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Postglacial adjustment of steep, low-order drainage basins,Canadian Rocky Mountains

T. Hoffmann,1,2 T. Müller,1,3 E. A. Johnson,1 and Y. E. Martin1

Received 7 May 2013; revised 4 November 2013; accepted 18 November 2013.

[1] It is generally argued that Pleistocene glaciation results in increased sediment flux inmountain systems. An important, but not well constrained, aspect of Pleistocene glacialerosion is the geomorphic decoupling of cirque basins from main river systems. This studyprovides a quantitative link between glacier-induced basin morphology, postglacial erosion,and sediment delivery for mountain headwaters (with basin area <10 km2). We analyze themorphology of 57 headwater basins in the Canadian Rockies and establish postglacialsediment budgets for select basins. Notable differences in headwater morphology suggestdifferent degrees of erosion by cirque glaciers, which we classify into headwater basins witheither cirque or noncirque morphology. Despite steeper slope gradients in cirque basins,higher-mean postglacial erosion rates in basins with noncirque morphology (0.43–0.6 mm a�1)compared to those in cirques (0.19–0.39 mm a�1) suggest a more complex relationshipbetween hillslope erosion and slope gradient in calcareous mountain environments thanimplied by the threshold hillslope concept. Higher values of channel profile concavity andlower channel gradients in cirques imply lower transport capacities and, thus, lower sedimentdelivery ratios (SDR). These results are supported by (i) postglacial SDR values for cirquesand noncirque basins of <15% and >28%, respectively, and (ii) larger fan sizes at outlets ofnoncirque basins compared to cirques. Although small headwater basins represent the steepestpart of mountain environments and erode significant postglacial sediment, the majority ofsediment remains in storage under interglacial climatic conditions and does not affectlarge-scale mountain river systems.

Citation: Hoffmann, T., T. Muller, E. A. Johnson, and Y. E. Martin (2013), Postglacial adjustment of steep, low-orderdrainage basins, Canadian Rocky Mountains, J. Geophys. Res. Earth Surf., 118, doi:10.1002/2013JF002846.

1. Introduction

[2] High mountains experience some of the largest sedi-ment fluxes on earth [Milliman and Syvistki, 1992] and aremajor sediment sources for large river systems. Sedimentflux in mountains, however, is strongly controlled by the spa-tial organization of geomorphic processes, which comprisestheir location within the basin, as well as their successionand connectivity along the flow path [Fryirs et al., 2007;Tunnicliffe and Church, 2011]. Sediment storage modifiesthe efficiency of sediment conveyance by changing the virtualvelocity of sediment through the system [Hinderer, 2012;

Straumann and Korup, 2009] and causes distinct mismatchesbetween mountain erosion and sediment delivery to neighbor-ing lowlands [Church and Slaymaker, 1989]. Identification ofsources and sinks of sediment and understanding controls ofthe spatial organization and connectivity of geomorphic pro-cesses are fundamental to obtain deeper understanding of tran-sient landscape evolution [Brardinoni and Hassan, 2006],river management and restoration [Owens, 2005; Slaymaker,2003], and riverine ecology [Buffington et al., 2004;McCleary and Hassan, 2008].[3] Many high mountains have been sculpted by strong

glacial erosion during the Pleistocene that resulted in valleywidening and overdeepening [Brocklehurst and Whipple,2002, 2006; MacGregor et al., 2000] and the formation ofglacial cirques [Evans, 2006; Hooke, 1991], U-shaped val-leys, and widespread glacial deposits [Barsch and Caine,1984; Hallet et al., 1996; Montgomery, 2002]. The retreatof glacial ice exposes oversteepened hillslopes that are suscep-tible to rockfalls, deep-seated landsliding, gully erosion, anddebris flows, and can also result in valley aggradation andreworking of valley deposits through debris flow activity andfluvial processes [Brardinoni et al., 2012; Dadson andChurch, 2005]. It has been argued that sediment fluxes causedby these processes remain elevated even several thousand

Additional supporting information may be found in the online version ofthis article.

1Biogeoscience Institute, University of Calgary, Calgary, Alberta,Canada.

2Department of Geography, University of Bonn, Bonn, Germany.3Department of Geography, University of British Columbia, Vancouver,

British Columbia, Canada.

Corresponding author: T. Hoffmann, Department of Geography,University of Bonn, Meckenheimer Allee 166, DE-53115 Bonn, Germany.([email protected])

©2013. American Geophysical Union. All Rights Reserved.2169-9003/13/10.1002/2013JF002846

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JOURNAL OF GEOPHYSICAL RESEARCH: EARTH SURFACE, VOL. 118, 1–17, doi:10.1002/2013JF002846, 2013

years after the retreat of valley glaciers [Ballantyne, 2002a;Church and Ryder, 1972]. For example, Church andSlaymaker [1989] suggest that an increase in specific sedimentyield of British Columbian rivers with drainage basin sizefrom 1 to 104 km2 results from the redistribution of sedimentthat was delivered by glaciers 10 ka ago.[4] In general, research on the adjustment of topography to

deglaciation has focused on the accelerated geomorphicactivity of hillslope processes or on the reworking of tilland other glacial deposits [Ballantyne, 2002b]. Important,but not well-constrained, impacts of glacial landformsculpting include its effects on postglacial erosion and sedi-ment delivery from mountain headwaters [Barsch andCaine, 1984]. Oversteepened hillslopes and flat, steppedvalley bottoms formed by glacial erosion in alpine headwatersmay (i) modify the sequence of geomorphic processes andchannel reach morphology along the flow path of sediment,(ii) increase rates of gravitational mass movements due tooversteepened hillslopes, (iii) reduce the transport capacityof streams that currently occupy the formerly glaciated valleysbecause of reduced slope gradients and overdeepening alongthe thalweg, and (iv) decrease connectivity between hillslopesand the channel network [Brardinoni and Hassan, 2006;2007]. Thus, it may be hypothesized that Pleistocene glacialerosion in mountain headwaters may have two counter-balancing effects on postglacial sediment fluxes: first, it resultsin increased postglacial hillslope erosion and second, sedimentexport from headwater basins is decreased as a consequence oflow transport capacity conditions along streams. Consequently,mountain headwaters sculpted by glacial erosion may producelarge amounts of sediment during postglacial times, but exportvery little sediment beyond their basins outlets compared tomountain landscapes that are characterized by fluvial processesand morphology (e.g., V-shaped valley forms and concavelongitudinal profiles).[5] Several studies have estimated postglacial erosion and

deposition rates for single geomorphic processes [e.g.,Brardinoni et al., 2009; Hinderer, 2001; Hoffmann andSchrott, 2002; Sass, 2007] or have established postglacialsediment budgets for multiple geomorphic processes [e.g.,Campbell and Church, 2003; Schrott and Adams, 2002;Schrott et al., 2003; Slaymaker, 1993; Tunnicliffe andChurch, 2011; Wichmann et al., 2009]. Results of the abovestudies have demonstrated the effects of Pleistocene glacia-tion on geomorphic process rates and their spatial organiza-tion [Ballantyne, 2002b]. However, to our knowledge, nostudy has studied systematically the relative importance ofincreasing hillslope erosion vs. decreasing sediment deliveryfrom glacial headwater basins after deglaciation. This paperaddresses this gap in the research literature, focusing onmountainous headwater basins (< 10 km2). Despite theirsmall individual areal extent, their steep and unstable terrainimplies that these headwater basins often generate the highestamounts of sediment relative to more gentle downstreamlocations [Ballantyne, 2002b; Montgomery and Brandon,2002]. Collectively, the large number of these headwaterdrainage basins suggests a significant areal proportion withinthe larger drainage basins [Strahler, 1957], adding to theirerosional significance. The primary objective of this paperis to analyze the effects of glacial erosion on postglacial sed-iment production and delivery of headwaters in the CanadianRocky Mountains. First, we quantify the morphology of

headwater basins to investigate the effects of glacial erosionon basin morphology. This is achieved by comparing headwa-ter drainage basins with different degrees of Pleistoceneglacial imprint. Second, we estimate postglacial erosion andsediment delivery using a sediment budget approach for se-lected headwaters with variable morphology in the CanadianRockies. We discuss the findings in terms of how glacialerosion in headwater basins has affected postglacial erosion,sediment delivery, and connectivity.

