Charles F. Raymond

Albedo of dirty snow during conditions of melt

The evolution of spectrally averaged albedo (wavelengths between 0.28 μm and 2.8 μm) of snow surfaces treated with known initial concentrations of particles of submicron-sized soot and air fall volcanic ash was investigated during conditions of natural melt. Depending on the particle type and concentration, the initial applications reduced the surface albedo to values ranging from 0.18 to 0.41 which were substantially lower than the albedo of the untreated natural snow (about 0.61). Many of the soot particles flushed through the snowpack with the meltwater, and surface concentrations of soot greater than about 5 × 10−7 kg/kg did not persist for more than a few days. The migration of particles to depth caused the snow to brighten after the initial application, thus limiting the amount of albedo reduction and the consequent effects on melting. Nevertheless, the soot remaining near the surface had a substantial, long-term effect. The residual concentration of 5 × 10−7 kg/kg persisted for several weeks and, compared to the untreated surface, reduced the albedo by about 30% and increased melting by 50%. Particles of volcanic ash with diameters larger than about 5 μm remained at or near the snow surface. Although many of the smaller particles flushed through the snow with the meltwater, the surface albedo was not changed significantly by their removal. The different behaviors of the ash and soot are probably related to the difference in their particle size distributions in relation to the thickness of water films that form the transport paths under conditions of partial saturation that are characteristic of melting snow.

Hydrology, erosion and sediment production in a surging glacier: Variegated glacier, Alaska, 1982-83

Outlet streams of Variegated Glacier, Alaska, U .S.A ., were observed before, during and after the surge of 1982-83. Measurements of discharge, suspended sediment and dissolved load in the outlet streams are presented for the years 198284, and comparisons are made with data from previous years. The data are interpreted to yield characteristics of the basal hydraulic system. The surging region of the glacier was underlain by a basal hydraulic zone of low water velocity and high water storage, inferred to be a distributed-flow system. The ice down-glacier of the propagating surge front was underlain by a highvelocity, low-storage zone, inferred to be a conduit system. The volume of water stored above the surge front was the major hydraulic control on the surge. Basal bedrock erosion during the surge was extreme in comparison to non-surging glaciers. The sediment output was directly proportional to the basal sliding, with a dimensionless erosion rate (meters eroded from the bed divided by meters of sliding) of order 1.0 x 10--4. Total erosion during the 20 year surge cycle was on the order of 0.3 m of bedrock, with approximately two-thirds occurring during the 2 years of the surge peak, and the bulk of this during the peak 2 months .

Shear margins in glaciers and ice sheets

Analytical and numerical techniques are used to examine the flow response of a sloped slab of power-law fluid (power n) subjected to basal boundary conditions that vary spatially across the flow direction, as for example near an ice-stream margin with planar basal topography. The primary assumption is that basal shear stress is proportional to the basal speed time a spatially variable slip resistance. The ratio of mean basal speed to the speed originating from shearing through the thickness, denoted as r, gives a measure of how slippery the bed is. The principal conclusion is that a localized disturbance in slip resistance affects the basal stress and speed in a zone spread over a greater width of the flow. In units of ice thickness H. the spatial scale of spreading is proportional to a single dimensionless number R n ≡(r/n+1) 1/n+1 derived from n and r. The consequence for a shear zone above a sharp jump in slip resistance is that the shearing is spread out over a boundary layer with a width proportional to R n . For an ice stream caused by a band of low slip resistance with a half-width of wH. the margins influence velocity and stress in the central part of the band depending on R n in comparison to w. Three regimes can be identified, which for n=3 are quantified as follows: low r defined as R 3 1w, for which the boundary layers from both sides bridge across the full flow width and the driving stress in the center is supported almost entirely by side drag; intermediate r, for which the driving stress in the center is supported by a combination of basal and side drag. Shear zones that are narrower than predicted on the basis of this theory (≃R 3 ) would require localized softening of the ice to explain the concentration of deformation at a shorter scale.

Switch of flow direction in an Antarctic ice stream

Fast-flowing ice streams transport ice from the interior of West Antarctica to the ocean, and fluctuations in their activity control the mass balance of the ice sheet. The mass balance of the Ross Sea sector of the West Antarctic ice sheet is now positive—that is, it is growing—mainly because one of the ice streams (ice stream C) slowed down about 150 years ago. Here we present evidence from both surface measurements and remote sensing that demonstrates the highly dynamic nature of the Ross drainage system. We show that the flow in an area that once discharged into ice stream C has changed direction, now draining into the Whillans ice stream (formerly ice stream B). This switch in flow direction is a result of continuing thinning of the Whillans ice stream and recent thickening of ice stream C. Further abrupt reorganization of the activity and configuration of the ice streams over short timescales is to be expected in the future as the surface topography of the ice sheet responds to the combined effects of internal dynamics and long-term climate change. We suggest that caution is needed when using observations of short-term mass changes to draw conclusions about the large-scale mass balance of the ice sheet.

