Bedload Transport and Channel Bed Changes in the Proglacial Skeldal River, Northeast Greenland

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Bedload Transport and Channel Bed Changes in the Proglacial Skeldal River, NortheastGreenlandAuthor(s): Tim StottSource: Arctic, Antarctic, and Alpine Research, Vol. 34, No. 3 (Aug., 2002), pp. 334-344Published by: INSTAAR, University of ColoradoStable URL: http://www.jstor.org/stable/1552492 .

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Arctic, Antarctic, and Alpine Research, Vol. 34, No. 3, 2002, pp. 334-345

Bedload Transport and Channel Bed Changes in the Proglacial Skeldal River, Northeast Greenland

Tim Stott Physical Geography and Outdoor Education, School of The Outdoors, Leisure and Food, Liverpool John Moores University, I. M. Marsh Campus, Barkhill Road, Liverpool, L17 6BD, U.K. [email protected]

Introduction Studies of bedload transport processes in arctic proglacial

fluvial environments are outnumbered by those from temperate alpine regions where steep channel gradients, unconsolidated sediments and relatively high magnitude diurnal fluctuations in discharge result in dynamic braided channel systems (Maizels, 1979, 1983, 1995). These proglacial environments are associated with high rates of bedload transport, deposition within the chan- nels leading to bar growth in some areas, and associated bank erosion and channel change in others. The steep channel slopes give rise to high excess shear stresses, and the readily available supply of unconsolidated sediment, combined with the large diurnal fluctuations in discharge, means that sediments are trans- ported frequently and so they remain loose rather than becoming imbricated. The assumption that these conditions apply to all proglacial fluvial environments, however, may be unfounded since most studies seem to have been focused in alpine environments in Europe (e.g., Fenn and Gumell, 1987; Maizels, 1979; Gumell et al., 1988; Warburton, 1992, 1994; Gumell, 1995; Lane et al., 1995) and in North America (Fahnestock, 1963; Ferguson et al., 1989; Goff and Ashmore, 1994), with relatively few proglacial studies in Arctic areas. Most studies have fo- cussed on subarctic regions: Nicholas and Sambrook Smith (1998) reported on bedload transport rates on the Virkisa River in southern Iceland and Ashworth and Ferguson (1986) in Lyngsdalselva River in arctic Norway, and Church and Gilbert (1975) presented data for rivers in Alaska and Scandinavia. In the Antarctic, Mosley (1988) measured bedload transport rates on the Onyx River, as did Inbar (1995) on Deception Island. Studies of bedload transport in locations north of the Arctic Cir- cle are relatively few: Church (1972) carried out extensive studies of sandurs on Baffin Island and Busskamp and Hasholt (1996) monitored bedload transport processes on Ammassalik Island in Southeast Greenland.

334 / ARCTIC, ANTARCTIC, AND ALPINE RESEARCH

Abstract This study presents data on bedload transport rates for the proglacial Skeldal River in northeast Greenland monitored during August 1998. The study adds to a very limited data set from which we can assess bedload transport rates in true Arctic environments. Bedload transport rates were determined using a sediment budget approach based on cross-section resurvey, supplemented by "at-a-point" bedload measurements using a Helley-Smith bedload sampler. Bedload transport rates were low and ranged from 0.02 to 0.04 kg m-1 s-1. These rates fall within the extremes reported for arctic and subarctic environments. Bedload sampling conducted through flood events suggests that there is a broad (but not significant) relationship with discharge. Maximum bedload transport rates derived using a sediment budget approach based on cross-section resurvey were found to be 50% lower than maximum rates measured "at-a-point" using a Helley-Smith bedload sampler during meltwater flood events. Bedload transport rates derived in this study vary both spatially and temporally by up to an order of magnitude. Upstream bedload supply parameters are believed to be more influential than specific stream power in con- trolling the transport patterns and changes in channel bed elevation observed.

Information about the sediment load of proglacial streams and rivers is becoming increasingly important in an applied con- text as glacial meltwaters have been developed for hydroelectric power generation (0strem, 1975). An understanding of the con- trols on magnitude and frequency of meltwater flows in proglacial rivers is crucial for the prediction of channel morphology and sedimentology, sediment dynamics and yields, as well as for the interpretation of the sedimentary record of formerly glaciated regions. Previous studies from true arctic proglacial rivers tend to report lower rates of sediment transport and greater channel stability than their alpine equivalents. This study adds to the very limited data set available to date to allow assessment of bedload transport and channel change for arctic proglacial environments.

Aims The aim of this study was to use two independent tech-

niques for measuring bedload transport, both successfully uti- lised in previous studies, to assess the bedload transport rates in a reach of the Skeldal River, northeast Greenland, during part of the melt season in August 1998. The objectives were to:

(1) monitor bedload transport rates in a true proglacial arctic river directly through diurnal flood cycles using a Helley- Smith bedload sampler to gain an insight into the dynamics at the diurnal timescale, and

(2) to assess bedload transport rates and channel change indirectly using a sediment budget approach based on repeated channel cross-sectional surveys over a 3-wk period during the melt season.