2. Study Area

[6] The study area is located in the Front Range of theCanadian Rocky Mountains, Alberta, Canada (Figure 1).The topography, characterized by NNW-SSE aligned ridges(rising up to 3100 m above sea level) and valleys (minimumelevation 1400 masl), is the result of tectonic uplift during theLaramide Orogeny and strong glacial erosion during thePleistocene [Osborn et al., 2006]. The thin-skinned foldand thrust deformation of the Laramide Orogeny during theLate Cretaceous to Early Tertiary involved mostly veryresistant Paleozoic carbonates that were thrust eastward overmuch less resistant Mesozoic sandstones, siltstones, andshales. The carbonates are characterized by steep slopes(slope gradient mostly> 0.6) and high elevations, generallyforming ridges that dip toward the SW and cut across sedimen-tary beds on the NE flanks. In contrast, the Mesozoic clasticsediments are distinguished by gentler slopes (slope gradientmostly< 0.6) and subdued relief that forms the valleys of thestudy area. The present-day topography of the region showsnotable evidences of glacial erosion during the Pleistocene,including cirques, U-shaped and hanging valleys withoversteepened rock walls [Osborn et al., 2006]. Except forthe highest peaks (above ~2700masl), the entire study areawas overridden by the Bow glacier (an outlet glaciers of theCordilleran Ice Sheet, CIS), the Kananaskis valley glacier,or local cirque glaciers during the last glacial maximum(LGM) [Jackson, 1980, 1987]. Several transfluences at thenorthern end of the Kananaskis valley indicate that theBow valley glacier fed the Kananaskis Valley with ice andthat the Kananaskis valley glacier overflowed the easternmost ridges. These ridges mark the transition between theFront Range and the Foothills of the Canadian RockyMountains. Since the deglaciation about 12 ka BP [Jackson,1980], oversteepened cliffs and dip slopes of the Paleozoiccarbonate have been prone to rockfalls and rock slides[Cruden and Hu, 1993]. Extensive debris cones at tributaryoutlets, which date back before the early Holocene (as indi-cated by the occurrence of the 6800 year old Mazama ash inthe upper part of one of the fan deposits), suggest increased de-bris flow activity and fluvial sedimentation at the transition be-tween the Pleistocene and Holocene [Jackson et al., 1982;Roed and Wasylyk, 1973]. Significant glacial advances in theHolocene occurred during the Little Ice Age that culminatedduring the 18th or 19th centuries [Beierle et al., 2003;Osborn and Luckman, 1988].[7] While carbonate slopes generally form steep bedrock

slopes with no or very limited cover of colluvial deposits,gentler slopes of Mesozoic rocks are generally covered byfine-grained regolith or colluvial deposits. The thickness ofthe colluvium varies in response to hillslope curvature andgradient, and the distance from the ridgelines. The majority

HOFFMANN ET AL.: POSTGLACIAL SEDIMENT FLUX

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of the colluvial slopes are covered by Lodgepole pine (Pinuslatifolia) - Engelmann spruce (Picea engelmannii) (dominat-ing below 1800 masl) and Engelmann spruce - Subalpinespruce (Abies lasiocarpa) (dominating above 1800 masl).The average precipitation at Kananaskis Lakes over a 13 yearperiod (elevation 1670 m) was 360 mm from October toMay, and 250 mm from May to October (Eastern RockiesForest Conservation Board, 1968).

3. Methods

3.1. Classification of Headwater Basins

[8] To identify the impact of Pleistocene glacial erosionon headwater basin morphology and its effect onpostglacial sediment flux, we selected 57 small drainagebasins (0.2 – 10.8 km2) in Kananaskis Country. KananaskisCountry is ideally suited for this study since it provides a largenumber of tributary basins along the major trunk valleys (theKananaskis valley, the Spray valley, and the Bow valley, seeFigure 1). The accumulation of massive alluvial and debrisflow fans at the tributary outlets indicates a high trapping effi-ciency at the transition of the tributaries to the main valley.The fans were generally well preserved because they formedat a relatively large distance (e.g., several 102 m) from thetrunk river and partly on top of post-LGM terraces, which pro-vided a stable base level since accumulation was initiated.

The basins are characterized by similar lithology (mainlylimestone and dolostone) but a large variety of basin mor-phologies. Different morphologies are mainly attributed todifferent degrees of glacial erosion (see below). Theseconditions are suitable to study the glacial imprint onpostglacial sediment fluxes.[9] The headwater basins were chosen and classified into

three basin morphologies using aerial photograph interpreta-tion, field observation, and interpretation of a shaded reliefmap (extracted from a 25m digital elevation model, DEM):[10] First, “glacial cirques” are characterized by glacial

landforms such as U-shaped cross sections with steep rockwalls (slope gradient> 0.72 m m�1), flat or stepped valleybottoms. Abundant talus slopes are located at the transitionbetween rock walls and valley bottoms. Lateral and terminalmoraines in these basins indicate the existence of local cirqueglaciers that often date back to the Younger Dryas/EarlyHolocene and Little Ice Age [Beierle et al., 2003; Gardneret al., 1983]. Lakes are present in some cirques and indicateoverdeepening caused by glacial erosion.[11] Second, headwater basins with V-shaped cross sec-

tions are dominated by gentle hillslopes that are generallycovered by a layer of regolith or colluvium that increases indepth from a few centimeters on ridge tops (as indicated byfrequent rock outcrops) up to several meters at the hillslopebase. Erosion of regolith and colluvial deposits by debris

Figure 1. Location (inset) and shaded relief map of Kananaskis River basin in the Front Ranges of theCanadian Rocky Mountains, and location of the 57 headwater basins. Numbers refer to those used in sec-tion “postglacial sediment budgets.” The 2100 masl elevation line represents the approximated upperboundary of the Cordilleran Ice sheet and the Kananaskis valley glacier during the last glacial maximum.


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flows and subsequent deposition in debris cones and fans isabundant. Abundant debris flow deposits in the valleybottom indicate that sediment transport by debris flows isgenerally much more efficient than fluvial transport. We rec-ognized no evidence of prior wet-based glaciation (e.g., lateraland terminal moraines) in the V-shaped basins, indicating thaterosion by local cirque glaciers during the Pleistocene did nottake place or was insignificant and did not modify themorphology of the basins. We refer to these basins as“noncirque,” even though we are aware that they werecovered by glacial ice, either by the Bow or theKananaskis valley glacier, at least during the LGM. Thus,the term “noncirque” means “no characteristic cirquemorphology” and does not exclude the prior existence ofglacial ice and erosion.[12] Third, “complex drainage basins” are characterized

by a mixture of glacial cirque characteristics in the upperpart of the basins and noncirque characteristics in thelower part.

3.2. Geomorphometry of Headwater Basins

[13] For each basin, we extracted a set of basin-widegeomorphic attributes including (i) distribution of slopegradients, (ii) mean geophysical relief, (iii) hypsometricintegral (HI), (iv) equilibrium line altitude (ELA), and(v) thalweg steepness and concavity. All attributes werederived from a 25 m DEM. The Kruskal-Wallis test wasused for each attribute to test if central tendencies betweenthe classified basin types are significantly different[Hollander andWolfe, 1999]. Results are given by the p-valueof the Kruskal-Wallis test (pkw-value). Significant differ-ences between basin types are given at the 5% level (i.e.,pkw> 0.05). Furthermore, we tested the effects of thelimited DEM resolution by comparing the parametersderived from the 25 m DEM with those derived from a1 m airborne light detection and ranging (LIDAR)-DEM(downscaled to 5m) that was available for the northern partof the study area (covering basins 1 to 14 in Figure 1).Results of the uncertainty assessment are given in thesupporting information.[14] Slope distributions, mean slope gradients, and the ar-

eal fraction of each basin that is steeper than 0.72 m m�1

were determined from the 25 m DEM. The slope gradientof 0.72 m m�1 represents the transition from soil/sediment-covered hillslopes to bedrock hillslopes (see below).Bedrock coverage and mean slope gradient were frequentlyused as a proxy of long-term erosion rates in alpine environ-ments [e.g., Binnie et al., 2007; Norton et al., 2010a].[15] The mean geophysical relief is given by the “missing

volume” of each basin between the smooth surfaceconnecting the highest points in the current landscape andthe present-day DEM divided by surface area, and is gener-ally used as a proxy of long-term erosion [Small andAnderson, 1998]. It is different from the “ordinary” relief,which is the difference between the maximum and minimumelevation within each basin. Herein, we do not use absolutevalues of geophysical relief as a metric of erosion, becauseof the large uncertainty associated with the link betweenlong-term erosion and geophysical relief. Rather, we usedifferences in geophysical relief between basin types asan indicator of distinctive erosional histories during thePleistocene. Ridge to ridge topography was interpolated

based on the extraction of ridgelines that outline a givenbasin and the interpolation of ridgeline elevations using a lin-ear triangulation in ArcGIS [similar to Brocklehurst andWhipple, 2002].[16] The hypsometry of each basin was derived by the

hypsometric curve (i.e., normalized cumulative frequencyof elevation or in other words the proportion of basin area be-low a certain elevation ranging from zmin and zmax) and thehypsometric integral:

