Kilometer-scale variations in ice velocity and surface topography are used to investigate the style of glacier deformation in the main body of Ice Stream B, West Antarctica. The pattern is very different from that reported for other glaciers. For the 250 km2 area studied on Ice Stream B, most of the observed deformation occurs within two narrow bands, in which there is large across-flow compression and slow lateral shearing. The bands underlie valleys in the ice-surface topography. Measured upward displacement of ice adjacent to the rapidly compressing bands appears to be linked to the creation of the ice stream’s topography. The most likely cause for the observed pattern of strain rates and surface topography, and their changes over time, is deformation guided by longitudinal bands of ice with an aligned crystal fabric.

The aim of deep ice drilling on Law Dome, Antarctica, has been to exploit the special characteristics of Law Dome summit, i.e. low temperature and high accumulation near an ice divide, to obtain a high-resolution ice core for climatic/environmental studies of the Holocene and the Last Glacial Maximum (LGM). Drilling was completed in February 1993, when basal ice containing small fragments of rock was reached at a depth of 1196 m. Accurate ice dating, obtained by counting annual layers revealed by fine-detail δ18О, peroxide and electrical-conductivity measurements, is continuous down to 399 m, corresponding to a date of AD 1304. Sulphate concentration measurements, made around depths where conductivity tracing indicates volcanic fallout, allow confirmation of the dating (for Agung in 1963 and Tambora in 1815) or estimates of the eruption date from the ice dating (for the Kuwae, Vanuatu, eruption ~1457). The lower part of the core is dated by extrapolating the layer-counting using a simple model of the ice flow. At the LGM, ice-fabric measurements show a large decrease (250 to 14 mm2) in crystal size and a narrow maximum in c-axis vertically. The main zone of strong single-pole fabrics however, is located higher up in a broad zone around 900 m. Oxygen-isotope (δ18O) measurements show Holocene ice down to 1113 m, the LGM at 1133 m and warm (δ18O) about the same as Holocene) ice near the base of the ice sheet. The LGM/Holocene δ18O shift of 7.0‰, only ~1‰ larger than for Vostok, indicates that Law Dome remained an independent ice cap and was not overridden by the inland ice sheet in the Glacial.

Pit-wall samples were collected from two sites about 2 km apart on Agassiz Ice Cap, Ellesmere Island, Northwest Territories, Canada, in 1992, 1993 and 1994, and from a site a further 1 km distant, in order to study spatial and seasonal variations in snow chemistry. Two of the pits were dug in wind-scoured zones and one in an unscoured zone. Although a large part of the winter snow is removed from the scoured zones (which do not show very negative δ18O values) the winter/spring anion peaks are still evident; this may be due to the predominance of dry deposition in mid-winter. The Cl− and SO42– ions peak in late winter/early spring, while NO3− peaks both in late winter/early spring and in summer. Vertical concentration profiles of all anions did not significantly alter over a 2 year period, indicating that there are no serious post-depositional changes due to evaporation, snow melting or photochemical reactions. However, comparisons between stake/board snow-accumulation measurements and those derived from the least scoured pit indicate that a single pit will represent annual accumulation rates for a local area only.

