Ice shelves form where ice flows off the Antarctic ice sheet onto the sea to produce rather flat slabs of floating ice which, for the theoretician, are the simplest of all large ice masses. Boundary conditions are well defined, conditions change very slowly over distances that are large compared with ice thickness, and horizontal velocities are independent of depth. Unconfined ice shelves can be used as giant creep machines to investigate the ice flow law at low stresses. Further inland, where movement is hampered by obstructions such as grounded ice rises and by shear between the ice shelf and its sides, the ice shelf transmits a backpressure which tends to restrict drainage from the ice sheets that feed it. Wastage from ice shelves is principally by calving and by bottom melting. There has been no direct measurement of bottom-melting rates, but indirect evidence suggests that, near the seaward edges of ice shelves, bottom-melting rates may exceed one metre per year, with significant melting within about 100 km of the ice front. Further inland there may be bottom freezing, and analysis of cores taken from the Amery Ice Shelf indicate that bottom-freezing rates average 0.5 m a–1 over a distance of 200 km. Such high freezing-rates are probably exceptional, and, beneath the Ross Ice Shelf, freezing appears to be insignificant even at a distance of 400 km from the ice front.

Because of their accessibility ice shelves have been studied in considerable detail, but many problems remain. In particular we need to improve our understanding of basal flux, ice-shelf dynamics near the grounding line, the calving of icebergs, and the state of equilibrium of ice rises. In addition there is a clear need for basic data from the Filchner-Ronne ice shelf.

As part of the Ross Ice Shelf Geophysical and Glaciological Survey, a detailed map of ice thickness has been produced from airborne radar measurements closely tied to the network of survey stations on the ice-shelf surface. The map, drawn with a 20 m contour interval, reveals a highly complex pattern of thickness variations reflecting presumably, at least in part, complex ice-shelf dynamics. Many features of the thickness variation pattern appear to be associated with zones of grounded ice, but not all. Features of interest include many ice thickness minima, with closures up to 120m; a narrow, greatly elongated ridge-trough system 450 km or more in length; a few ice thickness maxima; steep regional gradients of 10 m/km in freely floating ice; highly contorted contours suggesting a large-scale “turbulence”; and at least two remarkable step-like changes in ice thickness. The irregularity of many of these features suggests dynamic non-equilibrium, i.e. the existence of transients in the dynamic system, so that the ice shelf as a whole suggests a state of rather rapid change. Flow-bands constructed on the basis of the strengths of the echo from the ice-water interface clearly delineate the outflow from the main East Antarctic outlet glaciers in the grid eastern part of the shelf. A discontinuous flow band originating in a small mountain glacier (Robb Glacier) suggests a variable mesoclimate in the vicinity of the glacier within the last thousand years. Strong reflections near the ice front suggest bottom melting of saline ice previously frozen on to the underside of the ice. Several rifts or incipient rifts in the ice shelf characteristically show two lateral bands of strong reflections with a non-reflecting zone in between.

Radio-echo data have yielded information on the dynamics of the Ross Ice Shelf. Distinctive features present on the radio-echo records have been used to delineate several flow lines on the ice shelf. Measurement of the power returned from the ice-water interface has revealed regional variations in the r.f. dielectric loss of Ross Ice Shelf ice. These variations are used to indicate zones of bottom melting and freezing.

Seismic compressional-wave data from short refraction shooting carried out during the 1974–75 and 1976–77 seasons at J-9, the site of the Ross Ice Shelf Drilling Project, have been compared. Significant dissimilarities were found to exist between the two sets of data. The measurements were made at locations about 2 km apart, with three unreversed profiles 60° apart recorded during the 1976–77 season and one unreversed profile during the 1974–75 season. The resulting velocity–depth profiles, and hence the derived density–depth profiles, differ by as much as 8%, with the 1976–77 results indicating a maximum velocity, corresponding to solid ice, at a shallower depth than the 1974–75 data. Both profiles were subjected to the same analysis, and a comparison of travel-time curves shows the differences to be real. Densities measured on cores from a 100 m bore hole drilled in 1974–75 about 50 m from the center of the 1974–75 profile agree well with densities computed from that profile. The density difference is believed to be due to the passage of the ice through the high-stress system associated with the interaction between Ice Stream B, flowing in from the West Antarctic ice sheet, and the Ross Ice Shelf. A reversed refraction profile carried out at station B.C. about 30 km up-stream, shows evidence of dipping layers that may be similarly caused.

Seismic short-refraction studies were carried out at five stations on the Ross Ice Shelf during the 1976–77 summer season as part of the comprehensive Ross Ice Shelf Geophysical and Glaciological Survey. Measurements of the velocities of compressional waves were made at each location. Compressional wave velocities were measured along more than one azimuth at three sites, and shear wave velocities (both components) at two. Travel-time curves were fitted to an exponential expression by means of a non-linear least-squares regression technique. The errors in the apparent velocities are estimated to be about ±50 m s–1 at short distances, diminishing to about ±10 m s–1 near the ends of the profiles. Compressional-wave velocities show only slight variations with azimuth and only over certain depth intervals, showing that constant-velocity surfaces are essentially horizontal. Shear-wave velocities, however, exhibit large variations according to azimuth and polarization, indicating that transverse isotropy is violated at least in the upper 30–40 m of the ice shelf. It is believed that the anisotropy is caused by structural details in the firn perhaps modified by preferred crystal orientation and that it may arise at least partly from anisotropic stresses in the ice shelf.