HI ¼ Zmean � Z min

Z max � Z min(1)

where zmin, zmean, and zmax are the minimum, mean, andmaximum elevation within each basin, respectively. Thehypsometric integral is the area under the normalizedhypsometric curve and ranges by definition between 0 and1. It has been used to demonstrate changes in the importanceof fluvial vs. glacial processes [Kirkbride and Mathews,1997; Montgomery et al., 2001] or to characterize the pro-gressive glacial modification of initially fluvial landscapes[Brocklehurst and Whipple, 2004; Sternai et al., 2011]. HIsof glaciated landscapes (i.e., U-shaped valleys: HI≈ 0.2) aretypically lower than those of fluvial landscapes (i.e., V-shapedvalleys: HI≈ 0.5) [Sternai et al., 2011].[17] Based on the hypsometric curve and an assumed accu-

mulation area ratio (AAR; i.e., ratio of glacial area above andbelow the glacier equilibrium line) of 2:1 [Meierding, 1982;Porter, 2001], we approximated the equilibrium line altitudethat corresponds to the full coverage of each basin by a self-contained, local cirque glacier (herein defined as basin-specificELA). This was compared with the ELA of present-dayglaciers, which is estimated using the glacial snow line oforthophotos taken in 1999 and 2008 at ~2700 masl. We arguethat the difference between the basin-specific ELA and thepresent-day ELA is a proxy for the likelihood of each basinto achieve full coverage by a local cirque glacier. Basins withsmall differences (ΔELA-values) would tend to be glaciatedearlier by a cirque glacier than basins with large ΔELAs. Inthe case of large ΔELAs (i.e., low lying basins), it is morelikely that they are covered with ice from the main trunkglacier (e.g., Bow or Kananaskis valley glacier), rather thanbeing occupied by a local cirque glacier [Rutter et al., 2006].While the AAR method relates to mapped glacier outlinesand the ice surface [Ramage et al., 2005], the estimated hypso-metric curve refers to the surface of each headwater basin.Therefore, ΔELA values represent only a first approximationof the necessary drop of the present-day ELA for a headwaterto be fully covered by a local cirque glacier. In fact, ΔELAvalues are typically overestimated by the mean thicknessof cirque glaciers (approx. <100 m) that occupy theheadwater basins.[18] Slope-area relations of river longitudinal profiles are

frequently used to derive the erosive power and sedimenttransport conditions in river systems [Korup, 2006; Nortonet al., 2010b; Wobus et al., 2006]. Contributing basin area,A, was calculated using D8 flow accumulation [Garbrechtand Martz, 1997]. We extracted river longitudinal profilesfrom the intersection of contour lines at 20 m elevation inter-vals and downstream flow paths [Snyder et al., 2000;Wobuset al., 2006]. We estimated channel steepness, ks, and con-cavity, θ, for the main channel of each basin by linear


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regression of S [m m�1] and A [m2] using a log-log transfor-mation with contributing area> 104 m2 to the channel outlet:

S ¼ ks � A�θ (2)

(compare Figure S1). The minimum contributing area waschosen since most basins showed a break in the scaling at104 m2. Here we use ks and θ to measure the impact of glacialcirque erosion on the channel longitudinal profile, without ad-dressing the implications of uplift and erosion [Brocklehurstand Whipple, 2007]. However, since ks is not dimensionlessand depends on θ, it cannot be compared amongst differentbasins. To compare the channel steepness for cirque andnoncirque basins, we calculated the normalized channel steep-ness, ksn, and the representative slope, Sr. The value of ksn iscalculated by equation (2) with a fixed θ =0.25 as providedby the mean of all basins. Sr is calculated by normalizingcontributing area using a representative area, Ar=1.4� 106

m2, which was chosen based on the median of all basins:

S ¼ SrA

Ar

� ��θ

(3)

Sr is the expected channel slope at Ar and is scale independent.Therefore, ksn and Sr allow for comparison of channelsteepness between different basins [Sklar and Dietrich, 1998].[19] Uncertainties associated with the basin-wide parame-

ters were tested by comparing parameters derived fromDEMs with 25 m and 5 m grid spacing. No significantlimitations due to the low resolution of the 25 m DEMused in this study were revealed (see Figure S2 supportinginformation for more information on the uncertaintyassessment).[20] We analyzed the areal extent of alluvial and debris

flow fans below the outlet of the headwater basins at the tran-sition to the main trunk valleys. The areal extent of the fanswas derived from aerial photographic interpretation, field ob-servation, and interpretation of a shaded relief map. The latterwas extracted from the 1 m LIDAR-DEM [for details on theLIDAR-DEM see Hopkinson et al., 2009, 2012] and coveredthe northern part of the study site (e.g., basins 1 to 14 inFigure 1). To evaluate the effect of basin size, fan area wasgiven in absolute size (m2) and specific size (m2/m2), whichis normalized to basin size. We analyzed only those basins(n = 37) for which a high preservation potential of postglacialfan sediments could be inferred and for which evidence ofmajor erosion is lacking (e.g., fan deposition on top ofwell-established postglacial terraces, which escaped chan-nel erosion at the fan location through the trunk channel).Under these conditions, data from basins dominated by de-bris flows in the Italian Eastern Alps showed that debris fanarea can be expressed as a linear function of sediment yield[Brardinoni et al., 2012], justifying our approach to usefan size as a first proxy of sediment delivery from the head-water basins.

3.3. Compilation of Postglacial Sediment Budgets

[21] To further test the hypothesis of decreased sedimentdelivery from glacial cirques, postglacial sediment budgetsof five headwater basins in the Kananaskis Country were de-termined. These basins were selected from those used in theanalysis of basin morphology to represent typical headwater

basins with noncirque (basins #02 and #06 in Figure 1) andglacial cirque morphology (basins #24, #26, and #47 inFigure 1). Following Slaymaker [2003], sediment budgetsaccount for the sources, storages, and export of sediment.In this study, we focused on postglacial sediment storageand estimated the volume (m3) of sediment stored in the fansat the basin outlets (SVfan) and within the basins (SVbasin).Assuming that sediment storage in the basins and the fanswas initiated after glacial retreat, the mean postglacial ero-sion rate, ER [mm a�1], since time T of deglaciation (~12.5ka BP) for each basin is given by:

ER ¼ ρsρR

SV basin þ SV fanð ÞA�T (4)

where the density ratio of sediments and rock is ρS/ρR ~ 0.62(based on 50 measured average densities ρS = 1600 kg m�3

and ρR = 2600 kg m�3) and A (m2) is the contributing basinarea (excluding fan area) [Hinderer, 2012; Schrott et al.,2003]. We assumed the same T for each basin, due to therapid retreat of the glaciers at end of the Late Wisconsinan[Jackson, 1980; Osborn and Gerloff, 1997].[22] The mean postglacial sediment delivery ratio, SDR,

for each basin is then calculated:

SDR ¼ yield

erosion¼ SV fan

SV basin þ SV fan(5)