Snow stability during rain

The mechanical response of snow packs to penetrating liquid water was observed over two winter seasons in the central Cascade Mountains, Washington, U.S.A. Following the onset of rain, three evolutionary regimes of snow behavior were identified: immediate avalanching, delayed avalanching, and return to stability. Immediate avalanching occurred within minutes to an hour after the onset of rain and the time of release could be predicted with an accuracy of less than an hour from meteorological forecasts of the transition from snow to rain. These avalanches usually slid on surfaces substantially deeper than the level to which water or associated thermal effects had penetrated. The mechanism by which alteration of a thin skin of surface snow can cause deep slab failure has not been identified, but several possibilities involving a redistribution of stress are discussed. Delayed avalanches released several hours after rain started. The delay varied, depending on the rate of increasing stress associated with the additional precipitation, and on the time taken for water to penetrate and weaken a potential sliding layer. It is difficult to define accurately the evolving distribution ofliquid water in snow which makes it difficult to predict accurately the time of avalanching. Depth profiles of the rate of snow settlement showed that a wave of increased strain rate propagated into the snow in response to penetrating water. This type of measurement could prove useful for predicting when snow stability is reaching a critical condition. Avalanche activity was rare after continuation of rain for 15 h or more. This return to stability occurred after drainage structures had evolved and penetrated the full depth of the snowpack. Established drain channels route water away from potential sliding surfaces and are also relatively strong structures within a snowpack.

Interfacial water in polar glaciers and glacier sliding at −17°C

We have observed sliding at a cold (-17 oC) ice-rock interface beneath Meserve Glacier, Antarctica, and the segregation of ice into clean lenses amidst the dirty basal layers of this glacier. We interpret these as manifestations of thin water films at ice-rock interfaces. We use Shreves theory for sub-freezing sliding to esti- mate the nominal film thickness to be at least tens of nanometers. Such water films should exist around rocks in most polar ices, and likely have high solute concen- trations due to solute rejection during regelation and due to exchange with veins and grain boundaries where impurities reside.

Energy balance of ice streams

Analysis of the cross-flow transmission of force from the central parts of a well-lubricated ice stream to its margins shows that there is a corresponding shift in the lateral location of motion-induced heat generation. The rate of basal heat generation in the center can be substantially smaller than the local rate of potential energy loss given by driving stress times the speed of downslope motion. The basal heating is a maximum for an intermediate level of lubrication for which speed is about 40% of the speed over a frictionless bed and base stress is about 25% of the driving stress. Stable and unstable balances between meltwater production and drainage on the bed are identified. A stable steady state with a speed less (more) than that giving maximum heat generation is termed drainage-(production-) limited, since an increase in speed would lead to increased (decreased) basal melting and must (need not) be balanced by increased drainage. It is shown that gradual evolution of the basal water drainage system and the factors affecting basal melting can cause discontinuous jumps between fast- and slow-moving states. A simplified analysis applied to six cross-sections of West Antarctic Ice Streams B, D, E and Rutford Ice Stream shows them to be diverse in the level of support from the sides and corresponding shift of mechanical heating sideward from their central parts. The cross-sections of Ice Stream B near Upstream B may be production-limited, because of especially high lubrication and related support from the sides. Cross-sections in the upper part of Ice Stream D, Ice Stream E and Rutford Ice Stream are in a drainage-limited condition. Substantial reduction of basal heat generation by side drag (in most cases) and expected high heat flow into the basal ice associated with low thickness (in some cases) tends to favor basal freezing. Nevertheless, all of the examined cross-sections except one are expected to experience basal melting with a modest geothermal heat-flux density of 60 mW m or less in some cases. The lower part of Ice Stream B is an exception, where the analysis indicates that geothermal flux density must exceed 80-100 mW m to maintain melting. If this high geothermal flux is not present, then the base of the lower part of Ice Stream B may be freezing, which would suggest continued deceleration of this part of Ice Stream B.

The accumulation pattern across Siple Dome, West Antarctica, inferred from radar-detected internal layers

The spatial distribution of accumulation across Siple Dome,West Antarc-tica, is determined from analysis of the shapes of internal layers detected by radio-echo sounding (RES) measurements. A range of assumed accumulation patterns is used in an ice-flow model to calculate a set of internal layer patterns. Inverse techniques are used to determine which assumed accumulation pattern produces a calculated internal layer pattern that best matches the shape of internal layers from RES measurements. All of the observed internal layer shapes at Siple Dome can be matched using a spatially asymmetric accumulation pattern which has been steady over time. Relative to the divide, the best-fitting accumulation pattern predicts 40% less accumulation 30 km from the divide on the south flank of Siple Dome and 15^40% more accumulation 30 km from the divide on the north flank. The data also allow the possibility for a small time variation of the pattern north of the divide. The mismatch between the calculated and the observed layer shapes is slightly reduced when the accumulation rate north of the divide is higher in the past (>5 kyr BP) than at present. Sensitivity tests show that the predicted change in the spatial accumulation pattern required to cause the slight Siple Dome divide migration (inferred from previous studies) would be detectable in the internal layer pattern if it persisted for 42 kyr. Our analysis reveals no evidence that such a change has occurred, and the possible change in accumulation distribution allowed by the data is in the opposite sense. Therefore, it is unlikely that the Siple Dome divide migration has been caused by a temporal change in the spatial pattern of accumulation. This conclusion suggests the migration may be caused by elevation changes in Ice Streams C and D at the boundaries of Siple Dome.