Field Area Greenland is the largest island in the world, with an area

of 2,175,600 km2. About 84% of the area is covered by the Greenland Ice Sheet, with altitudes extending to more than 3000

? 2002 Regents of the University of Colorado 1523-0430/02 $7.00



m a.s.l.; nearly everywhere the ice sheet is surrounded by a strip of land with a maximum width of about 200 km. Because glacial erosion is one of the most sediment productive types of erosion (Hasholt, 1996), the expansion or contraction of the Greenland Ice Sheet, and of the ice-free coastal zone, has a major impact on the amount of sediment delivered to the surrounding oceans. With increasing interest in studies of climate change using ice and ocean bottom cores, knowledge of the sediment transport from Greenland is of major importance in interpreting deltaic and offshore sedimentation patterns.

In a review of sediment transport in Greenland, Hasholt (1996) summarized the findings of 11 studies which have been conducted in Greenland to date: nine of these were conducted in rivers draining to the southwest coast (the most populated part of Greenland). On the east coast, the first sediment transport studies in Greenland were initiated in 1972 on Ammassalik Is- land (Hasholt, 1976) and further studies have continued in that area (Hasholt, 1992, 1993, 1994; Busskamp and Hasholt, 1996). However, apart from these intensive studies on Ammassalik Is- land, the only other work on sediment transport reported for East Greenland is from the Zackenberg area (Jakobsen, 1992; Rasch et al., 2000) and for the Skeldal River in Scoresby Land (Stott and Grove, 2001).

The Skeldal River (72?15'N, 24?15'W) is situated in Scores- by Land, northeast Greenland (Fig. 1A). The igneous and metamorphic complex of the Staunings Alps is overlain by Paleozoic sediments and modem glacial and fluvial materials. A major north-trending fault, the Staunings Alps fault, separates the Ca- ledonian igneous-metamorphic complex and pre-Carboniferous sediments of west Skeldal, where peaks rise to approximately 2000 m, from the Permo-Carboniferous and Tertiary-Quaternary sediments which rise to about 1100 m east of the fault (Washburn, 1965; Lasca, 1969). The regional climate is polar-arid with annual precipitation around 300 mm, of which 200 to 250 mm is in the form of snow. Annual temperatures range from a minimum of -45?C to a maximum of +23?C. The spring thaw usually begins in May, and the first winter snows fall during September.

Skeldal is an east-trending valley extending about 30 km from its head in the glaciers of the Staunings Alps, to its mouth on Kong Oscars Fjord some 8 km northwest of Nyhavn. At its mouth the Skeldal drains an area of 560 km2 and the drainage area upstream of the study reach shown in Figure 1A is estimated to be 220 km2, around 44% of which is glacierized. Figure lB shows the study area (140 m above ordnance datum) in relation to the Skelbrae glacier and Figure 1C shows the location of monumented cross sections, bedload sampling point, and stage recorder. Mean channel slope of the study reach was estimated to be 0.008 m m-1. Figure 2 shows the study reach with the Staunings Alps in the background.

The Helley-Smith sampler was deployed at 2-m intervals (10 verticals in all) along Section F (Fig. 1C) between 1700-1745 on 5 August when discharge was estimated to be fairly constant at 6.0 m3 s-1 (Fig. 4) in order to assess variability in bedload transport rate across the channel. Bedload samples were subse- quently dried and weighed.

The second study aim was met by establishing six channel cross sections, A to F on Figure 1C, which allowed identification of the nature and extent of channel change in the reach. Indirect estimates of bedload transport and channel change were made by repeat survey of the six monumented sections from which channel planform and cross-sectional changes could be quantified and estimates of bedload transport rates determined. The reach was approximately 100 m long with cross sections spaced every 20 m on average. Sections were surveyed using a quick set level and staff and channel elevations were recorded at fixed cross stream intervals of 0.5 m on five occasions over the study period. Sec- tions ranged from 28 to 41 m in length. The first survey was made on 29 July before a discharge record was established, with subsequent surveys on 1, 6, 14, and 21 August (Fig. 3).