Without any dating control, ER and SDR representpostglacial averages. We, therefore, focus on the spatial con-figuration and the control of basin morphology on postglacialaverages of ER and SDR.[23] Sediment storage was determined using geomorpholog-

ical mapping, which provides information on the spatial extentof geomorphological forms, processes, and surficial materials,in combination with geophysical soundings that were used toestimate sediment thicknesses. The geomorphic mapping wasbased on 1m orthophoto interpretation and field reconnais-sance. For basin #02 and #06, a 1 m-LIDAR-DEM was alsoavailable. The hillshade, extracted from the LIDAR-DEM,allowed us to map geomorphic forms and processes under thedense forest cover of basin #02 and #06. The mapping wasconducted using the mapping system for alpine environments[Kneisel, 1998] and following an available terrain inventoryof the Kananaskis area [Jackson, 1987] that is based onHowes and Kenk [1997]. According to these mapping systems,we differentiated six geomorphic processes: (i) glacialprocesses, (ii) diffusive hillslope processes (including soilcreep, shallow landslides, animal borrowing, tree fall), (iii) rockfalls/slides, (iv) snow avalanches, (v) debris flows, and (vi)fluvial processes. Associated surficial materials are (i) till, (ii)regolith and colluvial creep deposits, (iii) talus and debrisdeposits, and (iv) alluvial deposits (Table 1). In our definition,colluvial deposits are limited to sediments related to diffusive(creep) processes, while coarse-grained deposits of rockfalls/slides, snow avalanches, and debris flows are referred totalus and debris. Colluvial deposits are differentiated fromtalus and debris based on texture, stratigraphy, locationalcontext (e.g., rock fall talus require a rock wall steep enoughfor rock fall to occur), and their surface form. Accordingly,we further subdivided the mapped geomorphic units based


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on their form to eight classes: rock wall, slope, slope bottom,cone, channel, valley, fan, and moraine [compare alsoBallantyne and Harris, 1994; Otto and Dikau, 2004].Typical combinations of processes, surficial materials, andforms in the study site are summarized in Table 1. A sketchof the geomorphic units is shown in Figure 2. Shallowsediment thicknesses (up to 1m) were mapped in the fieldand estimated by outcrops or digging pits. For instance, thick-nesses for bedrock slopes (BR) covered with a thin layer ofregolith or colluvial deposits ranged from 0 to 0.5 m as repeat-edly measured in several pits; in these cases, a mean thicknessof 0.25 m was assigned (key BR/Cv in Table 1). On colluvialslopes or talus slopes, with occasional bedrock outcrops, sed-iment thicknesses varied between 0.5 and 1 m (as measured inseveral pits and outcrops), with a mean value of 0.75 m (keyCv/BR and TS/BR in Table 1).

[24] To estimate sediment thickness in locations of mas-sive sediment deposits (e.g., talus slopes, debris cones, tilldeposits, and valley fills), we utilized refraction seismic andground penetrating radar [for details on data collection andprocessing see Otto and Sass, 2006; Otto et al., 2009; Sass,2006; Schrott and Hoffmann, 2008]. Since the applicationof geophysical soundings in steep terrain is very time con-suming, we were only able to measure four refraction seismicsoundings and four radar soundings in three basins (basin#02, #06, and #47). Except for the refraction sounding inthe valley fill of basin #06, the sediment was distinguishablefrom the underlying bedrock in each sounding. Thicknessesof talus slopes were generally 5 to 10 m in the upper partsof the deposits and increased to 30 m in the lower parts atthe transition to the valley. Measured sediment thicknessesof debris cones were generally less than 20 m (Table 1).

Figure 2. Sketch of geomorphic units mapped in the field. Geomorphic mapping included the mappingof landforms and sediments and following Jackson [1987] and Kneisel [1998]. Dominant processes (e.g.,rock fall, debris flow, and snow avalanche) were derived following Ballantyne and Harris [1994].

Table 1. Mapped Geomorphic Process and Form Unitsa

Geomorphic Process/Material Form Units Description Key Min Depth Max Depth

bedrock rock wall • steep bedrock wall BR 0 0steep slope • bedrock with shallow debris cover (mainly weathered bedrock) BR/Cv 0 0.5

diffusive hillslope processesb slope • shallow regolith or colluvium with bedrock outcrops (mainly at ridges) Cv/BR 0.5 1slope • colluvial veneer Cv 10

slope bottom • colluvial deposits of slope bottom Cbs 15fall/slide talus slope • talus deposits with major contributing rock wall TS 30

talus cone • cone-shaped talus deposits below bedrock chutes TC 30nival / snow avalanche channel • avalanche track/channel AVch 15

cone • cone-shaped, depositional area of avalanche tracks AVco 20debris flows channel • generally less than 10m wide DFch 15

debris cone • debris flow cone DFco 20valley • valley fill dominated by debris flows; width> 10m DFva 35fan • fan or cone at the basin-outlet DFf -

fluvial channel • channel with running water during most time of the year Ach 15fan • alluvial fan Af -

glacial moraine • terminal or side moraine Gm 25veneer • veneer of lodgment till either on hillslopes or in the valley bottom Gv 20

aMinimum and maximum sediment thicknesses were either mapped in the field or approximated based on ground penetrating radar and refraction seismic.bIncluding processes such as: rainsplash, tree throw, animal burrowing, shallow landsliding, and creep processes resulting from cyclic wetting and drying

as well as freeze/thaw and shrink/swell cycles.


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We assume that the uncertainty of the estimated sedimentthicknesses is approximately 25% [Hoffmann and Schrott,2003; Schrott and Hoffmann, 2008]. Valley fill sedimentsin the lower part of basin #06 were thicker than the penetra-tion depth of the refraction sounding, suggesting that thick-nesses are greater than 30 m. The geophysical soundingswere used to approximate the dip of the sediment/bedrockinterface at the transition between bedrock and sedimentarydeposits (given as α in Figure 3) and to evaluate maximumsediment thicknesses for the different geomorphic depositsand form units (compare Table 1). All estimated sedimentthicknesses include coarse- and fine-grained sediment.[25] Sediment thicknesses were interpolated for the entire

basins based on the mapped thicknesses, estimated dip of thesediment/bedrock interface and maximum thicknesses ap-proximated by geophysical sounding. First, we assigned themapped sediment thickness (e.g., BR=0 m, BR/Cv = 0.25 m,and Cv/BR=0.75 m) to areas that were estimated in thefield using geomorphic field mapping (see above). Second,the Euclidian distance to the nearest point with a mappedthickness was calculated and multiplied with the tangents of α(Figure 3). To evaluate uncertainty of α, we calculated two

scenarios with α=10° and α=20°. Third, the distance-depen-dent thickness was added to the nearest mapped thickness.Fourth, to avoid unrealistically high sediment thicknesses (inthe case of large distances to areas with shallow sediments),the thickness for each geomorphic process unit was limited tothe assigned maximum thickness that was derived from geo-physical exploration (maximum values are given in Table 1).Sediment volume in each mapped geomorphic process unitfor each basin was calculated based on area and mean inter-polated sediment thickness.[26] Consistent with other postglacial sediment budgets

studies, sediment export from the basins is assumed to beequal to the sediment stored in the fan at the outlet of thebasin [e.g., Hornung et al., 2010; Oguchi, 1997; Tunnicliffeand Church, 2011]. For basins #02, #06, and #47, thevolumes of stored fan sediment were estimated based onfan topography (derived from the LIDAR-DEM and the25m DEM) and the fan base [Campbell and Church,2003]. The debris fan of basin #02 is situated on top of aglacio-fluvial terrace that has been deposited shortly after de-glaciation of the main valley approx. 12.5 ka BP ago[Jackson, 1980]. The elevation of the terrace is located at1400 masl. The fan has recently been incised in the lower partby the channel of basin #02 and has been locally undercut bythe trunk channel (Kananaskis River) after deposition. Thus,the surface of this feature has been reconstructed based on theslope gradient of the fan and the intersection with the terracelevel at 1400 masl; this information is used to derive the vol-ume eroded by the trunk channel. The base elevation of thefans for basin #06 and #47 is assumed to be equal to the min-imum elevation of the fan, since they drain into the mainchannels at approximately the same elevation. No majorsigns of incision or undercutting are visible for these fans.At the outlet of basins #24 and #26, no fans have been built.Additionally, bedrock-covered valley bottoms and lakeswithin the basins having no or only very small lake deltassuggest that sediment transport at valley bottoms is insignif-icant since deglaciation of the basins.[27] Erosion rates and SDRs, which are calculated using

equations (4) and (5), are related to a number of uncertaintiesthat are either associated with the assumptions of the sedi-ment budget concept or the uncertainties of the estimated

Figure 3. Schematic cross section representing the interpo-lation of sediment thicknesses based on the dip α and themaximum sediment thickness dmax of different geomorphicdepositional environments (for more explanation, see text).

