯ow types

As snow accumulates on a mountain slope, thenormal stress in the snowpack increases with theincreasing thickness, or load. Concurrently, theyield strength of the snowpack increases, due todensi®cation. Strong winds may increase thesnowpack density, and a wind-blown snow isalmost invariably denser than one deposited byfreefall. When the stress increases faster than thestrength, a gravitational failure occurs. Other

triggering mechanisms include falling cornicesand snowpack destabilization by creep. The winddrift of snow is very important, for it affects thesnowpack thickness, builds cornices on the leeside of mountain ridges, and ®lls the mountainravines and gullies with thick snowpacks thatmay later obstruct runoff.

Snow typically fails on slopes of 30±40°, but awet snow can fail on a surface of 10°, if relativelysmooth or icy. The snowpack may slough by aseries of small retrogressive failures, or collapseen masse as an avalanche (Perla, 1978). Twotypes of snow¯ow avalanche can be distin-guished, with possible intermediate varieties(Hop®nger, 1983).

Dense snow¯ows. Cohesive or other densesnow is usually subject to a line failure, leadingto a translational slide. A broad slab of shearingsnow then descends the slope, leaving well-de®ned crown and ¯ank scarps in the snowpack.The result is a dense snow¯ow, also called`¯owing avalanche [Ger. Fliesslawine]' or `granu-lar-snow gravity ¯ow' (Hop®nger, 1983). Thesesnow¯ows may be cohesive or cohesionless, andrange from pseudolaminar to turbulent. The snowinvolved may be a dry granular or sintered snow,a damp granular or semibonded snow, a wetgranular snow, or a slush. A hard snow slab(sintered or semibonded) may resist fragmenta-tion, and the avalanche may carry large snowblocks to the depositional area (Fig. 22C). If fullyfragmented, the avalanche will consist of gravel-sized hard clods mixed with a granular orpowdery snow (Fig. 22B, 3A). Cohesive snowallows `snowrollers' (snow balls) to form andbreak repeatedly.

Dense snow¯ows are comparable to debris-¯ows, and are rheologically modelled as cohesiveor cohesionless plastics (Salm & Gubler, 1985;Lang et al., 1989; Nishimura & Maeno, 1989), or aviscoelastic material (Voellmy, 1955; Mellor,1978; Salm, 1982) in the case of very slowmovement. Their steep-fronted depositional lobesand leveÂed tracks (Figs 22A & B, 23A and 24) aresimilar to those of debris¯ows. In fact, slush¯owwas originally de®ned as `a mud¯ow-like ¯owageof water-saturated snow' (Washburn & Goldth-waite, 1958). However, the snow¯ows have shear-dependent viscosity and behave more like non-Bingham plastics (Lang & Dent, 1983; Lang et al.,1985). In a ¯ow of roughly constant density (say,c. 400 kg m±3), the snow hardness may range overan order of magnitude and the kinematic viscositymay vary from c. 400 to c. 2500 m2 s±1 (Maenoet al., 1980; Lang et al., 1985).



The dense snow¯ows on mountain slopes havetypical velocities of 10±30 m s±1 and are mainlysupercritical (Fr > 1). The mean density rangesfrom c. 150 kg m±3 for damp snow to c. 350 kg m±3

for wet or hard snow (Mellor, 1978), and the snowparticle concentration is typically in the range of33±44 vol.%. These snow¯ows are `groundavalanches' (Hop®nger, 1983), with the lowerboundary cut at the ground surface or deeply inthe snowpack.

Dense snow¯ows often arrive on the lowerslope by turning into ¯owslides: the ¯ow `freez-es' plastically from the top downwards, but

keeps gliding on a thin, shearing basal layer;and when the front brakes, the snow massdeforms by means of discrete listric thrusts andlongitudinal shear zones (Salm & Gubler, 1985).The resulting longitudinal and transverse shears/ridges (Fig. 22D, 23A; Mears, 1980, ®gs 6,7 & 8)resemble those observed in debris¯ows that havecompleted their movement in a sliding fashion(Nemec, 1990b, ®gs 28, 29 & 30). Slush¯ow canmove on a smooth slope of less than 2° (Nobles,1966), but stop on a rough, bouldery slope ofmore than 70° (Luckman, 1971) (Fig. 24).

Fig. 22. Features of fresh snow¯ow deposits. (A) Dense snow¯ow lobes on a colluvial fan at Gravem (locality 17 inFig. 1), spring 1993. (B) Snow¯ow deposit composed of pebble- to boulder-sized, hard snow clods mixed with apowdery snow entrained en route by the avalanche; note the leveÂe to the right (downslope view, shovel for scale);Gravem, spring 1993. (C) Large, hard snowballs at the surface of a melting snow¯ow lobe; Gravem, spring 1993.(D) Longitudinal shear zone, rich in meltout debris, in the lateral part of a dense snow¯ow lobe; Hjùrundfjorden(locality 6 in Fig. 1), spring 1988.



Powder snow¯ows. A fresh, powdery snow or aloose older snow that has lost cohesion (drygranular snow) is usually subject to a pointfailure, which quickly expands. The term `pow-der snow¯ow' is used here broadly for such low-density snow¯ows, composed partly or entirely ofpowdery snow and most often turbulent. Thesesnow¯ows come to rest on slopes ranging from34° to horizontal, but most commonly on those of17±28° (McClung & Schaerer, 1983). The turbu-lent snow¯ows, also referred to as `airborne snowavalanches' (Hop®nger, 1983), have generallylonger runouts.

A turbulent cloud typically develops at the topof a powder snow¯ow and grows at the expenseof the latter. The entrainment of dry snow by theambient air begins when their relative velocityreaches c. 7 m s±1, and the turbulent cloudbecomes fully developed when the velocity ex-eeds c. 10 m s±1. The trailing cloud acceleratesand turns into an overriding surge, which mayattain speeds of 30±70 m s±1, possibly up to125 m s±1, thus outrunning and overtaking thedepleted parental ¯ow. The thickness of theresulting turbulent avalanche may vary from lessthan 10 m to more than 100 m, depending on thesnow volume involved and the travel distance.The density may vary from c. 10±20 kg m±3 in thelower part to little more than 1±2 kg m±3 at thetop of the ¯ow (Mellor, 1978).

Many powder snow¯ows burst into a ground-hugging, highly turbulent cloud shortly after theirrelease, probably due to an internal shock wave(Hop®nger, 1983). For a snow¯ow with meanvelocity U and thickness h, the kinematic wavethat travels through the ¯ow has a celerityuk � 1á5 U, whereas the corresponding long dy-namic wave has a celerity ud � (qh)1/2 and effec-tively propagates with a velocity ud* � U + (qh)1/2

where q is the acceleration due to gravity. When theFroude number of a ¯ow, Fr � U/(qh)1/2, exceeds2, the kinematic wave moves faster than thedynamic wave, which may generate a `roll-wave',or growing shock bore. A velocity of 10 m s±1, towhich an avalanche can acceleratevery quickly on asteep slope, corresponds to Fr > 2. Ultra-rapid ¯owregime, with Froude numbers as high as 6±8á5, canbe reached by large powdersnow¯owson very steepslopes (Schaerer & Salway, 1980).

A turbulent powder snow¯ow is typically a`surface avalanche' (Hop®nger, 1983), whoselower boundary is cut shallowly in the snow-pack, rather than at the ground surface. These¯ows are akin to turbidity currents (Hop®nger,1983, ®g. 4), and are often modelled as such.They have mean frontal velocities of 12±60 m s±1,more typically 20±40 m s±1, and mean densitiesof 50±300 kg m±3 (Mellor, 1978; McClung &Schaerer, 1983). The density in the basal partmay be 2±4 times greater than the mean. Themaximum velocity is near the base, at a height ofroughly 1/10 of the ¯ow thickness. For a ¯owthickness h and a typical range of slope inclina-tions of 60° to 30°, the maximum velocity zone istheoretically at a height of 2á1h1/2 above the base(McClung & Schaerer, 1983, 1985). The maximumvelocity of the ¯ow body is 2±2á5 times the frontalvelocity. The concentration of snow is low,

Fig. 23. Features of fresh snow¯ow deposits on a col-luvial fan at Bùndergjerde, Skorgedalen (locality 13 inFig. 1). (A) Upslope view of the upper part of a densesnow¯ow lobe, showing large axial furrow, leveÂes andtransverse listric shears; the furrow was scoured byyounger, bypassing snow¯ows. The backpack to the leftis c. 50 cm long. (B) Rock boulders, an uprooted birchtree, numerous broken branches and abundant ®nerdebris melting out at the surface of a dense snow¯owlobe; the lobe at this stage is still c. 3 m thick. The lenscap is 5 cm. Photographs from late spring 1993.



typically 5±25 vol.%, which renders internalfriction very low. For example, the dynamicangle of friction for a concentration of 5 vol.%is less than 1°, and such a snow¯ow will behavelike a Newtonian ¯uid.