Comparison of these repeat surveys permitted identification of areas of erosion and/or deposition at each cross section over the intervening time period. These data were later used to con- struct a sediment budget for the 100-m channel reach and to determine downstream variations in estimates of bedload transport rates. This method of calculating rates and patterns of channel change has been widely adopted in recent studies of sediment transport in fluvial environments (Ferguson and Ashworth, 1992; Goff and Ashmore, 1994; Martin and Church, 1995; Nicholas and Sambrook Smith, 1998). The method allows volumetric changes in sediment storage to be determined for a given time period based on the geometry of the survey network and measured amounts of net erosion or deposition at each cross section. The rate of change of volumetric sediment storage over time may then be equated with a downstream sediment transport rate increment using the following simple relationship (Ferguson and Ashworth, 1992):

A Qs/Ax = - V(TAx) (1)

where V is the volumetric change in sediment storage between two cross sections separated by a distance A x that occurs over time T, and A Qs is the sediment transport rate increment over that length of channel (in m3 s-1). The volumetric change in sediment storage between two sections can be determined from

V = A Ax (2)

where A is the net cross-sectional area of erosion or deposition. If the sediment transport rate is known at one section along the study reach then transport rates can be calculated for the whole reach by applying the sediment budget equation

Qi- = ei-, + A Q5 Methods

Direct bedload transport rate measurements were taken ear- ly in the study during a 48-h intensive sampling program from 1200 on 1 August to 1130 on 3 August 1998 (sampling point shown on Fig. lB, sampling duration indicated on Fig. 4). Sam- ples were taken using a modified Helley-Smith bedload sampler (Helley and Smith, 1971) with a 0.1 m X 0.1 m square orifice, 30? flare angle, and net bag with 0.25-mm gauge. Samples were collected every half hour throughout the 48-h sampling period by wading in from the bank and placing the sampler on the bed for 30 s and then removing it. Water velocity over the same 30-s period was also measured at the bedload trap orifice by fixing a Braystoke flow meter to the Helley-Smith wading rod.

(3)

where Q5 is the volumetric bedload transport rate and A Q, is the change in transport rate between sections i and i - 1 which is calculated in equations 1 and 2. Where no independent measurements of sediment transport rates are available, estimates of transport rates throughout the study have been made by applying one of two conditions. These are that sediment transport rates should be non-negative at all sections (Griffiths, 1979) or that the rate of sediment transport is zero at one of the boundaries of the study reach (Mclean, 1990; Nicholas and Sambrook Smith, 1998). It should be noted that this method gives only a minimum estimate of the transport rate as there is the possibility of material moving through the reach that is not recorded by the small changes in bed elevation. Additional uncertainty in estimates of transport rates

T. STOTT / 335



FIGURE 1. A: Study location showing Mestersvig airstrip, Skeldal and the study area. B: Plan of study area in relation to Skelbrae glacier snout. C: Large-scale plan of study reach showing location of cross sections, bedload measurements, and stage recorder.

derived using this approach will result from the averaging of scour and fill both temporally (over the interval between surveys) and spatially (across a section). However, despite these limitations the sediment budgeting approach to estimating rates of bedload transport may offer more reliable estimates of long-term bedload trans-

port rates than the instantaneous estimates derived using the point- based Helley-Smith sampler.

In the absence of any independent data quantifying bedload transport for the whole study period, the 30-min Helley-Smith sampling being undertaken only for one 48-h period (1-3 Aug.),

FIGURE 2. View of study reach at low flow looking downstream. The field worker for scale is surveying section F.




E E

n- = a ._ r

- - - 0) 0) 0) 0) 0) 0 0) 0) 0) 0) 0) 0 0

? CM (D co o < < < < < < < < < < < < < C\ C\ \ \ 0 CO M ' CO LO N.l 0 .- C) L N. a) CO) 0)

0 0 0 0 0 C- - - - - 0 CM CM

the condition of non-negative transport is applied which for this reach is also equivalent to the assumption of zero transport at the downstream channel cross section. As with other studies that have applied one of these two conditions (e.g., Griffiths, 1979; Mclean, 1990; Ferguson and Ashworth, 1992; Goff and Ash- more, 1994; Martin and Church, 1995; Nicholas and Sambrook Smith, 1998), some uncertainty in the estimates of transport rates must result from this assumption. Bedload fluxes estimated from such volumetric changes are therefore likely to be lower than those estimated from bedload samples. However, the extreme stability of the channel banks (as identified in repeat cross-section surveys) and the imbricated nature of the channel bed sediments (observed during surveying) are factors that support the application of the non-negative transport condition. It is concluded that the errors associated with this assumption will be small in comparison with other inaccuracies associated with the spatial and temporal averaging of erosion/deposition areas as outlined above, and it does allow comparison of transport rates estimated from this method, to be compared with those derived from the Helley-Smith sampling exercise.