During ice-structure interaction, ice will fail in a brittle manner dominated by two processes. The first corresponds to the formation of macrocracks and the consequent spalling-off of large ice pieces. The second includes an intense shear-damage process in zones, termed critical zones, where high pressures are transmitted to the structure. The shear-damage process results in microstructural changes including microcrack formation and recrystallization. A range of tests on laboratory-prepared granular ice have been conducted to determine the fundamental behaviour of ice under various stress states and stress history, particularly as it relates to changes in microstructure. The test series was designed to study three aspects: the intrinsic creep properties of intact, undamaged ice; the enhancement of creep and changes in microstructure due to damage; and the effects of different stress paths. Tests on intact ice with triaxial confining pressures and low deviatoric stresses, aimed at defining the intrinsic creep response in the absence of microcracking, showed that an accelerated creep rate occurred at relatively low deviatoric stresses. Hence, a minimum Creep rate occurred under these conditions. Recrystallization to a smaller grain-size and void formation were observed. Ice damaged uniaxially and triaxially prior to testing showed enhancement of creep under both uniaxial and triaxial loading conditions Creep rates in triaxially damaged ice were found to be non-linear with high deviatoric stresses, corresponding to a power-law dependence of creep rate. Uniaxially damaged specimens contained microcracks parallel to the stressed direction which tended to close under triaxial confinement. Damage under triaxial conditions at low confining pressures produced small recrystallized grains near zones of microcracking. At high confining pressures, a fine-grained recrystallized structure with no apparent cracking was observed uniformly across the specimen. The recrystallization process contributes significantly to the enhanced creep rates found in damaged specimens.

Pressure and tracer measurements in boreholes drilled to the bottom of Ice Stream B, West Antarctica, are used to obtain information about the basal water conduit system in which high water pressures are developed.These high pressures presumably make possible the rapid movement of the ice stream. Pressure in the system is indicated by the borehole water level once connection to the conduit system is made. On initial connection, here also called “breakthrough” to the basal water system, the water level drops in a few minutes to an initial depth in the range 96–117 m below the surface. These water levels are near but mostly somewhat deeper than the floation level of about 100 m depth (water level at which basal water pressure and ice overburden pressure are equal), which is calculated from depth-density profiles and is measured in one borehole. The conduit system can be modelled as a continuous or somewhat discontinuous gap between ice and bed; the thickness of the gap δ has to be about 2 mm to account for the water-level drop on breakthrough, and about 4 mm to fit the results of a salt-tracer experiment indicating downstream transport at a speed of 7.5 mm s−1. The above gap-conduit model is, however, ruled out by the way a pressure pulse injected into the basal water system at breakthrough propagates outward from the injection hole, and also by the large hole-to-hole variation in measured basal pressure, which if present in a gap-conduit system with δ = 2 or 4 mm would result in unacceptably large local water fluxes. An alternative model that avoids these objections, called the “gap opening” model, involves opening a gap as injection proceeds: starting with a thin film, the injection of water under pressure lifts the ice mass around the borehole, creating a gap 3 or 4mm wide at the ice/bed interface. Evaluated quantitatively, the gap-opening model accounts for the volume of water that the basal water system accepts on breakthrough, which obviates the gap-conduit model. In order to transport basal meltwater from upstream it is then necessary for the complete hydraulic model to contain also a network of relatively large conduits, of which the most promising type is the “canal” conduit proposed theoretically by Walder and Fowler (1994): flat, low conduits incised into the till, ∼0.1 m deep and perhaps ∼1 m wide, with a flat ice roof. The basal water-pressure data suggest that the canals are spaced ∼50–300 m apart, much closer than R-tunnels would be. The deepest observed water level, 117 m, is the most likely to reflect the actual water pressure in the canals, corresponding to a basal effective pressure of 1.6 bar. In this interpretation, the shallower water levels are affected by loss of hydraulic head in the narrow passageway (s) that connect along the bed from borehole to canal(s). Once a borehole has frozen up and any passageways connecting with canals have become closed, a pressure sensor in contact with the unfrozen till that underlies the ice will measure the pore pressure in the till, given enough time for pressure equilibration. This pressure varies considerably with time, over the equivalent water-level range from 100 to 113 m. Basal pressure sensors 500 m apart report uncorrelated variations, whereas sensors in boreholes 25 m араrt report mostly (but not entirely) well-correlated variations, of unknown origin. In part of the record, remarkable anticorrelated variations are interspersed with positively correlated ones, and there are rare, abrupt excursions to extreme water levels as deep as 125 m and as shallow as 74 m. A diurnal pressure fluctuation, intermittently observed, may possibly be caused by the ocean tide in the Ross Sea. The lack of any observed variation in ice-stream motion, when large percentagewise variations in basal effective pressure were occurring according to our data, suggests that the observed pressure variations are sufficiently local, and so randomly variable from place to place, that they are averaged out in the process by which the basal motion of the ice stream is determined by an integration over a large area of the bed.