Some powder snow¯ows are non-turbulent.They are typically mixtures of powdery snowand an older granular snow, and are akin tocohesionless debris¯ows. Gubler (1989) has com-pared the existing models for dry non-turbulentsnow¯ows and concluded that the Salm & Gubler(1985) model of `partly ¯uidized' snow¯ow pre-dicts the observed velocities and runouts mostcorrectly. The model invokes a lower `¯uidized'layer (liquidized by the basal shear), with anexponentially convex velocity pro®le, overlain bya nonshearing `rigid plug'. Their relative thick-nesses depend on the total ¯ow thickness and thegradient and roughness of the slope surface, asthe latter factors control the ¯ow speed andrunout. The convex velocity pro®le, resemblingthat of a grain¯ow (Middleton & Southard, 1984),suggests a concentration gradient and is attribut-ed to the snow texture, considered to be a mixtureof powder snow and granular snow rich in hardclods of ®ne pebble size. The powder snow andair, incorporated en route by the ¯ow, act as aninterstitial ¯uid in the basal part. The energydissipation in the lower layer is mainly due topseudolaminar shear, so long as the slope rough-ness elements have a relief/spacing ratio smallerthan 0á03, but is more due to the snow-clodcollisions (Bagnoldian `viscous regime') when theslope roughness is greater. The constitutive rela-tionship used in this model invokes a non-Newtonian behaviour with the power coef®cient

for the strain rate taken to be 0á6 (contractional¯ow).

Snow¯ow avalanches have an enormous impactforce and high transport competence, which ren-ders them capable of carrying huge and heavyobjects. Published measurements indicate that theimpact force may vary from 10 to 20 kPa forturbulent powder snow¯ows to 1000±2000 kPa forsome large denser snow¯ows, with a reportedmaximum of 3900 kPa. For example, Voellmy(1955) calculated an impact force of 8á3 kPa for asnow¯ow that displaced a 120-t locomotive by20 m. Hop®nger (1983) reported a large avalanchethat detached a three-storey barracks from itsconcrete basement and carried it across topograph-ic ridges and gorges over a distance of 1á5 km.

Snow¯ow deposits

Snow¯ow avalanches are capable of transportinglarge amounts of rock debris, including huge

Fig. 24. Multiple tongues of debris-rich slush¯ows on a colluvial fan inSkjerdingsdalen, late spring 1988.Note that the narrow tracks ofsnow¯ows to the left have cobblyleveÂes, one-clast thick, but showlittle debris deposition at the termi-ni; the snow¯ows were nonerosive(vegetation not removed), carriedlittle debris and formed the leveÂesmainly by shouldering aside thepre-existing slope debris. The largeblocks and smaller debris scatteredin the forefront are attributed tolarge powder snow¯ows.

Fig. 25. Features of modern snow¯ow deposits. (A±B)Longitudinal grooves, interpreted as tool-marks formedby the dragging of large, angular clasts in snow¯owtraction; note the `rolling' a(t) fabric of scattered pebblesand cobbles. (C) Upslope view of a `patchy' lobe ofcoarse gravel, one-clast thick, showing clast imbricationand both a(t) and a(p) fabric; the `mixed' fabric indicatesclast reorientation by subsequent snow¯ow avalanches.(D) Close-up view of a(p)a(i) clast fabric (depicted bythe notebook, 20-cm long) in a `patchy' gravel lobe de-posited by dense snow¯ow. (E) Large boulders and ®nerdebris concentrated at the surface of a melting lobe ofdense snow¯ow; the snow deposit at this stage is still c.4á5 m thick. (F) Rock debris and humic soil material richin uprooted plants, at the surface of a melting lobe ofdense snow¯ow, which is still c. 4 m thick here.



boulders. Some snow¯ows bear little or no debris,as is the case with many powder snow¯ows in themid-winter time, whereas others are very rich in

clastic material (Figs 23B, 25E & F). The contentof debris in a single ¯ow may be extremelyvariable (Fig. 22D; Luckman, 1971). Snow¯ows



transport debris that has accumulated on thesnowpack due to rockfalls and related processes,including wind-blown ®nes; debris that has beenremoved from the mountain slope/ravine andincorporated en route by the ¯ow; and debris thathas been swept by the ¯ow from the apical part ofa colluvial fan.

Dense snow¯ows have high shear strength, andtheir `rigid plugs' can support rock clasts as largeas boulders. Non-turbulent powder snow¯owscan carry cobbles in a similar way. These densersnow¯ows thus carry rock debris in much thesame way as debris¯ows do, with the importantdifference that the snow `matrix' here melts outshortly after deposition and all the debris settlesto the ground. The turbulent powder snow¯owscarry debris similarly to rapid turbidity currents:granules, sand and ®ner grains are carried inturbulent suspension, and the coarser debris inbedload traction. On a steep and relativelysmooth or snow-covered slope, the bedload mayeasily include cobbles and boulders.

The sedimentary deposits of snow¯ows, asobserved on modern colluvial slopes, range fromblankets of scattered, unsegregated debris toirregular, `patchy' lobes of unsorted debris, nomore than one-cobble or -boulder thick (Figs 6, 24and 25C). The clasts are occasionally foundresting upon one another in precarious positions,due to the passive settling from the melting snowmass. The debris is surrounded by a blanket ofwaterlain sand (Fig. 24), usually rich in ®negravel, but is seldom fully illuviated or buried(Fig. 26B). The sand blanket, derived by meltwa-ter runoff, has low-relief internal scours andcommonly shows planar strati®cation and ripplecross-lamination, even within the larger interstic-es of the buried gravel patches (Fig. 26A). Whenemplaced on a subaqueous slope and soon buriedby a debris¯ow, the `patchy' debris lobe mayretain its openwork texture (Fig. 26C). Otherwise,it is ®lled with silt and ®ne sand from nearshoresuspension (Fig. 15B), occasionally bearing faunashells.

The deposits of successive snow¯ows, whenemplaced directly after one another, are amal-gamated and the indistinct boundaries render theindividual avalanche events dif®cult to recognizein an outcrop section. Although the discontinu-ous horizons of large clasts within a thickerblanket of waterlain sand (Fig. 6) almost certainlyrepresent separate avalanches, the actual numberof snow¯ow events recorded by such a compositeunit is likely to be larger, because not everysnow¯ow might necessarily carry coarse debris or

the large clasts may be few and unexposed in aparticular section. A sedimentary unit little morethan one-cobble thick may represent 5±15 con-secutive snow¯ows.

Fig. 26. Details of snow¯ow deposits in outcrop sec-tions. (A) Bouldery, clast-supported gravel whose largeinterstices have been ®lled with ripple cross-laminatedsand and poorer-strati®ed granule gravel; the knife is17 cm. (B) Debris deposited by a larger number ofsnow¯ows in an intertidal backbeach zone, illuviatedwith ®nes and heavily shattered by frost action (notethe open fractures in the foliated gneiss fragments);outcrop of the Younger Dryas palaeobeach in a collu-vial-fan delta; the knife (encircled) is 17 cm. (C)Openwork gravel deposited by snow¯ows that havedescended onto the subaqueous slope of a fan delta; theruler's arms are 23 cm each. Photographs from thecolluvial successions at Gardvik and Fiksdalstrand(localities 1 & 11 in Fig. 1).



The clast fabric of snow¯ow deposits varies ona local scale, and is disorderly when measuredmore systematically. Boulders and cobbles de-posited from turbulent powder snow¯ows mayoriginally have tractional a(t) fabric (Fig. 27A),but the scattered clasts on a steep and sand-®lledsubstratum are highly vulnerable to reorientationby the subsequent avalanches and meltwater ¯ow(Fig. 25C & D). The nonturbulent dense snow-¯ows and slush¯ows may create an internal a(p)fabric of clasts, due to laminar shear, but thefabric loses order when the debris melts out andsettles to the ground (Fig. 25E & F).

Among the characteristic features formed bysnow¯ows on colluvial slopes are low-relieflongitudinal grooves (Fig. 25A & B) and ribs(Fig. 27A). The occasional presence of a cobbleor boulder stuck at the downslope end of a groovesuggests that these furrows are tool-marks formedby the snow¯ow's dragging of large angular clasts.The ribs are due to a linear accumulation ofdebris, rather than erosion (Fig. 27A), but their

formation mechanism is unclear. These longitu-dinal ribs are merely one-boulder or -cobble thickand their width/height ratio is in the range of 2±5.Where coupled, they may be leveÂes made of thesur®cial debris that has been shouldered aside bya dense, highly elongate snow¯ow (Figs 13A and24). Where multiple (Fig. 27A), they may still beleveÂes of a series of such snow¯ows, or mayre¯ect some peculiar, helicoidal pattern of sec-ondary ¯ow in the turbulent boundary layer of alarge powder snow¯ow (cf. Allen, 1984, chapt. 1).Snow¯ows in western Norway often alternatewith heavy rainfalls, and it cannot be precludedthat the longitudinal ribs are due to strong runoff,with the sheet¯ow power maximized by the steepslope and a frozen or water-logged substratum.

Snow¯ows occasionally leave `debris horns' onthe upslope sides of large, immobile obstacles(Fig. 6), which is attributed to the local plastic`freezing' of a dense snow¯ow rich in rock debris.A disorderly `meltout' fabric and an openworktexture with waterlain basal sandy in®ll help to

Fig. 27. (A) `Patchy' gravel lobe deposited by a turbulent, powder snow¯ow on the grass-covered surface of acolluvial fan's ¯ank; Sunnylvsmoldskreddalen (locality 8 in Fig. 1), spring 1993. Note the `rolling' fabric of largeclasts, the longitudinal `debris ribs', one-clast thick, and the overall scatter of debris. The slope's height is c. 60 m. (B)`Debris shadow' on the lee side of large block, left by an erosive slush¯ow; this short ridge is a relict of the upper partof a pre-existing debris¯ow deposit. (A large, bouldery debris¯ow descended the adjoining, till-covered mountainslope of c. 40° during a heavy rainfall in the early spring 1991, and is thought to have been followed by a turbulentslush¯ow derived from the higher slope; the latter avalanche scoured the deposit of its predecessor.) Fiksdal (locality10 in Fig. 1); photograph taken a few days after the event.



distinguish these `debris horns' from the similarfeatures formed by debris¯ows (see earlier sec-tion).