Thermal effects on the location of ice stream margins

West Antarctic ice streams move rapidly over beds that are thawed by heat generated by the motion. Intervening zones of slow-moving ice are probably freezing at their bases, where heat generation is absent. Heat is also generated in the shear zone between the slow and fast moving ice. Through modeling of coupled heat and mass flow, we show that the position of the boundary between freezing and melting at the bed is unstable for an otherwise morphologically uniform bed. Above a threshold speed in the ice stream, heat generation in the shear zone allows melting beneath the slow ice and possible widening of the stream. Below this threshold, heat generation at the edge of the thawed zone is insufficient to maintain melting, leading to freezing and narrowing of the stream. The threshold speed appears to be a function of the temperature at the base of slow ice (relative to the melting point) and is also highly sensitive to the inflow of colder ice from the slow zone. For environmental conditions typical of West Antarctica, the threshold speed is about 100 m yr−1 in the absence of inflow and increases to typical ice stream speeds (500 m yr−1) for inflow speeds of a few meters per year. Thus the ice streams can be in a delicate thermal balance with the potential of melting outward or freezing inward. Rapidly moving ice streams could expand unless blocked by bed properties that would prevent fast motion even when there is melting at the bed.

Flow in a transverse section of Athabasca Glacier, Alberta, Canada

Measurements of ice deformation at the surface and at depth in the Athabasca Glacier, Canada, reveal for the first time the pattern of flow in a nearly complete cross section of a valley glacier, and make it possible to test the applicability of experimental and theoretical concepts in the analysis of glacier flow. Tilting in nine boreholes (depth about 300 m, eight holes essentially to the bottom) was measured with a newly developed electrical inclinometer, which allows a great increase in the speed and accuracy with which borehole configurations can be determined, in comparison with earlier methods. The measurements define the distribution of the velocity vector and the strain-rate tensor over 70% of the area of the glacier cross section. The main longitudinal component of flow has the following general features: (1) basal sliding velocity which exceeds 70% of the surface velocity over half of the width of the glacier, (2) marginal sliding velocity (not more than a few meters per year) much less than basal sliding velocity at the centerline (about 40 m yr(-1)), (3) marginal shear strain rate near the valley walls two to three times larger than the basal shear strain rate near the centerline (0.1 yr(-1)). The observed longitudinal flow is significantly different from that expected from theoretical analysis of flow in cylindrical channels (Nye, 1965). The relative strength of marginal and basal shear strain rate is opposite to that expected from theory. In addition, the longitudinal flow velocity averaged over the glacier cross section (which determines the flux of ice transported) is larger by 11% than the average flow velocity seen at the glacier surface, whereas it would be 2% smaller if the theoretical prediction were correct. These differences are caused to a large extent by the constant sliding velocity assumed in the theoretical analysis, which contrasts strongly with the actual distribution of sliding. The observed relation between marginal and basal sliding velocity is probably a general flow feature in valley glaciers, and may be caused by lateral variation of water pressure at the ice-rock contact. The observed pattern of longitudinal velocity over the section also shows in detail certain additional features incompatible with the theoretical treatment, even after the difference in boundary conditions (distribution of sliding velocity) is taken into account. Longitudinal strain rate (a compression of about 0.02 yr(-1) at the surface) decreases with depth, becoming nearly 0 at the bed in the center of the glacier. The depth variation cannot be explained completely by overall bending of the ice mass as a result of a longitudinal gradient in the curvature of the bed, and is at variance with existing theories, which require the longitudinal strain rate to be constant with depth. Motion transverse to the longitudinal flow occurs in a roughly symmetric pattern of diverging marginward flow, with most of the lateral transport occurring at depth in a fashion reminiscent of extrusion flow. The observed lateral velocities averaged over depth (up to 1.9 m yr(-1)) are compatible with the lateral flux required to maintain equilibrium of the marginal portions of the glacier surface under ablation (3.7 m yr(-1)) and are driven by the convex transverse profile of the ice surface. When the measured strain-rate field is analyzed on the basis of the standard assumption that the shear stress parallel to the glacier surface varies linearly with depth, the rheological behavior in the lower one-half to two-thirds of the glacier is found compatible with a power-type flow law with n = 5.3. However, the upper one-third to one-half of the glacier constitutes an anomalous zone in which this treatment gives physically unreasonable rheological behavior. In a new method of analysis, rheological parameters are chosen so as to minimize the fictitious body forces that appear as residuals in the equilibirum equations when evaluated for the measured strain-rate field. This new method requires no a priori assumptions about the stress distribution, although for simplicity in application, the mean stress is assumed constant longitudinally. This treatment shows that the anomalies in the near-surface zone are due to significant departures from linear dependence of shear stress on depth, and gives a flow-law exponent of n = 3.6, which is closer than n = 5.3 to values determined by laboratory experiments on ice.