A simplified planform map of the study site was obtained using a plane table (Fig. lB). The average gradient of the reach was 0.008 m m-'. Grain size data were collected on bar and bed surfaces by sampling 100 clasts at each location according to the method of Wolman (1954). The D50 of the bar surfaces ranged from 22 to 65 mm with a mean of 43 mm. Specific shear stresses were estimated along each section at 1-m intervals by measuring depth and flow velocity, using a standard Braystoke type current meter averaged over a 30 s period at 0.6 of depth, with all measurements being conducted between 1600-1930 on 5 August when discharge was estimated to have been fairly constant at around 6.0 m3 s-' (Fig. 4). Estimates of specific stream power (w) were based upon surveyed channel bed and water surface elevations, flow depth and mean flow velocity at 0.6 depth:

w = phSv (4)

where p is the fluid density, h is the flow depth, S is the bed slope, and v is the flow velocity. An unbroken discharge record was maintained from 1 to 21 August 1998. The gauging site was established at a stable 45-m-wide cross section at the downstream end of the study reach some 2 km from the Skelbrae Glacier (see Fig. 1C). Stage was recorded using a Druck pressure transducer (fixed to a stage board) with readings being taken at 1-min intervals and the 5-min mean was logged by a Grant's Squirrel 1000 series data logger downloaded on five occasions to an Apple portable computer. Calibration of the pressure trans-

FIGURE 3. Daily rainfall totals in Skeldal from 20 July to 26 August 1998 with channel survey times indicated.

ducer was undertaken during the 48-h intensive sampling period from 1200 on 1 August to 1130 on 3 August when the river stage and velocity were gauged at half hourly intervals. The channel margins at the gauging site maintained a stable cross section for the duration of the study. Velocity measurements (n = 96) were taken using a Braystoke flow meter at 0.6 of the depth during the full diurnal range in stage from 1 to 3 August, but larger rainfall induced events, such as that on the afternoon of 11 August, were not gauged and discharge for these events is estimated by extrapolating the linear relationship obtained between 1 to 3 August. Channel bed elevation and water depth was surveyed at 0.5-m intervals using a quickset level and staff on five occasions during the study period, and velocity at 0.6 m depth was measured at 2-m intervals (22 verticals) across the channel. River discharge for each profile was calculated using the velocity area method (Herschey, 1978). A linear stage-discharge relationship (r2 = 0.89) was most appropriate for the data. Daily rainfall totals were collected by a standard rain gauge lo- cated near the base camp (Fig. lB).

Results

ESTIMATES OF BEDLOAD TRANSPORT DERIVED FROM THE SEDIMENT BUDGETING APPROACH

Figure 1C shows a simplified planform map of the study site. Figure 3 shows the rainfall pattern from 20 July to 26 Au- gust and Figure 4 shows the discharge record for the period 1 to 21 August with the times of surveys and bedload sampling indicated. Figure 5 shows the six channel cross sections for the five survey dates and has specific stream power measured at moderate flow (-6.0 m3 s-') superimposed. There were no obvious changes to channel planform or bar positions during the study, though repeat surveying of sections revealed that changes in channel bed levels had occurred in the deepest parts of the channel (Fig. 5). As might be expected, the changes were greatest in areas of the channel where stream power was highest, notably between 3 and 10 m from the south channel bank along section A where the bed level fluctuated by up to 0.5 m over the study period. The bed elevation between 5 and 9 m along section B fluctuated by up to 0.35 m over the study period. In section C bed elevation fluctuated by up to 0.3 m over the study period between 6 and 15 m along this section. In section D the bed elevation fluctuated by up to 0.25 m over the study period between 10 and 20 m along the section. These changes in bed elevation are of the order of 5-10 times larger than the D50 (0.065 m b-axis) revealed from Wolman sampling of bar surfac-

T. STOTT / 337



1-5 August 1998 [- Discharge 50

45 SURVEY 2

40

35

30 [Half hourly bedload sampling period J. 0' 25

20

15s 1 0

0 12.00:00 0.00:00 12:00:00 0:00:00 12:00:00 0:00.00 12:00':00 0:00':00

1/8/98 2/8/98 2/8/98 3/8/98 3/8/98 4/8/98 4/8/98 5/8/98

5-9 August 1998 - Discharge

50 -

45 40 35

E 30 0 25

20 Is5 10 5

12:00:00 0:00:00 12:00:00 0:00.00 12.00:00 0:00:00 12:00:00 0:00:00 5/8198 8/8/98 6/8/98 718/98 7/8198 818/98 8/8/98 9/8/98

9-1 3 August 199- Discharge

45

40

35

30

&25

A20 15

105

12:0000 0:00:00 12:0000 0:00:00 12:00:00 0:00:00 12:0000 0:00:00 9/8198 1018198 10/819 11/8/98 1181898 12/819 12/8/98 138/98

13-17 August 1998 - Discharge

45

404 35

30

2s

20

1 5

10

12:00:00 0:00:00 12:00:00 0.00:00 12.00:00 0:00-00 12:00:00 0:00:00 13/8/98 14/8/98 14/8/98 15 /8/98 15/8/98 16/8/98 16/8/98 17/8/90

1(121 AUgUSt 1996 Discharge

40

35

E30

20 215

)1 0

12:00:00 0:00:00 12:00:00 0:00:00 12:00-00 0:00:00 12:00.00 0:00:00 17/8/98 18/8/98 18/8/98 19/8/98 19/8/98 20/8/98 20/8/98 21/8/98

FIGURE 4. Skeldal River discharge record for 1-21 August 1998 indicating the timing of monumented cross-section surveys and intensive bedload sampling.