Dry snow on the surface of polar ice ice sheets is first densified and metamorphosed to produce firn. Bubbly ice is the next stage of the transformation process which takes place below the depth of pore closure. This stage extends to the transition zone where, due to high pressures and low temperatures. air trapped in bubbles and ice begins to form the mixed air clathrate hydrates, while the gas phase progressively disappears. Here we develop a model of bubbly-ice rheology and ice-sheet dynamics taking into account glacier-ice compressibility. The interaction between hydrostatic compression of air bubbles, deviatoric (uniaxial) compressive deformation of the ice matrix and global deformations of the glacier body is considered. The ice-matrix pressure and the absolute-load pressure are distinguished. Similarity theory and scale analysis are used in examine the resultant mathematical model of bubbly-ice densification. The initial rate of bubble compression in ice sheets appears to be relatively high, so that the pressure (density) relaxation process takes place only 150-200 m in depth (below pore close-off) to reach its asymptotic phase, wherein the minimal drop between bubble and ice pressures is governed by the rate of loading (ice accumulation). This makes it possible to consider densification under stationary (present-day) conditions of ice formation as a special case of primary interest. The computational tests performed with the model indicate that both ice-porosity and bubble-pressure profiles in ice sheets are sensitive to variations of the rheological parameters of pure ice. However, only the bubble-pressure distinguishes between the rheological properties at low and high stresses. The porosity profile at the asymptotic phase is mostly determined by the air content in the ice. In the companion paper (Lipenkov and others, 1997), we apply the model to experimental data from polar ice cores and deduce, through an inverse procedure, the rhelogical properties of pure ice as well as the mean air content in Holocene and glacial ice sediments at vostok Station (Antarctica).

The partitioning of resistive force between the bed and sides of Ice Stream B, Antarctica, is obtained for three large areas that have been measured using repeat aerial photogrammetry. Problems associated with data errors and local variations in ice strength and velocity are reduced by considering the areally averaged budget of forces for each photo block. Results indicate that the bed under Ice Stream B must be very weak and unable to provide much resistance. Mechanical control on this ice stream emanates almost entirely from the lateral margins.

A mathematical model for simulating the densification of bubbly glacier ice is used to interpret the following experimental data from the Vostok (central Antarctica) ice core: two ice-porosity profiles obtained by independent methods and a bubble-pressure profile obtained by direct measurements of air pressure within individual bubbles. The rheological properties of pure polycrystalline ice are deduced from the solution of the inverse problem. The model and the inferred ice-flow law are then validated, using porosity profiles from seven other ice cores drilled in Antarctica and Greenland, in the temperature range from -55° to -20° C. The following expression is adopted for the constitutive law:

where ė and τ are the effective strain rate and stress, respectively, α is the creep exponent taken as 3.5, Rs is the gas constant and T(Ts) is the temperature (standard temperature). The numerical values obtained for the “linear” and “non-linear” viscosities are: μ1 = 2.9 ± 1.3 M Pa year and μ2 = 0.051 ± 0.019 M Paα year, and the apparent activation energy Q is confirmed to be 60 k J mole−1. The corresponding flow law is in good agreement with results of both mechanical tests and independent estimations based on the analysis of different natural phenomena associated with glacier-ice deformation. When the model is constrained by the porosity and bubble-pressure profiles from Vostok, the mean air content in Holocene ice is inferred to be about 0.088 cm3g−1. The corresponding mean air pressure in bubbles at the end of pore closure is about 0.083 M Pa, whereas the atmospheric pressure at this depth level would be 0.063 M Pa. The influence of the climatic change on the ice-porosity profile is discussed. It resulted in an increased air content in ice at Vostok during the Last Glacial Maximum: 0.096 cm3 g−1.