Snow¯ows may also form `debris shadows' onthe downslope sides of large boulders (Fig. 6).Some of these features are depositional, compris-ing coarse debris that has apparently beendropped from a turbulent, debris-laden snowavalanche in the ¯ow-separation zone on theobstacle's lee side. Other `debris shadows' areclearly erosional, representing the lee-side relictsof a sur®cial layer of collvium that has elsewherebeen wiped out by an erosive snow¯ow (Fig. 27B).The obstacle in either case must have been deeplygrounded, or its height was much less than 1/10 ofthe ¯ow thickness (see previous section).

The extreme hydraulic jump at the foot of a cliffmay cause a virtual crash-landing of snow¯owavalanches, with the snow mass splashing outand ejecting substratum debris in a blast-likefashion. A result of the powerful impacts is acharacteristic impact crater, referred to also as`plunge pool/pit' (Liestùl, 1974; Corner, 1980).These oval depressions are 1±5 m deep and 20±100 m in diameter, are often ®lled with meltwater(pond) and have a distinct outer rim of debrisejecta (Fig. 28). No such craters are found to berelated to the other avalanche types, althoughtheir impact on the foot-zone substratum, com-bined with the powerful thrust, can be enormous.For example, a georadar (GPR) survey in Roms-dalen indicates that some large debris¯ow ava-lanches have deformed the postglacial valley-¯oor alluvium to a depth of c. 30 m, forminghydroplastic folds, thrusts and sur®cial `pressureridges'.

Some illustrative examples of colluvial succes-sions rich in snow¯ow deposits are shown inFigs 20 and 29. Importantly, these deposits oftencontain plant fragments (Fig. 25D, E & F), andalso their waterlain in®ll commonly includesaccumulations of humic soil rich in plant detri-tus, derived by contemporaneous slopewash.This plant material and the marine shells fromsubaqueous facies have been used for the radio-carbon dating of the colluvial deposits.

Water¯ow processes

Waterlain deposits are of minor volumetric im-portance in most of the colluvial systems, whoseslope catchments are typically small. However,the role played by water¯ow is often quitesigni®cant, as is shown particularly well by thecolluvial-fan deltas.

The ¯ow of water, whether due to snow-melt orrainfall, occurs generally in two modes: as ashallow, uncon®ned or poorly con®ned sheet-¯ow, and as a channelized stream¯ow. Sheet¯owwashes the colluvial slope and hence is oftenreferred to as `sheetwash' by geomorphologists.The ¯owing water winnows mud and sand fromthe upper slope and deposits this sedimentfurther downslope, mainly by percolatingthrough the coarse sur®cial colluvium and ®llingits interstices. Sheet¯ow is more powerful whenthe sur®cial colluvium is frozen or water-logged.The shallow water¯ow in such conditions issupercritical, capable of transporting pebblesand dislodging isolated larger clasts. In short,sheet¯ow transfers the ®ner sur®cial sedimentfrom the fan apex to the lower slope, alters the

Fig. 28. An impact crater formed bysnow¯ow avalanches that havecrash-landed at the foot of a verysteep colluvial slope; Valldalen(west of locality 15 in Fig. 1), Au-gust 1996. Note the crescentic rim ofejected debris at the outer margin ofthe crater pond (c. 25 m in diame-ter).



primary texture of the sur®cial colluvium andtends to bury the openwork or scattered debris ofpreceding rockfalls and snow¯ows (Fig. 29). Ananalogous process, more often called `slopewash',occurs on the adjoining mountain slope, wherethe ¯owing water removes ®ne debris, includingcontemporaneous soil material, and deposits it onthe colluvial slope below.

Stream¯ow is a more pronounced mode ofwater¯ow (Fig. 6). There is seldom more than onechannel active on a colluvial fan (Figs 3 and 5D),and the channel shifts due to in®lling andavulsion, rather than lateral migration. Thesechannels are typically narrow gullies, developedfrom the ruts of earlier avalanches, whose trackshave been overtaken and modi®ed by water-¯ow. The channels are up to 1±1á5 m deep, haveV-shaped cross-sections and generally lack lev-eÂes. Some channels have leveÂes formed by earlierand/or contemporaneous debris¯ows, but theonly waterlain overbank deposits are thin splaysof sand and ®ne gravel. The modern channels aremainly ephemeral, carrying the seasonal meltwa-ter discharge and the runoff from heavy autumnrains. The isolated palaeochannels found in thecolluvial successions suggest that similar condi-tions occurred also in the past.

The water¯ow in the gullies is often slushy,carrying heavy sediment load and rapidly losingcompetence in the middle to lower reaches. Thesediment concentration is high, and the water mayalso bear slush at the beginning of a snow-meltseason. A `hyperconcentrated', high-density wa-ter¯ow (Fig. 21) is thought to be characterized byhigh apparent viscosity and a rheological behav-iour intermediate between plastic and Newtonian¯uid ¯ow (Beverage & Culbertson, 1964; Nemec &MuszynÄski, 1982; Rickenmann, 1990, 1991); thiswould mean a non-Newtonian pseudoplastic¯uid, which has no limiting plastic strength, butwhose viscosity coef®cient increases rapidly withthe decreasing turbulent shear stresses. The tur-bulence in such conditions tends to be suppressedand the deposition of debris occurs mainly byrapid dumping, with little or no tractional segre-gation.

Where the slope catchment involves a glacier orsemipermanent snowpack and the water dis-charge is relatively high and more perennial, thestream¯ow on a colluvial fan assumes the form ofa relatively wide and shallow channel, oftenbraided (Fig. 30). These channels may be a fewmetres wide, less than 0á5 m deep, and showmedial gravel bars. The ¯ow is rapid and thesediment transport is mainly tractional. The bars

Fig. 29. Detailed log from a valley-side colluvial aprondominated by snow¯ow avalanches, Korsbrekke (lo-cality 9 in Fig. 1). The colluvial system has prograded,obliquely, onto the submarine front of a Gilbert-typedelta formed by the valley-axis river system, supplyingdebris to the delta slope and forming the delta topset.Note the radiocarbon dates. Texture code CS and OWas in the caption to Fig. 9; MS= matrix-supported tex-ture.



are small, no more than 15±20 cm in relief, lackslipfaces and emerge as rhomboidal islands at lowstages. They form mainly as pebbly `shadows' onthe lee sides of cobble/boulder clusters, but thereare seldom more than 1 or 2 such bars per channelwidth. The stream is very effective in transferringsand and ®ne gravel to the fan toe. In the case of afan delta, the stream builds a secondary deltaiclobe, or underwater mouth-bar, on the subaque-ous slope. Conditions of such pronouncedstream¯ow activity occurred periodically in thepast, when many of the colluvial-fan deltasturned into stream-dominated Gilbertian systems(see subsequent section).

Stream¯ow has a natural tendency to scour andtransfer sediment to the lowest part of a fan, butthis does not necessarily mean that the process ispurely `destructive'. Stream channels often playthe role of a distributary that causes markedprogradation of the depositional system. Thechannels on colluvial fans shift less frequentlythan on common alluvial fans, and hence thebuild-out is more localized.

Deposits. The channel-®ll deposits in the col-luvial successions are bedded gravels with inter-layers of strati®ed sand, very coarse to mediumgrained. The beds are 10±30 cm thick, pebbly tocobbly, with a clast-supported texture and spo-radic boulders. Some beds are crudely strati®ed,but most are massive and normally graded(Fig. 31). Most of the gullies are thought to havebeen ®lled by episodes of high-density stream-¯ow, commonly followed by the tractional depo-sition of strati®ed sand and ®ne gravel (Fig. 31).The smaller gullies often show a simple ®ning-upward in®ll, with little or no internal bedding.Many gullies have apparently been plugged bydebris¯ow and slush¯ow avalanches.

The wider and shallower colluvial channelsshow tractional clast fabric and gravel segregationinto lenses (low-relief transient bars). Thesechannel-®ll deposits show internal scours andgreater textural variation, with common planarstrati®cation and isolated sets of low-angle cross-strata.

Debris creep

The knowledge on colluvial creep processes hasadvanced quite recently. For example, Kirkby(1967, p. 359), only three decades ago, felt stillcompelled to assert that `the role of creep in slopeevolution is largely a matter of speculation ¼,unsupported by evidence'.

Debris creep, also referred to as `soil creep' bygeomorphologists and engineers, is a very slowgravitational movement of the sur®cial mantle ofslope debris, which involves no basal detachmentplane of sliding and is perceptible only overrelatively long periods of time (Bryan, 1922;Sharpe, 1938; Kirkby, 1967). Creep is a cumula-tive effect of the innumerable tiny displacements,or discrete interparticle slips, driven by thedistributed load stress of the colluvial mantle.The percolation of water through the debrismantle, the heaving of clasts by frost action andthe biological activity of plant-roots and animalsare the most important physical agents of collu-vial creep (Schumm, 1956; Kirkby, 1967; Emery,

Fig. 30. Shallow braided streams on a colluvial-fandelta at SkaÊr (A) and a vegetated colluvial fan inRomsdalen (B). Photographs taken in August 1996, af-ter the main phase of meltwater discharge.