Section A

1 11 r A', .t -s-^*' -I'- ^ . .. A. P.:'.. " 1 .==~. , -7l N-:; ....~-~ ~ 2..

... .. I . ..... , 3 8 13 18 23

Distance across channel (m)

_ ^ Stream Section B pow er

29-Jul 3.5 - 8 E

j 3 6 _ " -- 31-Jul

o ........6-Aug

nc .5,.m m ^.,,,,, UJ, ,,,, ,,^ ,0 -C ---14-Aug

3 8 13 18 23 28 21-Aug Distance across channel (m)

Section C Stream power

29-Jul 3.5 8 E

? 3 6- | ;;, --- 31-Jul

E '

2.5 4 B 2 2

- 1.5 ..onr T....,^^^ ,,0 ---14-Aug

3 8 13 18 23 28 33 38 21-Aug


~ Stream Section D power

_ 3.5- - 8 29-Jul >u S

. 3 6 , , ---- 31-Jul

", E 2.5 " , I 4

r 2 '

ijIi.(1.- ^ ! 6-Aug

t? 1.5 ........... 6 MM? ,,AA A f A AA A, ,fP.,,,. 0 <fl --- 14-Aug

3 8 13 18 23 28 33 38 43 .....e21-Aug


- Stream Section E power

3.5'o~~~~~~~~~~~~~ 8~~2 ~29-Jul 3.5- - -

- 8 E

4) 3 6 2 7 ---- 31-Jul

4 E 2.5 ,. 4 2

llx^ 2\ . ~~Ii 11 20 a.0,&l~C 6-Aug

1.5 0 ...) 0. w 6-14-Aug 3 8 13 18 23 28 33 38 43


Stream Section F power

29-Jul 3 2.5 -8

.3 2

,.- -6 . --- 31-Jul

1 lII I I I I I I I JI eI I I I I I l . o O --14-Aug 3 8 13 18 23 28

Distance across channel (m) 21-Aug

FIGURE 5. Channel cross- section changes and stream

power as measured at 1800 on 5 August for six monumented cross sections (A to F) on the Skeldal River, August, 1998.

T. STOTT / 339

z

2) -j

a '

c c

3.5

3

2.5

2

1.5

8 E 6 -

4

0 .n

r _S trea m- power I

-29-Jul

--- 31-Jul

....... 6-Aug

--- 14-Aug

- 21 -Aug 28

I I I I I I II I I I II



A 12 ? 29

0.5 - - A A

~.. - .-----1 A U E - Auc

? -0.5 6 -- - A

-1 - - 14 21

-1.5 - m

Distance (m)

B

I-

0.9

0.8

0.7

0.6

0.5

0.4

0.3

0.2

0.1

0

/ - - X-

? - * .....-- Al - -

50

Distance (m)

2.5

_ '

o

E 1.5 C E

0 C -

1 a.)

o.5 0

2

100

FIGURE 6. A: Downstream variations in erosi tion as revealed by changes in area of six cros four time periods. The sections are labelled A tc right (upstream to downstream-see Fig. 1C). I changes in section-averaged bedload transport re using the sediment budgeting approach for fou The sections are labelled A to F from left to rig, downstream-see Fig. 1C).

es. Grain sizes in the deepest parts of the channel than this, so error limits of at least + 0.1 m she to these bed elevation changes.

Figure 5 shows the channel thalweg swingin, bank in Section A towards the center of the cha D. Section E has an island between 28 and 35 m.

power in sections E and F is not as high as in se and D and there is not such an obvious thalw shows the areal changes in cross-sectional area i]

direction, from left (section A) to right (section survey periods are indicated using different sym Maximum erosion of up to 1 m2 and deposition are shown. For example, the first survey perioc August) indicated by the solid line, showed alr erosion at section A, 0.6 m2 of deposition at sec of erosion at section C, and between 0.4 and 0.6 n at sections D, E, and F. In the second survey perio the pattern is more or less reversed, whereas the t

periods are dominated by erosion, though this is at some sections in the first and second survey pe survey period showed erosion at sections A, B, <

no net change at sections D and E and some e section F. In the fourth period there is almost 0..' tion at section A and then a consistent downstre erosion to 0.3-0.4 m2 at sections D, E, and F.

Figure 6B shows the computed transport rq

using the sediment budgeting approach, showing est in the first period (29 July to 1 August) at with a tail off towards the downstream end of ti

July-1 gust

kugust-6 gust

kugust-1 4

gust

August-

tions E and F The next highest transport rates were seen in the second survey period (1-6 August) and these increased towards the downstream end of the reach. The third and fourth periods showed similar transport rates at sections A, B, C, and D, but the fourth period had higher transport rates at sections E and F.