Twenty-seven studies on the thermal conductivity of snow (Keff) have been published since 1886. Combined, they comprise 354 values of Keff, and have been used to derive over 13 regression equation and predicting Keff vs density. Due to large (and largely undocumented) differences in measurement methods and accuracy, sample temperature and snow type, it is not possible to know what part of the variability in this data set is the result of snow microstructure. We present a new data set containing 488 measurements for which the temperature, type and measurement accuracy are known. A quadratic equation,

where ρ is in g cm−3, and Keff is in W m−1K−1, can be fit to the new data (R2 = 0.79). A logarithmic expression,

can also be used. The first regression is better when estimating values beyond the limits of the data; the second when estimating values for low-density snow. Within the data set, snow types resulting from kinetic growth show density-independent behavior. Rounded-grain and wind-blown snow show strong density dependence. The new data set has a higher mean value of density but a lower mean value of thermal conductivity than the old set. This shift is attributed to differences in snow types and sample temperatures in the sets. Using both data sets, we show that there are well-defined limits to the geometric configurations that natural seasonal snow can take.

Exposure of a full transverse cross-section of the terminus of Fireweed rock glacier, Wrangell Mountains, Alaska, revealed a thin layer of unconsolidated debris mantling a consolidated mélange of ice and rock. The main rock glacier is fed by three tributaries; at the terminus, contacts between the three are sharply defined. Ice content is >50% by volume. Bubble foliation and crystal morphologies of the ice matrix are similar to those reported from glacier ice. Folded ice-rich strata and lenses, foliation planes, and the long-intermediate axial planes of tabular-shaped englacial clasts dip sleeply toward the center line of the rock glacier. The planar structures generally parallel the steep walls of the gorge containing the trunk stream. These steeply dipping, longitudinal structures appear to result from transverse compression where the tributaries converge and the trunk stream narrows down-valley.

Bergschrund-like and moulin-like features at the heads of the middle and west tributaries, respectively, exposed admixed ice and rock similar to that of the terminus but lacking the well-developed flow texture. Talus from the unstable cirque headwall and polygenetic ice both accumulate at the base of the headwall and nourish the tributary rock glaciers.

Measurements from the subarctic snowpack are used to explore the relationship between grain growth and vapor flow, the fundamental processes of dry-snow metamorphism. Due to extreme temperature gradients, the subarctic pack undergoes extensive depth-hoar metamorphism. By the end of the winter a five-layered structure with a pronounced weak layer near the base of the snow evolves. Grain-size increases by a factor of 2–3. while the number of grains per unit mass decreases by a factor of 10. Observed growth rates require significant net inter-particle vapor fluxes. Stable-isotope ratios show that there are also significant net layer-to-layer vapor fluxes. Soil moisture enters the base of the pack and mixes with the bottom 10 cm of snow, while isotopically light water vapor fractionates from the basal layer and is deposited up to 50 cm higher in the pack. End-of-winter density profiles for snow on the ground, compared with snow on tables, indicate the net layer-to-layer vapor flux averages 6 x 10−7 kg m−2 s−1, though detailed measurements show the net flux is episodic and varies with time and height in the pack, with peak net fluxes ten limes higher than average. A model, driven by observed temperature profiles, reproduces the layer-to-layer flux pattern and predicts the observed weak layer at the base of the snow. Calculated layer-to-layer vapor fluxes are ten times higher than inter-particle fluxes, which implies that depth-hoar grain growth is limited by factors other than the vapor supply. This finding suggests that gain and loss of water molecules due to sublimation from grains takes place at a rate many times higher than the rate at which grains grow, and it explains why grains can metamorphose into different forms so readily.

A power law relation between stress and strain rate of the form was used to describe the response to strain rate of S1 ice loaded across the columns at –10° C. The rate exponent, n, decreased with increasing strain from about 4.6 at an observed peak on the load displacement curve to approximately 2.6 at a shortening of 2%. Analysis of these results and of the results of other authors on different forms of ice deformed at the same temperature suggests that the power law exponent, n is proportional to Fc/Fg The parameter Fc/Fg is the far-field basal dislocation climb force divided by the glide force.