Figure 4. Box andwhisker plot of (a) basin area and (b) mean elevation stratified based on basin morphology.For location of the 57 headwater study basins, see Figure 1. Outliers are either smaller than firstquantile� 1.5� IQR (interquantile range) or larger than the third quantile + 1.5� IQR. WhereIQR= third quantile � first quantile. pkw-values represent p-values derived from the Kruskal-Wallis test(see text). pkw< 0.05 gives significant differences between basin types.


7

field data [Hoffmann and Schrott, 2002; Tunnicliffe andChurch, 2011]. The primary uncertainties include theassumption that sedimentation was initiated at the time ofdeglaciation (i.e., basins were sediment free after cirqueglaciers retreated) and the errors of the estimated sedimentthicknesses. While rock fall, debris flow, and fan depositsare very likely of postglacial age [Ballantyne, 2002b; Roedand Wasylyk, 1973], the first uncertainty is mainly relatedwith (i) the unknown age of colluvial creep deposits in basinswith limited glacial erosion (e.g., basins #02 and #06) and (ii)the contribution of abundant till in glacial cirques (basins#22, #26, and #47). This uncertainty is taken into accountby subtracting creep and till deposits from SVbasin in equations(4) and (5), resulting in corrected sediment delivery ratios anderosion rates, which are given by SDR* and ER*. To estimatethe errors associated with the calculation of sediment storagevolume, we assume uncertainties for sediment thickness of25%. The uncertainty of the areal extent of storage units, thebulk density ratio (ρS/ρR), and time since deglaciation (T) are as-sumed to be 10%. We used Gaussian error propagation to cal-culate the errors of SV, SDR, and ER. Additionally, wecalculated sediment budgets for four scenarios of interpolatedsediment thickness. The first and third scenarios (“α=10°,max” and “α=20°, max”) are given by a dip of α =10° and

α =20° (compare Figure 3), respectively, and limited sedimentthicknesses for each form unit given in Table 1. The second andfourth scenarios (“α=10°” and “α =20°”) resemble the firstand third scenarios, but without limited sediment thickness.For more details on the uncertainty assessment, the reader isreferred to the supporting information.

4. Results

4.1. Geomorphometry of Headwater Basins

[28] The 57 headwater basins (Figure 1) used in thisanalysis range in size from 0.2 to 10.8 km2. Of these ba-sins, 24 are classified as noncirque, 23 as glacial cirques,and 10 basins show a more complex pattern with charac-teristics of both types. There is a significant difference(pkw = 0.002) in basin size between basins with noncirque(median 0.9 km2), cirque (median ~ 1.6 km2), and complexmorphology (median ~ 3.3 km2) (Figure 4a). In general,glacial cirques are located in higher elevations in the southand west ranges, with mean elevations of around 2400 m,while the majority of noncirque and complex basins arelocated below 2400 m (Figure 1 and Figure 4b).Significant altitudinal differences (pkw = 0.009), in conjunc-tion with higher hypsometric integrals of noncirque basins

Figure 5. (a) Box and whisker plot of hypsometric integral and (c) drop of ELA for full glaciations strat-ified based on basin morphology. The drop of ELA for full glaciations of the basins is based on a modernELA of 2700 masl and an accumulation area ratio AAR of 0.33. (b) Scatterplot of HI versus basin size.Colors represent basin morphology (green = noncirque, blue = cirque, red = complex).

Figure 6. (a) Box and whisker plot of geophysical relief GR stratified based on basin morphology and (b)geophysical relief versus drainage basin size A. Gray shaded area defines 90% confidence interval of linearregression given by GR = a+ b� log(A), where a and b are regression coefficients.


8

(pkw = 0.009; Figure 5a), indicate a lower necessary drop ofthe recent ELA to fully cover cirque basins by a self-con-fined cirque glacier compared to noncirque and complexbasins (Figure 5c).[29] The mean geophysical relief of basins with noncirque

morphologies (median = 48 m) is significantly smaller(pkw< 0.001) than those with cirque (median=124m) and com-plexmorphology (median=123m; Figure 6a). The plot of meangeophysical relief as a function of drainage area (Figure 6b)shows that differences between basins with noncirque andcirque morphology are insignificant for basin size <1 km2

and increase significantly for larger basins (e.g., 150 m and250 m at 5 km2 for noncirque and cirque basins, respectively).[30] Notable differences in slope gradient distribution and

slope parameters are observed for the three basin types(Figure 7). The slopes of noncirque basins tend to be normallydistributed (Figure 7a) with mean slopes of ~0.66 m m�1

(Figure 7b). In contrast, glacial cirques are characterized by pos-itively skewed distributions, with mean slopes of ~0.75 mm�1.The latter value is similar to the transition of soil/sediment cov-ered hillslopes to bedrock hillslopes at ~0.72 m m�1

(Figure 7c). The skewed distribution of glacial cirques resultsin a significantly larger fraction of areas steeper than 0.72 mm�1 (Figure 7d). In general, glacial cirques exhibit a larger

scatter of slope gradients, which result from the larger propor-tion of flat valley bottoms on the one hand, and the large frac-tion of steep headwalls on the other. Slope distributions ofcomplex basins are also positively skewed and similar to glacialcirques, but with significantly lower mean slopes and a muchsmaller fraction of slopes steeper than 0.72 m m�1.[31] Stream concavities of cirque basins (θ = 0.29) and

complex basins (θ = 0.27) are significantly larger than thosefor noncirque basins (θ = 0.20) (pkw = 0.003, Figure 8a).Notable differences (pkw = 0.027) are also evident for thereference slope, Sr, amongst basin types, with low values(Sr=0.23 and 0.21) for cirque and complex basins and higherslopes (Sr=0.28) for noncirque basins (Figure 8b). In contrast,normalized channel steepness, ksn, is similar in all basin types(pkw=0.678, Figure 8c).[32] Absolute and specific values of debris fan size at the

outlets of headwater basins show significantly larger valuesfor noncirque basins in comparison to glacial cirques(Figure 9). Median values of specific fan sizes for noncirqueand complex basins represent ~15% and ~13% of the basinarea, respectively, and only ~5% in case of the glacialcirques. Figure 9 (inset) demonstrates that fan size is notscale dependent (i.e., no correlation between fan size and ba-sin size), indicating that differences between basin types are

Figure 7. (a) Slope distribution stratified based on basin type and (c) bedrock vs. sediment coverage.Colored polygons in Figure 7a define the range of slope distributions and bold colored lines in Figure 7arepresent the mean slope distribution for each basin type. (b) Box and whisker plots of mean slope and(d) fraction of areal extent steeper than the slope gradient of 0.72 m m�1. This gradient represents thetransition from sediment-covered to bedrock-covered slope gradient.


9

unlikely an artifact of different basin sizes. Furthermore, fanslope of noncirque and cirque basins are similar, suggestingthat differences in the areal extent may represent significantlymore sediment stored in fans at the outlet of noncirque ba-sins. Thus, postglacial sediment yields are likely to be largerhere than in cirque basins.