DIRECT MEASUREMENT OF BEDLOAD TRANSPORT

August Figure 7A shows bedload transport rates measured at half-

hourly intervals during the intensive sampling program from 1 to 3 August. Even the maximum transport rates measured (0.04 kg m-1 s-1) are very low and rates in excess of 0.02 kg m-1 s-' are only really measured when specific stream power exceeds

-.-29 July-1 2.0 kg m-1 s-1. The considerably higher discharge on the after- August noon of 1 August as compared to the 2 August seemed to have

had little effect on the transport rate. Indeed, if anything, trans-

port rates were generally higher on 2 August. The relationship

-- 1 August- between bedload transport rate and specific stream power is plot- 6 August ted in Figure 7B. It should be noted that the measurements plot-

ted in Figure 7B were taken during a warm sunny spell when

discharge was being controlled by glacial melt only. As can be seen from Figures 3 and 4 the rainfall in the third and fourth

A --6 August- survey periods generated considerably higher discharges which 14 August were not sampled by this method. The rating relationship in Fig-

ure 7B is both site specific, and did not sample the higher dis-

charges measured later in the study period. It is therefore not ion and deposi- used to determine reach scale bedload transport rates, but rather is sections over used to show the temporal variation in bedload transport at one v F from left to . F from left to

point in the river. B: Downstream ates determined r time periods. Discussion ht (upstream to

Maximum bedload transport rates determined from cross- section resurveys occurred in the first period (29 July to 1 Au-

gust) at section B where rates reached 0.8 m2 d-l (and ranged could be larger between 0.5 and 0.8 m2 d-l1). Assuming a sediment bulk density

Duld be applied of 1800 t m-3, this represents transport rates of 0.01 kg m-1 s-' and 0.02 kg m-1 s-1, respectively) for four (sections A, B, C,

g from the right and D) of the six sections during the first period. However, bed- mnel in section load transport rates measured "at-a-point" by the Helley-Smith Specific stream bedload sampler deployed through two meltwater floods on I :ctions A, B, C, and 2 August peaked at around 0.04 kg m-I s-l. Nicholas and

veg. Figure 6A Sambrook Smith (1998) reported maximum transport rates of n a down-reach 0.68 m2 d-l for their first survey period and 0.26 m2 d- I for their F) and the four second period on the Virkisa River in southeast Iceland. In con- ibols and lines. trast to this, Ferguson and Ashworth (1992) report values, also of up to 0.8 m2 inferred from patterns of erosion and deposition, of 2.4 and 1.5 J (29 July to 1 m2 d-~ for the proglacial White River and Sunwapta River, re- most 0.6 m2 of spectively. The recognition by Nicholas and Sambrook Smith ction B, 0.4 m2 (1998) that the resurvey technique is likely to underestimate bed- n2 of deposition load transport rates, especially relative to instantaneous mea- Ad (1-6 August) surements at times of peak flow, is bom out by the findings of third and fourth this study where the Helley-Smith sampling measured peak bed- not as great as load transport rates of rates 0.04 kg m-1 s-1, double those esti-

riods. The third mated from the resurvey method. Nevertheless, both methods and C, virtually indicate transport rates that are an order of magnitude lower than rosion again at the comparable data of Ferguson and Ashworth (1992), but 5 m2 of deposi- broadly comparable with the transport rates reported by Nicholas -am increase in and Sambrook Smith (1998) using the resurvey technique on the

Virkisa River in southeast Iceland. ates determined The temporal and spatial trends in channel erosion/deposi- this to be high- tion and in bedload transport rates illustrated in Figures 6A and sections A-D, 6B suggest that there may be downstream migration of a wave

ie reach in sec- or lobe of sediment over the study period (Hoey, 1992; Nicholas




I Bedload Transport Rate - Discharge

25.0

'' 20.0

&4 15.0 L.

I 10.0-

5.0 -

N

I I - .

i II I I, 111.1 0.0 I III I ll 11III .1 . . 111111 I

12:00:00 18:00:00 00:00:00 06:00:00 12:00:( 01/08/1998 01/08/1998 02/08/1998 02/08/1998 02/08/1

0.10000

I II, I II I I, . . ... ---. ,11 00 18:00:00 00:00:00 06:00:00 998 02/08/1998 03/08/1998 03/08/1998

y = 0.0087Ln(x)- 0.0126

R2 = 0.2081 % *

*- -

.** * .*

0.00010

** 4*0 * 4

10.0

Log10 Discharge (m3 s-1)

FIGURE 7. A: Skeldal River bedload transport rate measured "at-a-point" (see Fig. IC) using a Helley-Smith sampler through two meltwater flood events on 1-3 August 1998. B: Log-log plot of bedload transport rate and discharge for half hourly bedload samples collected 1-3 August using a modified Helley-Smith sampler.