Layering affects the air flow through snow caused by surface pressure variations. The horizontal and total fluxes are high in hoar layers but the pressure perturbations and vertical components of the flow do not penetrate as deeply as in homogeneous snow. That is because the layers “pipe” the flow horizontally toward the area of low pressure. An ice layer at the surface reduces the total flow everywhere. The flow decreases as ice-layer thickness increases and, in general, flow changes with permeability. However, the magnitude of the effect is proportionately weaker when the ice layers are further from the surface. The residence time is reduced when hoar layers are present due to shorter flow paths, reduced penetration into the deeper snow and higher speeds.

Experiments on plate-like specimens have established that across-column cracks from within S2 (columnar) salt-water ice when compressed uniaxially along a direction inclined to the long axis of the grains. Wing cracks initiate from the across-column cracks and lengthen into axial splits when the ice is rapidly deformed; correspondingly; the macroscopic behavior changes from ductile to brittle. The across-column cracking is attributed to grain-boundary sliding, and the splitting to the suppression of crack-tip creep.

Seismic reflection data from two lines on Rutford Ice Streem are presented and are compared with data already published from a third line on the ice stream. The amplitude and phase of the reflections have been used io investigate the properties ol the sub-ice material. Multiple reflections on long record-length data allowed calibration of the reflection coefficient at the ice-bed interlace and determination of the acoustic impedance of the bed material. The characteristics of the bed material vary both along and across the ice stream. The average acoustic impedance of the bed material across the glacier at the upstream line is 3.88 x 106 kg−2 s−1. This decreases to 3.19 x 10−6 kg m−2s−1 52 km further downstream. These values are within the rang which is typical of soft sediments. Using acoustic impedance as an indicator of subgiacial porosity, some areas of the ice-stream bed are interpreted as dilatant water-saturated sediments undergoing pervasive deformation. In other areas, the bed is not deforming and basal sliding may be a more important process. The proportional the ice-stream width over which bed deformation occurs increases downstream.

This paper examines the origin of moraine ridges in East Antarctic coastal oases and derives depositional models appropriate for the reconstruction of Quaternary history. On the basis of morphology, structure and sedimentology, four principal types of ridge may be identified: (1) type A moraines from when the basal debris zone crops out near an ice margin; (2) type B moraines form when large recumbent folds develop in the basal debris zone; (3) type C moraines are ice-contact screes and fans which form when debris accumulates at steep or cliffed ice margins; and (4) type D moraines are thrust-block moraines that form when unconsolidated sediment is entrained by freezing, shearing and thrusting of sediment blocks at the base of the glacier. Simple calculations of the rate of debris accumulation at ice margins suggest that type A, B and C moraines take thousands of years to form and record stable ice margins. Type D moraines are structural features that may form relatively quickly when ice margins override unconsolidated sediment. Constructing models to explain the origin of the moraines is an important part of reconstructing the Quaternary history of Antarctic coastal oases, because the models provide a basis for reconstructing the position and behaviour of the ice sheet during advance and retreat

Parameterizing the incoming longwave radiation L ↓ in terms of the fourth power of the absolute temperature at the reference height is used in glaciology for several purposes. In this paper, the validity of this kind of parameterization is investigated for the Greenland ice sheer, both by observations and by numerical simulation with a meso-scale model, It is found that such a parameterization severely underestimates the increase of L ↓ in response to large-scale warming in an area where surface melting is important. This is explained by the systematic influence that is exerted on the shape of the temperature profiles by surface melting.