1978; Finlayson, 1985; Feda, 1992). The interpar-ticle slips occur at random and the creep processis highly incremental, with discrete and relativelybrief episodes of activity. Longer-term measure-ments indicate rates ranging from a few mil-limetres to a few centimetres per year, dependingon climate and material type (Fleming & Johnson,1975; Young, 1978a,b; Selby, 1982; Goudie et al.,

1985; Jahn, 1989; Lautridou et al., 1992; Auzet &Ambroise, 1996). In warm arid regions, signi®cantcreep may occur only after a series of heavyrainfalls have saturated the colluvium (Schumm,1956). In humid regions, creep is more commonand may reach speeds as high as 20 cm dayÿ1

(Radbruch-Hall, 1978).Although commonly de®ned as `a very slow,

more or less continuous ¯ow' (Bolt et al., 1975;Emery, 1978), creep in reality is highly discon-tinuous, and is not a ¯ow in the strict terms oflow-velocity friction dynamics (Baumbergeret al., 1994). Creep, even when driven at aconstant rate, involves a multitude of interparti-cle `stick' and `slip' states, or localized jerkymovements, due to the build-up and randomrelease of interparticle frictional stresses. Creepis thus characterized by marked velocity ¯uctu-ations, which disappear when the strain rateincreases and a slow frictional ¯ow begins. Thehigher strain rate eliminates the stick-slip modeof particle movements and renders them increas-ingly continual, although the transition fromcreep to slow inertial ¯ow is hardly discernableto an observer's eye, even when the ¯ow rate isof the order of 1 lm s±1 (or c. 30 m yr±1).Further, the kinetic friction coef®cient (de®nedas the tangential stress normalized to the mate-rial weight) in creep is `rate-weakening', whichmeans decreasing with the increasing strain rate,typically in a logarithmic fashion. In a frictional¯ow, the coef®cient is `rate-strengthening', in-creasing with the strain rate, no matter if the¯ow rate is high (McTigue, 1982; Haff, 1983;Campbell, 1990) or very low (Baumberger et al.,1994).

In rheological terms, creep is considered sim-ply as an increase of permanent strain with timeunder a constant stress (Lubahn & Felgar, 1961;Coates, 1967). Interestingly, the strain in creepmay be episodically reversible (Kirkby, 1967;Finlayson, 1981; Auzet & Ambroise, 1996), hencethe constitutive models include an elastic com-ponent (Emery, 1978, ®g. 3). It is now widelyagreed that the creep of clastic materials is anonlinear, elasticoviscous Maxwell-type phe-nomenon, whereas their ¯ow is a nonlinear orlinear, plasticoviscous Coulomb-type phenome-non. In contrast to mass¯ow, there is also no truesteady-state creep of clastic materials (Emery,1978). Under a constant deviatoric stress, thestrain rate either declines or increases exponen-tially with time. The increase often leads to enmasse sliding failure (Emery, 1978, ®g. 1) andpossibly mass¯ow (Holmes, 1978).

Fig. 31. Detailed log from a colluvial fan formed byalternating debris¯ow and stream¯ow processes; Bars-taddalen (locality 4 in Fig. 1). The mountain slope hereis 23±35° and the channel of the feeder stream, Tver-relva, has an inclination of 20±24°, but the depositionalsurface of the fan is inclined at 13° (apex) to 10° (toe),which renders it transitional to a common alluvial fan.



The velocity pro®le of a creeping debris layer,as predicted by Davison's (1889) theory andcon®rmed by ®eld studies (Auzet & Ambroise,1996), is characteristically concave, with thedownslope velocity component declining expo-nentially with depth, but dependent upon thesediment properties (represented by the Davison`soil constant'). Kirkby's (1967) theory predictsthe exponential velocity decline to occur below asur®cial rigid-plug zone, which means maximumshear strain at some depth, rather than near thesurface. This condition would more readily pro-mote sliding failure.

When driven by frost and thaw action, thecolluvial creep is called soli¯uction (or `soil¯ow'). Many authors distinguish `soli¯uction'from `soil creep', although there is little physicaldifference, apart from the sawtooth pattern ofclast movement due to the frost-heave and thaw-collapse of larger clasts (Washburn, 1979). In fact,the concept and ®rst quantitative theory ofcolluvial creep were introduced, by Davison(1889), in the context of freeze-thaw cycles.However, soli¯uction is due to speci®c condi-tions and is the most pronounced variety ofcolluvial creep, involving a good deal of slowfrictional ¯owage. Soli¯uction contributes to thesmooth, `¯owing' pro®le of periglacial colluvialslopes and has been a subject of extensive ®eldstudies (Frenzel et al., 1993). The studies show,for example, that a sur®cial colluvial mantle cancreep over a frozen substratum on slopes as gentleas 2±3°.

The colluvium in permafrost regions is oftenpermanently cemented with ice, except for ashallow summer thaw. With the continuousreplenishment by debris from upper slope, theincreasing load stress may cause the interstitialice to yield and ¯ow in a plastic fashion. A rock-glacier then creeps down the slope and spreadsslowly in the foot zone (Giardino et al., 1987;Barsch, 1988; Whalley & Martin, 1992; Leliwa-Kopystynski & Maeno, 1993).

Deposits. A pronounced colluvial creep, suchas soli¯uction, involves lobes or broader sheets ofsur®cial debris, ranging from c. 1 m to 50±80 min width. Colluvial deposits attributed to sol-i¯uction include well-bedded to poorly beddedgravels, ®ne pebbly to cobbly, as well as somecrudely bedded diamictons (Bertran et al., 1995).This wide range of deposits suggests that theircharacteristics may be far more related to theprimary slope material, or the local mode ofdebris production, than to the creep processitself. In short, the products of creep in ancient

colluvial successions are dif®cult to recognizedue to the lack of diagnostic criteria, or anestablished concept for the sedimentary signa-tures of creep.

An exception may be the concept of `stone-banked' soli¯uction lobes and sheets (Bertranet al., 1994, 1995; Van Steijn et al., 1995), invokedfor the alternating layers of openwork coarse-pebble gravel and matrix-supported ®ner gravel ofgreÁzes-liteÂes type. In this model, a lobe or sheetof sur®cial debris creeps downslope, at a rate of2±5 cm to 20±30 cm yr±1. Frost action makes thecoarsest clasts reach the surface, where theyfurther creep downslope, due to needle-icegrowth, at a rate of up to 1 m yr±1. This coarsedebris accumulates at the front of the lobe orsheet, and the resulting `stone bank' is progres-sively buried, in a caterpillar fashion, by theslow-creeping parental layer of matrix-rich de-bris. As the process continues and the front isreplenished with coarse clasts, an undulatorylayer of coarse openwork gravel forms and iscovered by a ¯at-topped layer of ®ner, matrix-richgravel. The stony pavement that creeps over thetop may develop an upward coarsening due to thefrost segregation of clast sizes, whereas theprogressive burial of the frontal `bank' results inan upward ®ning, with a matrix-rich upper part.When signi®cant illuviation occurs, the layer maybe covered with a thin sand or silty mud capping.The successive lobes or sheets lead to a repetitionof this bedding motif, with the bed geometryranging from lenticular (lobes) to more tabular(sheets) in outcrop sections parallel to the slopestrike.

However, the model should be applied withcaution if used as a possible guide for therecognition of soli¯uctional deposits in colluvialsuccessions. Firstly, the `matrix-rich' layers in thereported type sections are mainly clast-support-ed, though matrix-®lled (Bertran et al., 1994,1995; Van Steijn et al., 1995), which renders thebedded gravel similar to some other colluvialfacies, for which a soli¯uctional origin is unlikely(see deposits of low-viscosity debris¯ows earlierin the text) or can outrightly be precluded (seecolluvium D of Nemec & Kazanc�, 1999). Second-ly, many soli¯uctional deposits reported in theliterature are quite unlike those implied by the`stone-bank' model (Flint, 1971; Washburn, 1979;Frenzel et al., 1993). Thirdly, soli¯uctional de-posits are widely reported to have an a(p) fabric(Nelson, 1985; Bertran et al., 1997), which maynot be the case with the `stone banks' (Van Steijnet al., 1995).



The style of creep-related deformation on acolluvial slope may vary from extensional, with apronounced downslope spreading of the debrismantle (as lobes or sheets) and possible formationof transverse `gulls' (open cracks), to distinctlycompressional, with a sur®cial buckling or morelocalized bulging of the mantle. The creep ofcolluvium is liable to sudden acceleration after anepisode of heavy rainfall or snow melting, whenthe bulk weight of the mantle increases andinterparticle friction is reduced. Many debris-¯ows are thought to be generated in this way onmountain slopes (Holmes, 1978; Radbruch-Hall,1978). Likewise, the deposits of colluvial slidesand debris¯ows often tend to creep after theemplacement (Bolt et al., 1975; Van Asch, 1984),and this secondary creep may strengthen the a(p)clast fabric of a debris¯ow deposit.

In short, most colluvial slopes are probablysubject to creep processes, and there is littledoubt that these processes have played a role onthe colluvial slopes in western Norway. Creepwas probably important in triggering debris¯ows,or in displacing the fragmented bedrock intounstable positions and causing rockfalls (Haefeli,1966; Hausen, 1971; Wyroll, 1977). However, anydirect evidence of creep processes in a colluvialsystem dominated by avalanches is rather dif®-cult to recognize.

SEDIMENTARY SUCCESSIONS

This discussion of colluvial successions is fo-cused on coastal depositional systems, which

include both subaerial and subaqueous facies andthus bear also the record of relative sea-levelchanges (Fig. 32). These depositional systems arecolluvial-fan deltas (Blikra & Nemec, 1993b):coastal accumulations of land-derived sedimentformed by the progradation of colluvial fansdirectly into the sea (Figs 3 and 5D). The steepmorphometry and the predominance of avalancheprocesses, from the subaerial apex to the sub-aqueous toe, render these gravelly systems aspeci®c end-member category of the spectrum ofdeltas observed in nature (Nemec & Steel, 1988;Nemec, 1990a; Reading & Collinson, 1996).