et al., 1995). During the period 29 July to 1 August there was deposition of 0.6 m2 in section B whereas section C showed 0.4 m2 of erosion. In the second period (1-6 August) a large amount of sediment was removed from section B resulting in 1 m2 of erosion, mainly concentrated in the area of highest stream power (between 5 and 9 m along the section) and this material appeared to accumulate in section C, some 20 m downstream, where there

TABLE 1

Channel cross-section changes and flow thresholds for four survey intervals, Skeldal River, northeast Greenland, 1998

Net change in Time in hours when

cross-section area (n2) (mean for Q > thresholds

(m2) (mean for Length of period (m3 s-l) whole reach

Survey dates -6 sections) Hours Days 10 15 20

29 Jul.-l Aug. 0.24 48 2 22 8 0 1 Aug.-6 Aug. -0.35 120 5 37 11 0 6 Aug.-14 Aug. -0.10 192 8 47 19 2 14 Aug.-21 Aug. -0.19 168 7 61 25 10

was a change in cross-section area from 0.4 m2 of erosion to 0.3 m2 of deposition. During the third time interval (6-14 August) there was aggradation in sections D and E of 1 m2 and 0.8 m2, respectively, suggesting that the sediment lobe had moved into the downstream half of the reach and attenuated (Fig. 2). The lobe seems to have left the reach during the fourth period (14- 21 August) as shown by decreases in area (erosion) of section D, E, and F and a decrease in channel depth. This suggests that the lobe of sediment moved at least 100 m in order to have left the reach over the study period. Nicholas and Sambrook Smith (1998) reported such a feature which moved through their 250 m reach on the Virkisa River in southern Iceland at a similar rate. In the light of this they suggested that upstream sediment supply is probably the dominant control on bedload movement through the stable channel network.

The rainfall and discharge record in this study allowed some inferences to be made regarding the driving forces for the channel change observed. Figure 3 shows the rainfall distribu- tion from 20 July to 26 August 1998. Highest rainfall totals occurred during the study period on 14-15 August. The total rainfall for the period was approximately 71 mm. The annual mean is around 300 mm, of which 200-250 mm is in the form

T. STOTT / 341

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0.00100

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UE 0)

Q)

0 0Q 0

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0

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B

0.00001

1.0 100.0

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TABLE 2

Bedload transport rates from selected arctic and alpine studies

Bedload transport rate Channel slope

Data source Location (m m-') Bed grain size (m2 d- )a (kg m-1 s-1)

Alpine studies

Ferguson et al. (1989) White River, Washington, U.S.A. Dso 40-81 mm 4

Goff and Ashmore (1994) Sunwapta River, Canada 0.1-2.85a

Warburton (1992) Bas Glacier d'Arolla, Switzerland 0.002-0.01 Bed surface Dso 18-35 mm Up to 1.75

Sub-surface D50 4-10 mm

Ferguson and Ashworth (1992) White River, Washington, U.S.A. D50 40-81 mm 2.4 0.05a

Lane et al. (1995) Arolla, Switzerland 0.002-0.01 Bed surface D50 18-35 mm 0.04-0.74

Sub-surface D50 4-10 mm

Ferguson and Ashworth (1992) Sunwapta, Canada 1.5 0.03a

Mean

Arctic studies

Ashworth and Ferguson (1986) Lyngsdalselva, Arctic Norway 0.023 Median surface 69 mm 0.0002 during melt up to 3.5 during rain

Mosley (1988) Onyx River, Antarctica 0.0025-0.0074 Bulk samples: 0.38-2.16

Sandur median-2.7; channel bed = 4.6 mm

Stott, this study Skeldal River, NE Greenland 0.008 Bar surfaces 43 mm Up to 0.8 0.02a-0.04b

Nicholas and Sambrook Smith (1998) Virkisa, southern Iceland 0.014-0.018 Median within channel 34-80 mm 0.26-0.68 0.005-0.014a

Busskamp and Hasholt (1996) SE Greenland 0.005-0.02 Dso surface 10-40 mm 0.002

Stott and Short (1996) Gipsdalen, Svalbard 0.004 D50 surface 2 mm 0.0005

Mean

a Estimated from sediment budgeting approach. b Estimated from "at-a-point" Helley-Smith sampling.

C)

z--

Fl

C) Mean

4 1.47 0.88

0.05

0.39

0.03

1.14

1.65

1.27

0.03 0.01

0.002 0.0005

0.49



of snow. It seems, therefore, that the rainfall total for this summer was quite exceptional. The effect of the rainfall in August can be seen on the discharge record in Figure 4 when the highest discharge for the study period peaked at an estimated 46 m3 s-~ (though this was not accurately gauged) on 11 August. Further rainfall caused considerably higher than normal discharges on 12-13 August and again on 16-17 August.