To ascertain whether the velocity of Ice Stream B, West Antarctica, may be controlled by the stresses in its marginal shear zones (the “Snake” and the “Dragon”), we undertook a determination of the marginal shear stress in the Dragon near Camp Up B by using ice itself as a stress meter. The observed marginal shear strain rate of 0.14 a−1 is used to calculate the marginal shear stress from the flow law of ice determined by creep tests on ice cores from a depth of 300 m in the Dragon, obtained by using a hot-water ice-coring drill. The test-specimen orientation relative to the stress axes in the tests is chosen on the basis of c-axis fabrics so that the test applies horizontal shear across vertical planes parallel to the margin. The resulting marginal shear stress is (2.2 ± 0.3) × 105 Pa. This implies that 63–100% of the ice stream’s support against gravitational loading comes from the margins and only 37–0% from the base, so that the margins play an important role in controlling the ice-stream motion. The marginal shear-stress value is twice that given by the ice-stream model of Echelmeyer and others (1994) and the corresponding strain-rate enhancement factors differ greatly (E ≈ 1–2 vs 10–12.5). This large discrepancy could be explained by recrystallization of the ice during or shortly after coring. Estimates of the expected recrystallization time-scale bracket the ∼1 h time-scale of coring and leave the likelihood of recrystallization uncertain. However, the observed two-maximum fabric type is not what is expected for annealing recrystallization from the sharp single-maximum fabric that would be expected in situ at the high shear strains involved (γ ∼ 20). Experimental data from Wilson (1982) suggest that, if the core did recrystallize, the prior fabric was a two-maximum fabric not substantially different from the observed one, which implies that the measured flow law and derived marginal shear stress are applicable to the in situ situation. These issues need to be resolved by further work to obtain a more definitive observational assessment of the marginal shear stress.

The basal deformation of a gravity-driven linear creeping flow sliding frictionless over slowly varying bed undulations in two dimensions is analysed analytically, using results from second-order perturbation theory. One of the key results is that, close to sinusoidal bedrock undulations, up to two different spatial regions of local extrusion flow may arise. The offset and onset of extrusion flow is controlled primarily by the amplitude-to-wavelength ratio. Above the crest of a sinusoidal bed line, a local maximum of the surface-parallel velocity develops for ε : = ak < 0.138, where a is the amplitude and k is the wave number. As ε increases from zerо to this critical value, the vertical position of the velocity maximum moves from kz = 1 to kz ≈ 1.98, where z is the vertical distance above the mean bed line. Within and above the trough of a sinusoid, a region of local minimum of the surface-parallel velocity component develops, which shifts from kz = 1 towards the bed line as ε increases front zero to 1/2. Below this velocity minimum, and for some distance above the velocity maximum, the surface-parallel velocity increases with depth. This type of extrusion flow will cause a reversal of borehole-inclination profiles close to the bedrock.

The aim of this paper is to investigate the geophysical characteristics of the Antarctic ice sheet using radar altimetric observations. To do this, we use an altimetric waveform simulator, in situ observations, ERS-1 (European remote-sensing satellite) data and SPOT (Satellite pour l’observation de la terre) images. The small-scale study takes place at Dome C, Terre Adélie, which is a relatively flat region with gentle undulations and low wind speed. Despite this, the altimetric waveform parameters (height, energy, leading edge and trailing edge) are highly noisy. The effect of undulations on the waveform parameters is found to be dominant. The combination of a subsurface signal and a rough surface produces a linear effect on the altimetric backscattering or on the trailing edge of the waveform, but a strongly non-linear effect on the leading edge of the waveform or height estimation. As a consequence, the height measurement is very sensitive to the altimeter technical or orbital characteristics and is not reproducible from one mission to another. Observations show sastrugi fields that enhance the leading edge and affect the whole waveform. Observed local backscattering changes, probably due to local variations in surface microroughness, enhance the backscattered energy and may artificially create a topographic signal. The continental-scale study shows coherent patterns. Even if both surface and subsurface components affect the altimetric observation, the large-scale signal is mostly controlled by surface backscattering variations. The surface or near-subsurface characteristics of the snowpack may then be reached by altimetric observations.