The colluvial-fan deltas have relatively largesubaqueous parts and small, short subaerialapices (Fig. 32), which pass upslope into distinctavalanche tracks, often stemming from bedrockgullies or mountain ravines. The deposits areslope-waste gravels, poorly sorted and oftenpoorly bedded, as is typical of a colluvium.However, much of the material here has beenderived by the resedimentation of glacial till andpossibly kame terraces, which renders the col-luvium somewhat unusual: the gravel is com-monly subrounded to rounded, and does notnecessarily correspond to the local bedrock.

The colluvial successions at ®rst glance look`chaotic' and very poorly organized. However,they are clearly heterogeneous, and a closeexamination of the stratigraphic sections almostinvariably reveals distinct changes in the compo-nent facies assemblages (depositional processes)and their bedding architecture (fan-delta physio-graphy).

Fig. 32. The depositional setting ofcolluvial-fan deltas at a steep fjordmargin affected by the postglacialregional isostatic uplift; schematicdiagram pertaining to the study areain western Norway.



The ensuing review focuses on three fan-deltatypes, which are representative of the generalrange of colluvial systems in the region. Theyshow the varying role of different avalancheprocesses, re¯ecting the variable responses ofthe local mountain slopes to the postglacialclimate and its changes. The emphasis is on thefacies anatomy and depositional history of thecolluvial systems, and the developmental `stages'distinguished in the local successions arechronostratigraphically correlative, as indicatedby radiocarbon data.

Debris¯ow-dominated colluvial systems

An example of this type of colluvial system is thearray of coalescent colluvial fans that haveprograded into the sea at the western margin ofTomrefjorden (locality 12 in Fig. 1). The adjoin-ing mountain slope has a thick glacial-till mantleresting on gneissic bedrock. The fans are steepaccumulations of coarse debris at the termini ofdistinct avalanche tracks scoured in the tillmantle. The tracks reach an altitude of c. 600 m,and the main source zone for the colluvium is atthe height of 200±600 m, where the mountainslope is steeper than 30° and its till cover hasbeen markedly depleted. The slope has a lowpotential for snow-drift accumulation under theprevailing pattern of westerly winds. Snowavalanches here are very rare and have played anegligible role in the postglacial sedimentation.

Our analysis of this coastal apron, nowemerged, is based on one of its component fandeltas, well-exposed in a large gravel pit (Figs. 19,33 and 34). The colluvial facies assemblages andtheir depositional architecture indicate ®ve mainstages of the fan-delta development (Fig. 35).

Stage 1. The oldest part of the sedimentarysuccession, or the fan delta's `core' (Fig. 35),consists of bouldery to cobbly debris¯ow depositsrich in ®ne-grained matrix (log 4 in Fig. 33). Theyrepresent mainly high-viscosity debris¯ows,some apparently accompanied by a watery, low-viscosity trailing ¯ow. The trailing mass¯ows,carrying pebble gravel and coarse sand, wereturbulent, capable of accreting backsets of up-slope-dipping cross-strata onto the debris¯owlobes. This trailing ¯ow was probably generatedwhen the debris¯ow plunged into the water onthe steep fjord-margin slope. The debris¯ow bedsoften show normal grading at the top (Fig. 14A),which suggests subaqueous emplacement withstrong shearing against the ambient water. Un-graded debris¯ow beds tend to be amalgamated,

due to the lack of visible textural differencesbetween two or more successive deposits. Somebeds are separated by thin interlayers of well-sorted, ®ne-grained sand, which show normalgrading and are thought to have been depositedfrom wave-derived nearshore suspension. Thewave base in Norwegian fjords is seldom deeperthan 2 m, and the steep nearshore slopes tend tobe draped with storm-derived suspension ®nes.The sand layers are often strongly deformed byloading.

This earliest colluvial sedimentation probablyfollowed directly the deglaciation of the area(c. 12 700 years BP), when the till-covered, non-vegetated mountain slope was highly unstable.The regional data on postglacial relative sea levels(Mangerud et al., 1979; Svendsen & Mangerud,1987, 1990) and the stratigraphic position of thisfacies assemblage beneath a younger Gilbert-typedelta (Figs 33 and 35) indicate that the depositsmust be subaqueous. The sedimentary faciessupport this interpretation.

The colluvial system at this early stage wasapparently a steep cone, whose depositionalslope extended directly into the deep water, withlittle or no change in gradient (Fig. 35). Manymodern fjord-margin systems have such a physio-graphy. The debris¯ows, derived from the slopetill, were thus emplaced mainly subaqueously.The relative sea level at this stage was high(Fig. 35), and the subaerial apex of the cone wasprobably small and might not even be preserved.If preserved, it would now be buried beneath theyounger colluvium at an altitude higher than thegravel-pit outcrop.

Stage 2. The overlying facies assemblageconsists of moderately thick beds of clast-tomatrix-supported, pebble to cobble gravel(Figs 14B and 15), intercalated with thinnerbeds of ®ne-pebble gravel and pebbly sand,mainly normally graded or graded-strati®ed(logs in Fig. 33). Their deposition is attributedto subaqueous debris¯ows and high-densityturbidity currents, respectively. These steeper-inclined deposits represent the foreset of aGilbert-type fan delta that has downlapped thepre-existing subaqueous cone. The foreset de-posits show extensive internal truncations andalso local angular disconformities marked bydownlap bedding (Figs 33 and 34). The erosio-nal features are attributed to the falling of therelative sea level with the onset of the postgla-cial isostatic uplift, whereas the episodes ofdelta-slope downlap indicate lateral shifts of theaxis of sediment dispersal of the fan.



The predominance of clast-supported depositsof low-viscosity debris¯ows and the abundance ofturbidites suggest relatively `wet' conditions onthe adjoining mountain slope, with a greater roleof ¯owing water. There is evidence of streampalaeochannels and mouth-bar cross-strata at the

top of the foreset (Fig. 34). The pre-existingavalanche tracks on the mountain slope wereprobably overtaken by ¯ushy, ephemeral streamscarrying meltwater and/or rainwater, and thepredominance of stream¯ow promoted the devel-opment of a Gilbert-type system. The fan delta

Fig. 33. Outcrop section and detailed logs of the colluvial fan-delta complex at Tomrefjorden (locality 12 in Fig. 1).The cross-section is approximately normal to the palaeoshoreline.



had a steep and short subaerial plain, allowingdebris¯ow avalanches to pass directly onto thesubaqueous slope.

The foreset facies assemblage is internallyvaried, but shows two packages of mainlycoarse-gravel beds separated by a thinner packageof mainly sandy and ®ne-pebbly turbidites(Fig. 33; the younger part is to the lower-left inFig. 34). The bipartition suggests two phases ofslope wastage dominated by watery debris¯ows,separated by a phase dominated by stream¯ow(stages 2A & 2B in Fig. 35). A narrow belt oferosive beach deposits occurs at the top of foresetlobe 2A, covered erosively by the waterlain topsetof lobe 2B (Fig. 35). The altitude of this palaeo-beach (c. 80 m) corresponds to the sea-level standshortly after the deglaciation, which means thatthe sedimentation of stage 1 was relatively briefand was followed directly by development of thedelta lobe 2A. Lobe 2B extends to an altitude of c.

76 m, which suggests that its deposition com-menced c. 12500 years BP.

Stage 3. The colluvial system subsequentlyresumed the simple conical physiography thatcharacterized stage 1, although the relative sealevel had fallen considerably by that time(Fig. 35). The emerged part of the Gilbert-typedelta of stage 2 had been strongly truncated byerosion and overlain by a succession of clast-supported, subaerial debris¯ow deposits (Figs 19and 34). They are poorly bedded and commonlyamalgamated, but the consecutive debris¯owevents are recognizable (Fig. 19). Some gravelbeds show features suggesting deposition fromslushy debris¯ows or dense snow¯ows.

The slope of the colluvial fan at this stageextended `smoothly' into the sea (Fig. 35), with anarrow, aggradational beach zone that was easilybypassed by many avalanches. The palaeobeachunit is a relatively thick succession of coarse,

Fig. 34. Outcrop section of the colluvial fan-delta complex at Tomrefjorden (locality 12 in Fig. 1). Note that the upperright-hand part of the cross-section is parallel to the palaeoshoreline, whereas the rest of the section is roughlynormal to the latter.



clast-supported gravel that has been severelyshattered by frost action, probably in an intertidalbackbeach zone (Blikra & Longva, 1995). Thealtitude of the palaeobeach corresponds to therelative sea-level stand the Younger Dryas time(11 000±10 000 years BP).

Stage 3 indicates a dramatic increase in slope-waste processes, apparently due to the climaticdeterioration of the Younger Dryas time. Thecolder climate, with a decreased and more sea-sonal rainfall, probably reduced slopewash andrendered the moutain slope prone to mass fail-ures. Soli¯uctional creep and the effects of slushand seasonal snow-melt might be important intriggering the debris¯ows.

Stage 4. The sediment yield then considerablydecreased, probably due to reduced snowfall anddepletion of the mountain-slope source by thepreceding phase of intense wastage. This stage ofminimal sedimentation and no record of ava-lanche deposits is attributed to the phase ofregional climatic warming, known as the Ho-locene climatic optimum (c. 10 000±5000 yearsBP). The hiatus is regional.

Stage 5. The intensity of slope-waste processessubsequently increased, although the late Ho-locene colluvium is thin, more local, and duemainly to the redeposition of the older colluvium.The radiocarbon date of soil material separatingthe leveÂes of these younger debris¯ows from thedeposits of stage 3, at the margin of an avalanchetrack, indicates avalanches younger thanc. 2800 years BP. This suggests a relatively longbreak between the Younger Dryas sedimentationof stage 3 and the onset of the Holocene sedi-mentation. A few gullies have more recently beenincised in the apron surface. The streams carrymainly the spring meltwater runoff from the fan

apices to the shoreline, where underwater mouth-bars have formed on the steep subaqueous slope(Fig. 35).