The Helley-Smith bedload transport rates (measured half hourly from 1-3 August) plotted in Figure 7A would suggest that it is not unreasonable to choose a threshold for bedload transport of around 10 m3 s-~. Taking this as the "critical discharge" threshold for the movement of bedload (Werritty, 1984) and thus changes in channel dimensions, along with two higher thresholds, Table 1 presents the differences between the periods. It is assumed that even though the periods are of unequal length, that the time (in hours) when flow exceeded this critical discharge would be the important parameter influencing bedload transport and channel change. Only during the first period (29 July to 1 August) was there net deposition in the reach. This corresponded to a period of fine weather with no rainfall. A clear diurnal discharge pattern was observed, and although not gauged, the pattern was very similar to that for the 1-3 August as seen on Figure 4. The assumed critical discharge for bedload transport of 10 m3 s-1 was exceeded for 22-h, and the 15 m3 s-l threshold was exceeded for 8 h. The second, third, and fourth periods all had successively greater num- bers of hours when the discharge thresholds were exceeded and all showed net erosion. However, the magnitude of the net erosion does not reflect the duration of discharge exceeding the thresholds. Indeed, the second period shows greatest net erosion, yet has least time when discharge thresholds were exceeded. The fourth period has the second greatest net erosion but has the greatest duration of discharge exceeding the thresholds. The third period has the least net erosion and second greatest duration of discharge exceeding the thresholds. What is not clear in a study with such coarse temporal resolution, however, is whether or not changes of greater magnitude may have been cancelled out, particularly during the two periods where rainfall induced higher discharges.

It would seem therefore, that neither discharge magnitude nor the time over which discharge exceeds a critical threshold value can be used to successfully explain the erosion and deposition patterns observed in this study. It may be that upstream bedload supply parameters, possibly related to the stage in the melt season (Clifford et al., 1995) or bed imbrication, for example, may prove more fruitful for future investigations when attempting to explain the patterns observed.

Table 2 reports bedload transport rates from 12 studies, six from alpine settings and six from polar regions (including this study) for comparison. It should be noted that all studies reported took place during the summer melt season and so bedload transport rates will be considerably higher than annual means which may be reported elsewhere. Direct comparisons can only be undertaken with extreme caution since different methods have been used to derive these estimates. Information on channel slopes and grain size is given and the different methods used to derive transport rates are indicated. Also it should be borne in mind that Warburton (1990) notes, "bedload moves in 'threads' with transport rates varying by a factor of 10 in the space of a metre" and that "this bedload 'streaming' effect varies markedly between days." Pitlick (1988) also found considerable variation in bedload transport rates: cross-section mean values were reported to be of the same order of magnitude as the standard deviation of the measurements. Mosley (1988) reported a large degree of variability in bedload transport estimates in the Onyx River, Ant- arctica, derived from Helley-Smith samples, even when dis-

charge remained constant during the measurement period. Nev- ertheless, given this inherent large variability in bedload transport rates, it is still possible to see from the 12 studies reported (six from arctic and six from alpine rivers), that the mean bedload transport rate for the alpine studies is more than double that for the arctic studies. Table 2 does not include all recent studies and does not attempt to record all the factors affecting bedload transport rates, but nevertheless it does indicate some differences in the broad estimates of bedload transport rates made in arctic and alpine rivers in recent years.

Conclusions

(1) This study adds to the very limited data set of bedload transport rates for arctic environments. Bedload transport rates reported here fall within the extremes reported for arctic and subarctic environments, but are still significantly lower that those reported from alpine rivers.

(2) Maximum bedload transport rates derived using a sediment budgeting approach appear to be 50% lower than maximum rates measured "at-a-point" using a Helley-Smith bedload sampler during meltwater flood events. Though the different lengths of the sampling period for the two methods does not allow a true comparison of the two techniques to be made, it is suggested that the budgeting approach offered a more reliable technique than the Helley-Smith sampler.

(3) Bedload transport rates derived in this study vary both spatially and temporally by up to an order of magnitude. Bedload sampling conducted through flood events suggests that there is a broad (but not significant) relationship with discharge, but upstream bedload supply parameters are concluded to be more influential on the transport patterns and bed elevation changes observed.

Acknowledgments This work was carried out as part of the science program

on BSES Expeditions' northeast Greenland summer 1998 expedition. The expedition organisation by BSES Expeditions un- der the Chief Leadership of Pat Cannings is acknowledged in this work. Field assistance and load carrying by young expedi- tioners made this work in a remote location possible. Dr Greg Sambrook Smith and Dr Jeff Warburton reviewed an earlier draft of the manuscript and Dr John Pitlick and other anonymous ref- erees made useful comments at the reviewing stage, all of which greatly improved the manuscript.

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Manuscript submitted August 2000 Revised manuscript submitted May 2001




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Bedload Transport and Channel Bed Changes in the Proglacial Skeldal River, Northeast Greenland

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