4.2. Spatial Organization of PostglacialGeomorphic Processes

[33] Substantial differences in the spatial organization ofpostglacial geomorphic processes (i.e., their location anddownstream succession within the headwater basins) arisebetween basins with cirque and noncirque morphology. Asshown in basins #24, #26, and #47 (Figure 10), ridgelinesand neighboring rock walls of glacial cirques are mainlyexposed by bedrock and only very scarcely covered by a sed-iment layer. At the transition between steep rock walls andflat valley bottoms, large talus slopes and debris cones are de-posited dominantly by rockfalls and debris flows (Table 2).Debris flows originate in small hollows in the rock wall oron the apex of talus slopes. The flat valley bottoms are eitherexposed by bedrock (basin #24 and #26) or covered by tilldeposits (basin #47) that were deposited during the late gla-cial or the Early Holocene [Gardner et al., 1983]. In contrastto the distinct pattern of sediment storage in basins withcirque morphology, the spatial organization of storage in ba-sins with noncirque morphology (e.g., #02 and #06) is muchmore heterogeneous (Figure 10). Generally, slopes lowerthan ~0.72 m m�1 are covered by regolith or colluvial de-posits (Table 2). The shallow regolith cover near ridgelinesfades into colluvial creep deposits with increasing distancefrom ridgelines. Colluvial creep deposits increase in thick-ness downslope and in concave topography. Locations ofbedrock outcrops and talus deposits at the foot of bedrockslopes are more strongly controlled by variable rock bedding,which affects rock strength and thus slope gradients in basinswith noncirque morphology [McMechan, 1995]. Slope con-cavities are sources of debris flows that are deposited inthalwegs and in large fans at the basin outlet.[34] The different spatial organization of geomorphic pro-

cesses is also shown by the changing dominance of sedimentstorages (Table 2). Creep and debris flow deposits in the

thalweg and the outlet-fan dominate in noncirque basins.Rockfall and debris flow deposits are most widespread inthe glacial cirques. Basin #47, which has been intensely gla-ciated during the Late Glacial [Gardner et al., 1983] and stillhosts a small cirque glacier, is dominated by till deposits andmoraines in the valley bottom (Figure 10 and Table 2), withminor rockfall and debris flow deposits.

4.3. Postglacial Sediment Budgets

[35] Results of sediment budget calculations for the four sce-narios “α=10°,” “α =10°, max,” “α =20°” and “α =20°, max”are given in Figure 11 and summarized in Table S1 (seesupporting information). In general, sediment volumes and cor-responding erosion rates increase from scenarios with α=10°to scenarios with α =20°. Limiting sediment thicknesses tomaximum depths (scenarios “α=10°, max” and “α =20°,max”) reduces sediment volumes and erosion rates, especiallyfor α=20°. Compared to the geophysical results, sedimentthicknesses for scenario “α =20°, max” seem to be most rea-sonable. Scenario “α =20°” results in very high sediment thick-nesses that are not supported by the geophysical soundings.

Figure 8. Box and whisker plot of channel profile parameter stratified based on basin morphology: (a)concavity index θ, (b) reference slope Sr, and (c) normalized steepness index ksn. The reference slope Sris calculated based on a median basin area Ar = 1.4 � 106 m2. ksn is normalized to a concavity index ofθ = 0.25. See Figure S1 in supporting information for example profiles from mapped basin.

Figure 9. Box and whisker plot of (left) absolute and (right)relative size of the outlet fans stratified based on basin mor-phology. Fan sizes are given for a subset of 37 basins, whichare included in the geomorphometrical analysis. Relative sizeis given by the size of the fan in relation to the basin size. Theinset, which plots fan size versus basin size for the 37 head-water basins, does not show any relationship between theconsidered attributes.


10

Figure 10. Spatial distribution of mapped geomorphic process units of basin #02 and #06, characterized bynoncirque morphology, and basins #24, #26, and #47, which are typical glacial cirques. Blue lines representcontour lines with elevation difference of 100 m. See Figure 1 for locations of the headwater basins.

Table 2. Distribution [%] of Area and Sediment Storage Mass for Each Geomorphic Unit in Each of the Five Mapped Basinsa

Area [%] Sediment Mass [%]

domain # 02 # 06 # 24 # 26 # 47 # 02 # 06 # 47 # 24 # 26

bedrock 33.1 21.8 63.9 55.6 55.1 0.3 0.4 0.6 0.9 1.9diffusive hisllope 31.6 52.4 0.8 8.6 1.6 34.5 50.3 0.5 10.7 1.0rock fall (slides) 12.5 5.1 23.0 21.9 10.1 9.3 4.2 58.5 43.1 33.2snow avalanche 0.0 2.1 0.0 0.3 0.0 0.0 2.8 0.0 0.3 0.0glacial 0.0 0.0 2.7 0.2 19.8 0.0 0.0 8.0 0.4 0.3debris flows 14.3 11.0 9.1 11.3 8.0 26.8 14.5 32.4 44.4 32.7alluvial 0.0 0.0 0.0 0.7 1.1 0.0 0.0 0.0 1.1 1.6fan 8.5 7.5 0.0 0.0 4.2 29.2 27.8 0.0 0.0 29.4lake 0.0 0.0 0.5 1.5 0.0sum 100 100 100 100 100 100 100 100 100 100

aBold numbers indicate values larger than 20% (in the case of area) and larger than 25% in the case of sediment mass.


11

[36] Calculated erosion rates range between 0.19 and0.6 mm a�1 (excluding the rates given by scenario “α=20°”)(Figure 11 and Table S1) and tend to be higher in basins withnoncirque morphology (generally> 0.4 mm a�1) than those inglacial cirques (generally< 0.4 mm a�1). Corrected erosionrates (ER*), based on calculations that exclude tills and creepdeposits, are 30–90% of the uncorrected erosion rates (ER);the latter include all stored sediment types. Differences are es-pecially pronounced in basins #47 and #06, which are charac-terized by a large fraction of tills and creep deposits. Incontrast, sediments in basins #25 and #26 are mainly associ-ated with postglacial rockfall and debris flow deposits, causingonly a small difference between ER and ER* in these basins.[37] Postglacial SDRs range between 0 and 0.34 and

between 0 and 0.56, including and excluding creep and tilldeposits, respectively (Figure 11). Larger sediment volumesstored in the outlet fans of tributaries indicate higher sedi-ment delivery ratios of basins with noncirque morphology(ranging between 0.19 and 0.56) compared to basins withcirque morphology (ranging between 0 and 0.3). Thus,exclusion of creep and till deposits from the calculation inequation (5) increases SDR* relative to SDR-values by a fac-tor of 1.5 (basin #02) and up to 3.5 (basin #47).

5. Discussion

5.1. Glacial Imprint on the Morphology of MountainHeadwater Basins

[38] Geomorphic analysis of 57 headwater basins inKananaskis Country reveals major differences with respectto mean geophysical relief, hypsometric integral, absolute el-evation, and elevation relative to modern ELA (see supportinginformation for uncertainty assessment of basin wide morpho-logical parameters). Significant greater geophysical relief(Figure 6a) and lower hypsometric integrals (Figure 5a) of gla-cial cirques are consistent with previous studies that have used

mean geophysical relief [Montgomery, 2002] and basinhypsometry [Brocklehurst and Whipple, 2004; Sternai et al.,2011] to characterize glacial and fluvial landscapes. The tran-sition from fluvial to glacial landscapes through strong glacialerosion generally results in a modification from V-shaped toU-shaped valleys, with decreasing hypsometric integralsand increasing geophysical relief for basins of similar sizes.These results support significant differences in headwatermorphology in Kananaskis Country that are caused by varyingdegrees and/or mechanisms of long-term (esp. Pleistocene)erosion [Osborn et al., 2006]. Disparities between theseheadwater basins are in good agreement with other studiesinvestigating the link between glacial erosion and topography[e.g., Foster et al., 2008; MacGregor et al., 2000; Sternaiet al., 2011] and are indicative of a higher degree of glacialerosion in cirque basins compared to basins with noncirquemorphology. Differences in geophysical relief and hypso-metric integrals remain after correction for the effect of basinsize (Figures 6b and 5b). Glacial cirques show stronger in-creases of geophysical relief with basin size, supporting ourassumption of more pronounced glacial cirque erosion duringthe Pleistocene. In contrast to Montgomery [2002], we findlarger geophysical relief in glacial cirques compared tononcirque valleys for basin sizes as low as 1 km2, highligh-ting the importance of glacial erosion for the evolution ofmountain landscapes even in small headwater basins in theCanadian Rockies.[39] Impacts of glacial erosion on longitudinal river profiles