The relative sea level is known to have fallenstrongly after the Younger Dryas time, and thisfactor is likely re¯ected in the redeposition ofcolluvium and the incision of stream gullies.Stream¯ow processes suggest an increase inseasonal runoff. Some of the gullies have beenplugged with debris¯ow and/or slush¯ow depos-its and subsequently rejuvenated by water¯ow,with the avalanche deposits preserved as leveÂes.The plugging of channels probably re¯ects ayounger phase of climatic deterioration, whenthe gullies became pathways for mass¯ow ava-lanches. The radiocarbon date of a palaeosolburied beneath one of the leveÂes indicates de-bris¯ows younger than c. 1500 years BP. Severaldebris¯ows have also descended the colluvialslope along the stream gullies in the historicaltimes (Blikra, 1994).

`Mixed-type' colluvial systems

The two colluvial-fan deltas near Gardvik (Fig. 3),at the eastern margin of érstafjorden, afford anexample of colluvial systems formed by alternat-ing debris¯ow and snow¯ow avalanches. Thefans are steep accumulations of coarse debris atthe termini of two distinct avalanche tracts thatcan be traced to an altitude of c. 900 m on themountain slope (Fig. 3, top). As in the previouscase, the gneissic bedrock here is covered with athick glacial till, but the mountain slope faces thesouth-west and has a relatively high potential toaccumulate snow, drifted by winds from thenorth and north-east. The fan surfaces show clearevidence of recent snow avalanches, including

Fig. 35. Stratigraphic anatomy of thecolluvial fan-delta complex at Tom-refjorden. Schematic diagram; theframe's length is c. 850 m andheight c. 200 m; the total relativesea-level fall (1±5) is c. 80 m.



damaged trees/bushes and the occurrence ofscattered bouldery gravel and `patchy' debrislobes on grass-covered slopes. There are alsosome relatively fresh gravel lobes deposited bywatery or slushy debris¯ows (Fig. 3, bottom).

The two fan deltas are very similar, althoughthe stratigraphic analysis is based on the northernone, well-exposed in a large gravel pit (Figs 3 and36). The sedimentary facies assemblages and theirdepositional architecture indicate the followingdevelopmental stages (Fig. 37), correlative withthose of the Tomrefjord fan-delta complex(Fig. 35).

Stage 1. The oldest part of the colluviumconsists of bouldery to cobbly, matrix-supporteddebris¯ow deposits, similar to those in theTomrefjord fan delta (previous section). Theyrepresent the subaqueous part of a steep colluvialcone, which probably had a very small subaerialapex (unexposed or unpreserved) and was built

chie¯y by high-viscosity debris¯ows that de-scended from the mountain slope directly intothe fjord. The oldest radiocarbon date frommollusc shells found at the stratigraphic top ofthis steeply bedded facies assemblage is12400 years BP, which indicates deposition di-rectly after the deglaciation. This early stagewould thus be coeval with stage 1 of the Tomre-fjord fan delta.

Stage 2. Deposits correlative with stage 2 of thelatter fan delta are negligible here. A thin layer ofshell-bearing, well-sorted ®ne sand, dated to c.12300 years BP, indicates a period of markeddecrease in slope-waste processes. Scattered cob-bles and boulders with pebbly sand `patches'(thin lenses) suggest sporadic snow avalanches,possibly powder snow¯ows, reaching the sub-aqueous slope. Some of the coarse debris could bederived by snow¯ow avalanches and further shedfrom a shoreline ice-foot. The marine sand layer iscovered by a heterogeneous unit of chaotic,cobbly to bouldery gravel, whose clasts showclayey coatings and the interstices are partly ®lledwith ®ne, silty sand. This composite unit suggestsdeposition from a series of snow¯ow avalanchesslightly younger than 12300 years BP. The pausein colluvial sedimentation was probably due tothe temporal depletion of the slope source by thepreceding phase of mass wastage and possibly alocal climatic effect (e.g. change of prevalentwind direction).

Stage 3. The overlying succession of cobbly tobouldery gravel beds (Figs 36 and 37) indicates an

Fig. 36. Detailed log from the colluvial-fan delta atGardvik (from outcrop section C in Fig. 3, bottom).Texture code CS and OW as in the captions to Fig. 9 &29. The succession corresponds to stage 3 in Fig. 37.

Fig. 37. Stratigraphic anatomy of the colluvial-fan deltaat Gardvik (Fig. 3). Schematic diagram; the frame'slength is c. 450 m and height c. 175 m; the total relativesea-level fall (1±5) is c. 50 m.



alternation of subaqueously emplaced debris-¯ows and snow¯ows. This dramatic increase inslope-waste processes represents the YoungerDryas climatic cooling, as is indicated by theassociated palaeobeach, rich in frost-shatteredgravel, at an altitude of 25±30 m (Figs 26B and37). Some snow¯ows were apparently halted inthe beach zone, but most of the avalanchesbypassed the latter, causing considerable accre-tion and progradation on the subaqueous slope.The subaerial part of this facies assemblage isdominated by bouldery deposits attributed todense snow¯ows (Figs 26A and 36).

Stage 4. No deposits from the ®rst half of theHolocene epoch are found in this colluvialsystem, which indicates a relatively long breakbetween the Younger Dryas sedimentation andthe onset of the Holocene slope-waste processes.This hiatus is regionally correlative.

Stage 5. The overlying subaerial deposits(Fig. 37), covering also the emerged slope of thefan delta, represent mainly debris-laden snow-¯ows. Radiocarbon dates indicate that this stageof deposition commenced not much earlier than4300 years BP (Blikra, 1994). The facies charac-teristics and intervening palaeosol horizons sug-gest several pulses of snow-avalanche activityseparated by periods of slopewash, with theredeposition of contemporaneous soil. The avail-able radiocarbon dates from this facies assem-blage suggest a high incidence of snow avalanch-es shortly after c. 4300 years BP, then fromslightly before 3500 until c. 3200 years BP, andtwo more pulses after 3200 and c. 3100 years BP,respectively.

More recently, a stream channel has beenincised in the emerged colluvial system (Fig. 3).This phase of water¯ow activity is probablyyounger than c. 3000 years BP and indicates anincrease in seasonal rainfall and sur®cial runoff.The gully carries mainly water from the springsnow-melt and heavy autumn rains, causingvigorous redeposition of the colluvium. Thestream¯ow and gully-conveyed avalanches,mainly snow¯ows, have built a small, secondarydelta lobe at sea level (Fig. 37).

Rockfall-dominated colluvial systems

Colluvial fans and aprons dominated by rockfallavalanches are generally the steepest, and aresourced by very steep, weathered bedrock cliffswith little capacity for snow accumulation. Sys-tems of this type are more common in the outercoastal zone (Fig. 1), where the mountain slopes

have lesser altitudes (Fig. 5A). The stratigraphicsections of these colluvial systems are not well-exposed, and also the radiocarbon dates onrockfall deposits are relatively few. Therefore,the stratigraphic model of a rockfall-dominatedfan delta (Fig. 38) is a tentative compilation fromseveral localities, including the colluvial systemsat Midsund and Stormyr, Oterùya (localities 14 inFig. 1). The recognizable depositional stages(Fig. 38) are correlative with those distinguishedin the other fan deltas (Figs 35 and 37).

Stage 1. The oldest part of a rockfall-dominatedfan delta is typically an assemblage of matrix-richgravel beds resting directly on the mountainslope, with a marked backlap (Fig. 38). Theseare mainly deposits of high-viscosity debris¯ows.This early stage of deposition is attributed to thevigorous resedimentation of glacial till almostconcurrently with the deglaciation, for the glaci-genic mantle on a cliff-like slope must have beenextremely prone to gravitational failures.

Stage 2. There is little record of rockfallavalanches at the early postglacial stage, whereasthe overlying colluvial deposits apparently rep-resent the Younger Dryas. A phase of nondepo-sition and weathering is thought to have oc-curred, which is hardly recognizable in asedimentary succession, apart from a possible

Fig. 38. Stratigraphic anatomy of a postglacial colluvi-al-fan delta dominated by rockfall avalanches. Sche-matic compilation from several localities; the frame'slength is c. 360 m and height c. 250 m; the total relativesea-level fall (1±5) is c. 50 m.



drape of marine ®nes on the subaqueous slope ofa fan delta. We have few radiocarbon data toassess the time-span of stage 2 in rockfall-domi-nated systems, but the resedimentation of glacialtill in these cases was probably very early andrelatively fast, and the following hiatus in slope-waste processes was thus likely longer than in theother systems.

Stage 3. The overlying deposits are bouldery tocobbly, openwork rockfall gravels alternatingwith the clast-supported, cobbly to pebbly depos-its of low-viscosity debris¯ows. The lack of anyintervening palaeosols and the occurrence offrost-shattered gravelly palaeobeach facies at thealtitude of Younger Dryas palaeoshoreline indi-cate that this stage of slope wastage represents thewell-known phase of regional climatic cooling,dated elsewhere to c. 11 000±10 000 years BP.

Stage 4. There is no record of rockfall processesfrom the ®rst half of the Holocene epoch, whichindicates another major hiatus in the rockfall-dominated colluvial systems.

Stage 5. The younger colluvium is dominatedby rockfall deposits, commonly separated byinterlayers of humic soil material derived bycontemporaneous slopewash (Fig. 9). The radio-carbon dates from these palaeosols indicate de-position in the second half of the Holoceneepoch. This wedge-shaped facies assemblage isslightly steeper and has a lesser downdip extent,pinching out near the slope break of the YoungerDryas palaeoshoreline (Fig. 38). The limiteddownslope extent probably re¯ects the decreasedvolume and/or free-fall height of rockfall ava-

lanches, due to the source depletion and theprogressive backlap of the mountain cliff. Onlysome exceptionally large rockfalls occasionallyreach the low-lying modern shoreline and cometo rest on the steep underwater slope (Fig. 38).

PALAEOCLIMATIC RECORD

The stratigraphic analysis of the postglacial col-luvium in western Norway has shown that thelocal sedimentary successions bear an importantrecord of climatic changes (Blikra & Nemec,1993a; Blikra, 1994; Blikra & Longva, 1995; Blikra& Nesje, 1997; Blikra & Selvik, 1998). More than100 radiocarbon (14C) dates of humic palaeosols,wood fragments, mollusc shells and organiccarbon-rich mud have been derived from over20 localities. The colluvial successions show aconsistent, regionally correlative record indicat-ing periods of pronounced slope wastage andperiods of relative stagnation. The regional dataare compiled in Fig. 39 in terms of the mainavalanche categories, and are discussed furtherbelow.

Avalanches and climate

The avalanche processes in mountainous terrainsare generally controlled by both climate and thelocal slope conditions (topographic gradient,altitude range, orientation relative to main winds,and debris source type). Colluvial systems arecharacterized by highly episodic sedimentation

Fig. 39. The chronostratigraphy andintensity of the main colluvial pro-cesses, based on a compilation ofdata from more than 20 local suc-cessions, compared with the recordof glacier development and ice-frontoscillations in Sunnmùre area (Lar-sen et al., 1984; Matthews & KarleÂn,1992), Jostedalsbreen area (Nesje &Kvamme, 1991) and Hardangerjùku-len area (Dahl & Nesje, 1994). Theactivity of particular processes hasbeen estimated qualitatively on thebasis of the abundance of their de-posits.



and are thus more likely to respond to majorclimatic changes than to minor ¯uctuations. Forexample, a climatic change must bring suf®cientcooling, or enough rainfall, to be `recognized' bythe dynamics of a colluvial system. Further, thelocal slopes may have different thresholds ofresponse to a particular climatic change, and thisvariation may cause `random noise' in the region-al record. It is therefore very meaningful that theclimatic record shown in the postglacial col-luvium in Norway is legible and regionallycorrelative.

Climate is widely recognized to be the principalcontrol on colluvial sedimentation, with differentprocesses generally related to somewhat differentconditions (Campbell, 1975; Caine, 1980; Rapp &Nyberg, 1981, 1988; Eisbacher & Clague, 1984;Fitzharris & Bakkehùi, 1989; Jibson, 1989; Ko-tarba, 1991; Luckman, 1992; GonzaÂlez DõÂez et al.,1996; Wieczorek & JaÈger, 1996). The incidence ofsnow avalanches is generally related to coldwinters and depends on snowfall intensity, prev-alent wind direction and snow-drift rate. Thisrelationship is thought to be re¯ected, for exam-ple, in the high incidence of snow avalanchesduring the cold Younger Dryas (Blikra & Nemec,1993a; Blikra & Nesje, 1997) and the Little IceAge, AD 1450±1920 (Grove, 1972). Rockfalls, al-though directly related to bedrock weathering, arelikely here to have been triggered mainly by theair temperature variations (freeze-thaw cycles)and the effects of rainstorms and snowpack. Theincidence of debris¯ows might depend on awider range of factors, including local slopeconditions and the more recent anthropogenicactivity (Innes, 1983). Some debris¯ows could betriggered by the impact of rockfalls and snow-¯ows on fan apices. Debris¯ow deposits predom-inate at some localities but are sparse or absent atothers, which suggests an important role of localconditions. However, the main cause of debris-¯ows was the slumping of glacial deposits andupper colluvium, probably due to heavy rainfallor meltwater runoff and thus related mainly toclimate. For example, the increase in debris¯owprocesses at the end of the Little Ice Age has beenattributed to permafrost degradation and thehigher proportion of rainfall in annual precipita-tion (Van Steijn, 1996).

The following discussion puts special empha-sis on the regional record of snow¯ow avalanch-es, because these processes are more closelyrelated to climate and also because the radiocar-bon dates on snow¯ow deposits are more abund-ant.

Snow¯ow record

There is a good regional evidence of abundantsnow avalanches around 12 300±12 000 years BP(Fig. 39), when the climate is thought to havebeen cold, but ¯uctuating (Karpuz & Jansen,1992). This earliest record of snow avalanchesand the frost-shattered coeval palaeobeaches(Blikra & Longva, 1995) lend strong support tothe notion of cold climatic conditions in theOlder Dryas time, contrary to the opinion of someother authors (Broecker, 1992). The majority ofthe colluvial systems further show a distinctpulse of snow¯ow processes corresponding tothe Younger Dryas phase of climatic cooling,c. 11 000±10 000 years BP (Fig. 39). A majorhiatus separates these early postglacial phases ofpronounced snow-avalanche activity from thephase that followed the Holocene climatic opti-mum. A similar hiatus has been recognized in theadjacent Jostedalsbreen area, where the availableradiocarbon dates from the palaeosols and peatlayers separating avalanche deposits from theunderlying glacial till are consistently youngerthan the Holocene optimum (Nesje et al., 1991,1994).

The Holocene regional record of snow ava-lanches indicates a progressive climatic deterio-ration and increasingly colder winter conditionssince at least c. 4700 years BP. Shorter-termclimatic ¯uctuations are indicated by the inter-vening debris¯ow and waterlain deposits, includ-ing local peat layers and common occurrences ofhumic soil redeposited by contemporaneousslopewash. A regional climatic deteriorationsince at least c. 5000 years BP is indicated alsoby the published dates and palynological data onpalaeosols (Matthews & Dresser, 1983; Matthews& Caseldine, 1987).

Some of the local colluvial systems, withcatchments at high altitudes or particularly proneto snow-drift accumulation, show the ®rst evi-dence of Holocene snow avalanches as early asc. 7000 years BP. Other localities show the onsetof Holocene avalanche sedimentation as late asc. 4700 years BP, with at least two consecutivepulses of pronounced slope wastage. This appar-ent diachroneity of the initiation of slope-wasteprocesses after the Holocene climatic optimum ismeaningful, as it probably re¯ects the variedresponse threshold of the local mountain slopesto the regional climatic deterioration.

An increase in snow avalanches occurredbetween c. 3800 and c. 3000 years BP, althoughthe intervening local palaeosols indicate phases



of nondeposition and slope stability. The subse-quent decline in snow¯ows was followed by atleast three marked pulses between c. 2500 andc. 1800 years BP, separated by phases of mainlywater¯ow processes. The youngest prehistoricalrecord of snow avalanches indicates a regionalpulse around 1400 years BP. Historical evidence(Grove, 1972) and lichenometric data (McCarrol,1993) indicate an increase in snow avalanchesduring the Little Ice Age.

Debris¯ow record

The colluvial successions at some localitiesindicate a pulse of debris¯ows directly after thedeglaciation, around 13 000 years BP, prior to theonset of the Holocene sedimentation c. 10 000years BP (Blikra & Nesje, 1997). However, themain activity of these processes dates to the lateHolocene (Fig. 39). Some mid-Holocene debris-¯ows, dated as younger than c. 5000 years BP,accompanied snow avalanches on several fans orwere roughly coeval with such avalanches onother fans. The association of the two processes,although not a rule, is recognizable more often inthe younger Holocene colluvium. The deteriorat-ing climate was probably seasonal, with sharperwinters, and the increase in snowfall might havecaused snow avalanches in the winter and earlyspring, and debris¯ows during the snow-meltseason and summer/autumn rains (Rapp & Ny-berg, 1981, 1988). Jonasson (1991) has recognizedan increase in colluvial processes, particularlydebris¯ows, in other parts of Scandinavia fromc. 4600 until c. 3800 years BP.

Our radiocarbon data on the Holocene debris-¯ow deposits are quite limited, but the dates areconsistently younger than 5000 years BP andindicate several pulses of debris¯ow processes(Fig. 39). The dates are mainly from the localpalaeosols that separate debris¯ow lobes or leveÂesfrom the underlying colluvium or glacial till. Thegreater abundance of younger debris¯ow depositssuggests that some kind of geomorphic thresholdhas been reached by the mountain slopes around3000 years BP, possibly due to the climaticdeterioration, creep processes, slope-strippingby snow avalanches and anthropogenic defores-tation. Soil erosion and the increased weatheringof the glacial mantle, enhanced by vegetation, arelikely to have caused the more recent debris¯owson slopes of quite modest gradients. The crystal-line bedrock renders plant-root growth a destabi-lizing agent for the mantle. A similar increase indebris¯ow processes around 3000±2500 and 500±

300 years BP has been recognized in other parts ofEurope (Selby, 1982; Innes, 1983; Brazier &Ballantyne, 1989; GonzaÂlez DõÂez et al., 1996).Kotarba (1991) has noted an increase in debris-¯ows in the Polish Tatra Mountains at the end ofthe Little Ice Age.

The late Holocene increase in debris¯ow pro-cesses correlates with an increase in water¯owactivity on the colluvial slopes (Blikra & Nesje,1997). Likewise, the dating of soli¯uctional lobesin the Jostedalsbreen area (Nesje et al., 1989)indicates that the creep processes there com-menced c. 3200±2800 years BP. In Jotunheimenarea, Matthews et al. (1986) have recognized anincrease in slopewash processes from c. 3000years BP and an increase in soli¯uction rate fromc. 1000 years BP.