are more variable. In contrast to Brocklehurst and Whipple[2002], who studied the morphology of glacial versusnonglacial basins in the Eastern Sierra Nevada (California),glacial cirques in our study site were characterized by muchlarger concavities (θ> 0.3) than basins with noncirquemorphology (θ< 0.3, Figure 8a). According to their results,nonglaciated basins reveal much larger concavities (θ =0.32)than glaciated basins (θ =0.07). This apparent contradictionmost likely results from the different spatial scales.Brocklehurst and Whipple [2002] focused on basins between10 km2 and 36 km2 and defined a critical drainage area forchannel initiation of 105 m2. In our study, 80% of the basinsare smaller than 4 km2 and longitudinal profiles cover flow lineswith contributing area >104 m2. Although most basins show abreak in the slope-area relation at 104 m2, some fraction of thehillslopes might be included. In agreement with Brook et al.[2006] and Sternai et al. [2011], higher concavities of glacialcirques in our study are most likely the result of the transitionbetween steep hillslopes and flat valley bottoms.[40] Despite their distinctive morphology, all headwater ba-

sins were covered either by the Bow or the Kananaskis valleyglacier during the LGM (and certainly during former glacialmaxima). Differences in the rate and style of glacial erosionmust be related to the location with respect to the large valleyglaciers and climatic conditions that govern the formation ofcirque glaciers within the basin. Formation of glacial cirquemorphology (indicated by increasing geophysical relief anddecreasing hypsometric integral) is associated with thesustained occurrence of cirque glaciers that carve the basinsdue to spatially variable erosion rates, rather than the occur-rence of large valley glaciers that completely cover the basinsand imprint a different erosional pattern than local cirque gla-ciers. In agreement with other studies [Foster et al., 2008;Pedersen and Egholm, 2013], formation of glacial cirque

Figure 11. (Top) Calculated erosion rates ER in mm a�1

and (bottom) sediment delivery ratios SDR for noncirque ba-sins #02 and #06 and glacial cirques #24, #26, and #47.Circles and triangles represent values calculated with andwithout maximum sediment thickness, respectively. Bluecolors indicate calculations including creep deposits and tills(e.g., ER and SDR), green colors indicate calculations withoutcreep and till deposits (ER* and SDR*). Bars show errormargins as calculated by the Gaussian error propagation. Allvalues are summarized in Table S1.


12

morphology is more likely in high-elevation headwater basinsthat require only a small depression of the present-day ELA(Figure 5c) to be covered by self-contained, cirque glaciersfor a sustained period. In contrast, low elevation basins requirea substantial decrease of the ELA to form cirque glaciers (gen-erally>600 m) and thus experienced shorter periods of glacialice coverage and less erosion by local cirque glaciers.[41] Calculated ELA-depression relates to the topography

of the tributary basins rather than to mapped glacier outlinesand the ice surface [Ramage et al., 2005]. Therefore, ELA-depressions required for the formation of a self-containedcirque glacier in headwater basins are very likely to beoverestimated. However, we believe that the calculateddifferences suggest greater probabilities of cirque basins tobe covered by cirque glaciers over a longer time period withmore effective glacial erosion than in noncirque basins.Additionally, altitudinal differences are in good agreementwith observations from other mountain ranges [e.g., Egholmet al., 2009; Sternai et al., 2011], which suggest that strongglacial headwater erosion is limited to elevations around andabove the long-term values of ELA that act as local base levelsfor glacial erosion.[42] Furthermore, low elevations of noncirque headwaters

in the northern part of the study site indicate that these basinsmay have been covered/filled by glacial ice of the trunkvalley glaciers. For example, the Bow valley glacier reachedelevations up to ~2100 m in the northern part of the studyarea [Rutter et al., 2006] and therefore easily filled lowelevation basins in this location. Basins that were more proneto erosion by trunk glaciers may have experienced a differentpattern of glacial erosion (e.g., erosion of tributary ridgesrather than deepening and widening of the basin) than basinscontaining local cirque glaciers. This idea is also supported bythe mapping of geomorphic processes and deposits: morainesand till deposits are limited to basins with cirque and complexmorphology (e.g., basin #24, #26, and #47), while tributarybasins with noncirque morphology are lacking substantialamounts of glacial sediment (e.g., basin #02 and #06).[43] Even though a detailed reconstruction of the glacial

extent of the LGM glaciers is not available for the study area,and causes of different basin morphology remain somewhatspeculative, evidence based on basin morphology and loca-tion (e.g., with respect to elevation and location relative tothe main trunk glacier) are consistent with differing degreesof erosion by cirque glaciers in these basins.

5.2. Controls on Postglacial Erosion Ratesin Mountain Headwaters

[44] Erosion rates are frequently related to slope gradientsin different geomorphic settings [e.g., Montgomery, 2001;Safran et al., 2005; Strahler, 1950]. At low slopes, diffusivesediment transport increases linearly with slope gradient, butas slopes approach the threshold gradient erosion ratesincreases notably with only small changes in slope gradient[Carretier et al., 2013; Montgomery and Brandon, 2002].The explanation proposed for this threshold hillslope beha-vior is either the limiting stability of the regolith, which setsthe maximum achievable slope in soil-mantled landscapes[Carson and Petley, 1969], or the increasing frequency ofbedrock landslides in active orogens, as a consequence ofchannel incision and subsequent hillslope steepening beyondthe threshold gradient that is determined by the bedrockstrength [Burbank et al., 1996; Schmidt and Montgomery,1995]. An important corollary of the threshold hillslope con-cept is the shift (i) from transport-limited to weathering-limitedprocesses, as erosion rates exceed the weathering rates [Binnieet al., 2007] and (ii) from low to high bedrock coverage[Norton et al., 2010a]. Thus, hillslope gradient and bedrockcoverage on steep slopes are indicative of high erosion ratesin many mountain environments.[45] Geomorphic mapping in the Kananaskis Country re-

vealed a transitional slope gradient from gentle hillslopes cov-ered by soil and/or sediment (e.g., transport limited) to steepbedrock hillslopes (e.g., weathering limited) at 0.72 m m�1

(Figure 7c). Under similar lithological conditions, steepermean gradients (Figure 7b) and significantly larger fractionsof weathering-limited hillslopes (Figure 7d) suggest highertopography-driven erosion rates in glacial cirques than in dom-inantly soil/sediment-covered headwater basins with noncirquemorphology. Erosion rates derived from the postglacialsediment budgets of five headwater basins in KananaskisCountry do not correlate with mean basin slope, but ratherdecrease with maximum slopes (Figure 12a), the fraction ofslopes steeper than 0.72 mm�1 (Figure 12b), and with bedrockcoverage (Figure 12c). These relationships are significant, de-spite large uncertainties of the budget-derived erosion rates.This is also true when erosion rates ER* without colluvialand till deposits are considered (not shown). The assumptionthat the chosen headwater-fan systems are closed in terms ofpostglacial sediment fluxes results in an underestimation of

Figure 12. Correlation of erosion rates (based on scenario “α = 10°, max”) with (a) mean and maximumbasin slope, (b) areas with slopes steeper than 0.72 m m�1, and (c) bedrock coverage.


13

ER and ER*, since we cannot fully exclude that fine sedimentbypassed the lowermost boundaries of the outlet fans. Thisespecially applies to basins with noncirque morphology, sincethese basins host a large fraction of fine-grained colluvialsediment, suggesting that differences between erosion rates inbasins with cirque morphology and noncirque basin are likelyto be underestimated. However, there was no significant fieldevidence of downstream coarsening, which suggests a pre-ferential export of fine-grained sediment (see supporting infor-mation for details on the uncertainty assessment).[46] Lower erosion rates of steep bedrock slopes compared

to gentler soil-covered slopes (as derived from equation (4))may have two possible explanations. First, low denudationrates despite extensive high and steep rockslopes in glacialcirques are not in accordance with the nonlinear relationbetween mean slope gradient in basins and erosion, assuggested by the threshold hillslope concept. In fact, thesteepest bedrock hillslopes, which are mainly composed ofwell-drained massive limestone, have been relatively stable interms of gravitational mass movements during the postglacialand Holocene, as supported by the limited occurrence ofmassive rock failures in the study basins [Cruden, 2003].This observation is especially true for reverse dip slopes of sed-imentary rocks (which dominate in the mapped basins), wherebedding planes dip into the slope with a dip direction oppositeto the surface. In contrast to reverse dip slopes, dip and over dipslopes (i.e., bedding planes dip in a similar direction to slopesurface and slopes are equal or larger than the dip slope) inKananaskis Country have been locations of massive slopefailures mobilizing large amounts of rock and sediment[Cruden and Eaton, 1987; Hu and Cruden, 1992]. However,the impact of massive slope failures in our empirical sedimentbudgets is of minor importance relative to the overall denuda-tion rate (large deep-seated rock slides generally account forless than 30% of the rockfall domain).[47] Second, high erosion rates in noncirque basins might