Although some debris¯ows have apparentlybeen associated or coeval with snow¯ows, theregional compilation of data suggests that thepronounced pulses of the two processes tended toalternate with each other, rather than be coeval(Fig. 39). The climatic conditions conducive todebris¯ows were probably different from thosecausing snow avalanches. The record of snowavalanches is thought to re¯ect an increase insnowstorm frequency, whereas the record ofdebris¯ows, as well as soli¯uction, re¯ects milderclimatic conditions and an increase in high-intensity rainfall (Sandersen et al., 1996). Thelatter suggestion is supported by the record ofwater¯ow processes (Blikra & Nesje, 1997); localpalaeobotanical data indicate a warmer climateand forestation by abundant deciduous treesduring periods with few or no snow avalanches(Blikra, 1994; Blikra & Selvik, 1998); and radio-carbon data support the onset of peat formation inthe outer coastal zone of the Sunnmùre area(Solem, 1989).

Rockfall record

The radiocarbon data on rockfall deposits arerelatively few (Blikra & Nesje, 1997), but someperiods of pronounced rockfall processes cantentatively be recognized (Fig. 39). The oldestrockfall deposits correspond to the Younger Dryasphase of climatic cooling, and are separated by aregional hiatus from the rockfall deposits ofmiddle to late Holocene age. Rockfall avalanchesin the middle Holocene were infrequent andoccurred probably as a seasonal phenomenonrelated to the early phase of climatic deteriorationand/or increased precipitation. Nesje, Blikra &Anda (1994) have used the Schmidt-hammer



method to estimate the age of the deposits of twolarge rockfalls in Norangsdalen, western Norway,and concluded that these were younger than c.6000 years BP.

The Holocene colluvial record in the study areasuggests an increase in rockfall processes sincec. 3800 years BP, which probably re¯ects the lateHolocene climatic deterioration, with sharper,colder winters and more rainy summers. Therockfall deposits are locally separated by palaeo-sols, which indicate climatic ¯uctuations andrelatively long phases of nondeposition and slopestability. Rockfall processes are enhanced byfreeze-thaw cycles, but often triggered by thun-derstorms, heavy rainfall and/or snowpack melt-ing (Sandersen et al., 1996).

DISCUSSION

The colluvial record has further been comparedwith the proxy record of atmospheric temperaturevariations based on the mountain glacier ¯uctu-ations in the Sunnmùre, Jostedalsbreen and Jot-unheimen areas (Nesje & Kvamme, 1991; Mat-thews & KarleÂn, 1992; Nesje & Dahl, 1993) and theHardangerjùkulen area further to the south (Dahl& Nesje, 1994). The two types of palaeoclimaticrecord, although based on quite different criteria,show a remarkably good agreement (Fig. 39). Theglacier-derived record indicates an overall cli-matic deterioration from at least 6000 years BP,and the pattern of ice-front oscillations corre-sponds quite well with the colluvial record. TheHolocene phases of pronounced snow-avalancheactivity can readily be correlated with the neo-glacial stages; except, perhaps, for the phase ofclimatic cooling around 9000 years BP, for whichno avalanche record has been recognized in thecolluvial successions. The record from Jostedalsb-reen indicates neoglacial stages dated to c. 6000,3700±3100, 2600±2100, c. 1800 and < 1400 yearsBP. The Jotunheimen record shows a neoglacialstage around 6400±5900 years BP, whereas therecord from Sunnmùre indicates neoglacial stagesaround 3400±3000, 2200±2100, < 1600 and< 1000 years BP (Matthews & KarleÂn, 1992). Therecord from Hardangerjùkulen indicates neogla-cial stages around 6300±5200, 4750±3900 (severalphases) and < 3800 years BP. Based on theradiometric data on peat growth and humi®cat-ion, Matthews (1991) has inferred periods ofpronounced climatic deterioration around 5700,4300±3200 and 1750±1300 years BP. All thesedata correspond rather well with the colluvial

record. It is worth noting also that the resolutionof the colluvial record is, in fact, signi®cantlyhigher than that of the other data types.

Nevertheless, the correspondence of the collu-vial and glacial records is by no means perfect,and the discrepancies require a comment. Firstly,one must be aware that a glacier-derived recordhas a limited resolution, bears many uncertaintiesand necessarily re¯ects the local conditions only.Notably, there are many discrepancies betweenthe records from the different glaciers (Fig. 39).For comparison, snow avalanches are seasonalphenomena, and their dependence on local con-ditions is far less critical when the overall recordfrom a larger number of localities is compiled.Secondly, the climatic aspects re¯ected in theglacier dynamics and the colluvial-slope dynam-ics may not be the same, and also the responsethresholds of glaciers and colluvial systems withrespect to climatic changes may be different.Arguably, the record of ice-front oscillationsre¯ects the net rate of snow accumulation, whichmeans both the winter snowfall rate and thesummer air temperatures. The record of snowavalanches, in contrast, is likely to re¯ect thesnowfall rate and the occurrence of snowstorms,which means the winter conditions alone.

If the latter notion is correct, the relatively goodcorrespondence between the two types of palaeo-climatic record may indicate that the glacieroscillations in western Norway have been con-trolled mainly by the winter conditions. Further-more, the discrepancy between the records fromHardangerjùkulen and Jostedalsbreen for 6300±4000 years BP may suggest that the glacier dy-namics in the former area were more related to thesnowfall rate, whereas the effect of annual tem-perature variations was more important in thelatter area. In short, it is crucial to comparedifferent types of climatic proxy record, becausetheir correspondence and discrepancies may bequite meaningful.

CONCLUSIONS

Colluvial depositional systems are one of theunexplored frontiers of clastic sedimentology.Colluvium, in striking contrast to its sister alluvi-um, is hardly mentioned in sedimentologicaltextbooks and has virtually been ignored in therecent development of clastic facies models. Mod-ern colluvial deposits, due to their unsorted andseemingly chaotic nature, have commonly beenregarded as nearly intractable `slope breccias', and



the colluvial systems have thus drawn littlesedimentological interest. Relatively few ancientcolluvial successions have been reported (seereview in Nemec & Kazanc�, 1999), although theirpreservation potential should not necessarily bemuch lower than that of `footwall' alluvial fans orvalley-®ll alluvium, which have commonly beenrecognized in the ancient record. One may rea-sonably suspect that colluvial deposits in basinalstudies were often lumped as proximal fanglome-rates, simply because the sedimentological criteriafor their distinction were not available.

The postglacial colluvium in western Norway,related to the steep mountain slopes of valleysides and fjord margins, comprises deposits ofslope-waste avalanches ranging from rockfalls/debrisfalls to debris¯ows and snow¯ows, withsome role played also by water¯ow and debris-creep processes. The sediment provenance here isspeci®c, involving gneissic bedrock and itsglacigenic mantle, and the colluvial facies assem-blages range from subaerial to subaqueous. Thestudy has focused on the processes of sedimentmovement and their depositional products, withspecial emphasis on snow avalanches, whose roleas a depositional agent was unexplored by sedi-mentologists.

The chronostratigraphic analysis of the localcolluvial successions indicates that these depos-its bear an important proxy record of regionalpalaeoclimatic changes. This aspect is particular-ly well-shown by the colluvial-fan deltas, whosevaried facies assemblages and bedding architec-ture re¯ect the combined record of climate andrelative sea-level changes.

Although it is true that colluvial systemsdepend upon the local slope conditions and thata particular genetic facies may not be a reliablepalaeoclimatic indicator (Van Steijn et al., 1995),our ®eld studies lead us to contend that thestratigraphic changes in colluvial facies assem-blages re¯ect signi®cant climatic changes. Thisnotion is supported by the regional correlation oflocal facies changes and their good correspon-dence with other types of regional palaeoclimaticdata (see also Nemec & Kazanc�, 1999).

The palaeoclimatic record deciphered from thecolluvial successions corresponds quite well withthe record of alpine glacier ¯uctuations in west-ern Norway, and seems to have, in fact, a muchhigher resolution than the latter type of climaticproxy data. Therefore, it may be concluded thatthe colluvial systems have acted as sensitiverecorders of regional climatic changes. However,it should be emphasized that colluvial systems

are also highly dependent upon local slopeconditions, and the palaeoclimatic record is thusmost legible when facies data from a largernumber of localities are compiled in chronostrati-graphic terms.

The study demonstrates that colluvial fans andaprons, with their speci®c range of processes andhigh sensitivity to changes in climatic conditions,are a fascinating type of depositional system.Comparative studies should show how the signalof Quaternary climatic changes has been recordedin colluvial successions in other regions, in awider range of climatic zones (Nemec & Kazanc�,1999).

ACKNOWLEDGMENTS

The study was sponsored by the GeologicalSurvey of Norway (NGU), with additional ®nan-cial support from the Norwegian Research Coun-cil (NFR) and the University of Bergen. EinarAnda, Oddvar Longva and Dag Ottesen arethanked for helpful discussions and ®eld assis-tance. The radiocarbon dating of samples wasdone by S. Gulliksen at the Radiological DatingLaboratory in Trondheim. Figures were drafted byB. SvendgaÊrd and I. Lundquist (NGU). Themanuscript was critically reviewed by BjoÈrnAndersen, Brian Jones, Eiliv Larsen, GeorgePostma and Henk Van Steijn. This project repre-sents activities of the Norwegian Research Groupfor Quaternary Sedimentology (NFKS).

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Manuscript received 16 February 1997;revision accepted 28 January 1998



Postglacial colluvium in western Norway: depositional processes, facies and palaeoclimatic record

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