be partly attributed to higher retention of soil moisture onhillslopes that are covered by a thin layer of soil, regolith orcolluvium, which in turn increases chemical weathering anderosion on soil-covered slopes. Thus, erosion rates on slopeswith a “stimulating” soil/sediment layer might be higher thanon well-drained bedrock slopes [see for instance Gabet andMudd, 2009]. If this is true, strong glacial erosion in cal-careous mountains may lead to the formation of hillslopesthat are too steep to support the formation of a soil/sedimentlayer and, thus, may experience lower erosion rates than soil/sediment-covered hillslopes that likely dominate in basinswith lower glacial erosion rates [Martin, 2000; Martin andChurch, 1997]. However, to further support this hypothesis,more field data linking bedrock coverage and erosion ratesare required.[48] To summarize, even though postglacial erosion rates

are independent of mean basin slope as suggested forthreshold hillslopes [Burbank et al., 1996; Montgomeryand Brandon, 2002], we argue that hillslope topographyin Kananaskis does not represent threshold hillslopes andsediment production at over steepened hillslopes is rela-tively low. Thus, it can be questioned whether the exponen-tial relationship between slope gradient and soil erosionas suggested by the threshold hillslope concept [Binnieet al., 2007; Montgomery and Brandon, 2002] is appropri-ate for the prediction of erosion rates in sedimentary,

calcareous mountain environments, such as the CanadianRocky Mountains. In these lithologies, a better representa-tion of bedding structures and their impact on deep-seatedbedrock landsliding, rockfalls, and debris flows is required[Cruden and Hu, 1988]. As indicated by our results, simplerelationships with the degree of Pleistocene glacial erosionand postglacial adjustment are not found [see e.g., Dadsonand Church, 2005].

5.3. Spatial Organization of Geomorphic Processesand Disconnectivity of Mountain Headwaters

[49] Variable sediment transport conditions in differentbasin types are attributed to different slope-area relationsexpressed by θ and Sr (Figure 8 and S1). Higher concavitiesof glacial cirques in our study are most likely the result of thetransition from steep rock walls and flat valley bottoms. Thelatter is supported by low median reference slopes, Sr, whichrepresent the channel steepness at certain reference basinssizes (here Ar = 1.4� 106 m2). Values of Sr decrease from0.28 in noncirque basins to 0.23 in glacial cirques and 0.21in complex basins. Assuming similar precipitation in allbasins and applying a mean value of θ (0.19 and 0.3) and amean value of Sr (0.23 and 0.28) in cirques and noncirquebasins, respectively, greater Sr values in noncirque basinsresult in greater stream power compared to glacial cirquesat basin sizes larger than ~3� 104 m2 (Figure S1f). Thus,all else being equal, glacial cirques are less able to exporttheir sediment than steeper basins with noncirque morphol-ogy at a similar contributing basin area [Bull, 1979].[50] The low ability of glacial cirques to export their sedi-

ment, in turn, is expressed by (i) different spatial organizationof geomorphic processes and their deposits (Figure 10),(ii) notably lower SDRs (Figure 11), and (iii) significantlysmaller sizes of alluvial fans at the outlets of headwater basins(Figure 9). Sediment deposits in noncirque basins (primarilycreep deposits and debris flows, Table 2) increase in thicknessfrom ridgelines to the thalweg, and from convex to concaveslope positions. In contrast, most sediment in glacial cirques(primarily rockfalls and debris flow deposits, Table 2) residesat the transition between steep rock walls and the flat valleybottom. This organization has major implications for hill-slope-channel coupling of postglacial sediment cascades. Inbasins with noncirque, V-shaped morphology (e.g., basin #02and #06), hillslope processes transport sediment to the thalweg,where debris flows and fluvial processes convey it to outlets oftributary basins. High sediment delivery in these basins is dem-onstrated by postglacial SDRs of ~ 30% (see supporting infor-mation for detailed uncertainty assessment of given SDRs). Incontrast, glacial cirques with U-shape morphology (e.g., basins#24, #26, and #47) relocate sediment from hillslopes/rock wallsto valley bottoms, where it remains as a deposit during thepostglacial without being delivered to the basin outlet. The lim-ited sediment delivery is given by zero SDR values for basin#24 and #26 and SDR=0.08 for basin #47. This finding issupported by significantly smaller fan sizes at the outlets of gla-cial cirques (n=13) compared to noncirque headwater basins(n=17) (Figure 9). It remains unknown how long the trappingefficiency of glacial cirques remains high, and how values ofSDR will change as the cirque fills up with sediment in thefuture. However, without a significant increase of the powerto transport coarse sediments (e.g., due to glacial erosion atthe beginning of the next, potential glaciation) or significant


14

weathering of the deposited sediment, it is likely that cirque-SDRs will remain low for several 104 years.[51] In summary, sediment transport conditions as expressed

by slope-area relations (Figure 8 and Figure S1f) and the spatialorganization of processes are in agreement with low sedimentdelivery and long-term sediment storage in glacial cirques.Although these basins represent the steepest part ofmountain environments and potentially produce largeamounts of postglacial sediment, only a limited fractionof postglacial sediments is exported from glacial cirquesduring the postglacial.

6. Conclusion

[52] Herein, we have presented evidence regarding the im-pact of glacial erosion on the morphology and postglacial sed-iment transport for headwater basins in Kananaskis Country,Canadian Rocky Mountains. This study provides a quantita-tive link between postglacial erosion rates, sediment deliveryratios, and headwater morphology. Notable differences in thehypsometry, geophysical relief, and elevation amongst the57 headwater basins point to variable landform histories.Although all basins have been covered by glacial ice duringthe LGM, significantly greater geophysical relief (90–180m)and lower hypsometric integrals (0.42–0.5) suggest morepronounced erosion by local cirque glaciers in high-elevationbasins in the southern part of Kananaskis Country than inlow-elevation basins in the northern part.[53] Distinct slope distributions of glacial cirque basins

point to higher erosion rates compared to headwater basinswith noncirque morphology. However, this observation isnot supported by erosion rates derived from postglacial sedi-ment budgets, which show rates between 0.4–0.6 mm a�1

and 0.19–0.4 mm a�1 for noncirque and cirque basins,respectively. These findings suggest a more complexrelationship between hillslope erosion and slope gradient incalcareous mountain environments than implied by thethreshold hillslope concept.[54] Strong contrasts between steep hillslopes and gentle

valley bottoms impose different transport conditions and lessefficient sediment delivery from glacial headwater basinsthan those that escaped strong erosion by local cirqueglaciers. These findings are supported by much smaller fansizes at the outlet of glacial cirques and significantly lowerpostglacial SDR (typically ~0.3 for noncirque basins and<0.2 for glacial cirques).[55] Our findings suggest that the formation of glacial

cirques during the Pleistocene has major implications forpostglacial sediment sources and fluxes in glaciated, alpineenvironments. Generation of relief by glacial erosion modifiesthe spatial organization of geomorphic processes anddecouples glacial headwater basins from main river valleys.Although small headwater basins represent the steepest partof mountain environments and potentially produce largeamounts of postglacial sediment, only a limited fraction ofthese sediments is exported from glacial cirques. The majorityremains in storage under postglacial climatic conditions anddoes not affect large-scale mountain river systems.Therefore, no major sediment removal is expected until cli-matic or tectonic conditions modify basin topography in sucha way that a higher connectivity of headwater basins emerges.

[56] Acknowledgments. This work was supported by NSERCDiscovery Grants to E.A. Johnson and Y.E. Martin. We thank theBiogeoscience Institute of the University of Calgary for logistic and labsupport at the Barrier Lake field station. Simon Brocklehurst, AnnRowan and two anonymous provided helpful comments on an earlierversion of this manuscript.

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