During the 1976—77 season of the Ross Ice Shelf Geophysical and Glaciological Survey, a series of vertical electromagnetic sounding profiles of subsurface features was completed at station J-9. The survey comprised three five-kilometer north-west-south-east profiles separated by one kilometer and six two-kilometer north-east-south-west profiles, and was carried out on the surface using 35 MHz and 50 MHz radar systems. Folded-dipole antennae were used and oriented to detect reflectors both along and perpendicular to the profile path. This was done to facilitate the interpretation of the data, which indicated a complex system of bottom crevasses. Measurements of the positions, heights, and shapes of these crevasses showed at least two sets of crevasses varying in both strike and size. The larger crevasses, about 120 m high and oriented more or less normal to the flow direction, are probably associated with the movement of ice stream B across the grounding line between the West Antarctic ice sheet and the Ross Ice Shelf. A satisfactory explanation for the secondary set of crevasses, about 60 m high and forming an angle of 60° ±10° with the first set, has not yet been found.

Electrical resistivity sounding, using the four-electrode Schlumberger array, has been carried out at 11 locations on the Ross Ice Shelf. The apparent resistivity curves generally show four characteristic zones. The first, at distances from 1 to 10 m, reflects the near-surface zone of seasonal temperature changes and inhomogeneities. The second zone, from 10 m to 100 m, reflects primarily the increasing density with depth in the upper 50 m of the ice shelf, modified, in some locations, by temperature variations. The third zone, from 100 m to a distance roughly equal to the ice thickness, is affected principally by the temperature gradient in the solid ice. In the fourth zone, at distances greater than approximately twice the ice thickness, the apparent resistivity usually decreases rapidly with distance, owing to the highly conductive sea-water beneath the ice shelf. At some stations associated with ice streams and outlet glaciers, however, an increase at large spacings indicates much more resistive basal ice.

Using data from seven locations on the grid eastern half of the shelf that do not show obvious evidence of a basal resistive zone, including temperatures to 100 m at two of the sites, the mass-balance rate at the bottom of the ice is estimated to be within a few tenths of a meter per year of zero at distances between 90 and 530 km from the ice front, assuming steady-state condition over most of the ice shelf. However, the assumption of steady-state is questionable at locations close to outlet glaciers, and must be treated with great caution. The temperature measurements at the two sites, along with previously observed temperatures at the RISP drill site, make it possible to estimate the activation energy in the solid ice. The models fitted to the observed values suggest an “apparent” activation energy in the solid ice closer to 0.15 eV (14 kJ mol−1) than to 0.25 eV (24 kJ mol−1). This difference is believed to be due to a decrease in the ionic impurity content with increasing depth in the ice by a factor of two or more.

Seasonal variations in sodium concentrations have been measured on surface-pit snow samples and on firn and ice core samples from the Ross Ice Shelf, Antarctica. Site locations include J-9 (82° 22’ S., 168° 40’ W.), Roosevelt Island dome (79° 22’ S, 161° 40’ W), C-7-1 (78° 30’ S., 177° 00’ W.), and C-7-3 (78° 20’S., 179° 51’ E.). The predominant source for the Na is sea salt, indicated by greater concentration levels at seaward sites. Al concentrations of the order of only 10–9 g/g show that the input of continental dust is comparable to that at inland Antarctic locations, and that dust contributes only a negligibly small fraction of the Na on the shelf. Maximum Na concentrations occur in the winter or early spring, as is the case for Greenland ice. The annual accumulation-rate at J-9, determined by counting Na concentration peaks with depth, is 90 kg m–2 year–1, in agreement with rates determined radiometrically. Annual cycles in Na concentration are also detectable at depth in the J-9 ice core. It is suggested that Na concentration is a useful diagnostic criterion for distinguishing between East Antarctic ice (10–8 g Na/g), West Antarctic ice (30 × 10–9 g Na/g), and ice that fell as snow on the shelf itself (> 30 × 10–9g Na/g). The transition between snow that is chemically characteristic of the ice-shelf regime to snow of an inland regime is expected to occur near the 500 m elevation contour. This position is up to 200 km inland of the grounding line. A model is presented for the large-scale decrease in Na concentration with distance inland within the ice-shelf regime. Since deeper ice in Ross Ice Shelf cores originated generally further from the ocean, the up-stream origin of shelf ice may be estimated from the chemical profile. The Little America V ice-core chemistry profile shows no discontinuity as would be expected if a recent surge of West Antarctic ice had occurred.

The Weddell Sea pack ice undergoes several unique advance–retreat characteristics related to the clockwise transport in the Weddell Gyre, the physical setting for the pack ice, and the free boundary with the oceans to the north. From satellite-derived ice charts, the annual cycle of the pack ice advance and retreat is depicted. The Weddell pack advance is characterized by a strong east-moving component as well as the north advance seen in other regions such as East Antarctica.

Physical characteristics of the pack ice at the summer minimum ice edge are presented. Indications are that deformation is a significant component of the ice accumulation, deformed ice accounting for c. 15 to 20% of the area covered in the year-round pack. Ablation characteristics are inferred from observations made during field work and from satellite imagery. These observations indicate that surface-melt ablation typically seen on Arctic pack is not seen on the Weddell pack inside the summer edge.

Using the physical-property data and transport inferred from ship and iceberg drifts, a new annual ice accumulation η < 3 m is inferred over the continental shelf in the South compared to η < 2 m previously estimated (Gill, 1973). The implication is that salt flux into the ocean over the shelf may be significantly larger, thereby increasing the production of Western Shelf Water, a component of Antarctic Bottom Water.

This paper describes a simple model for predicting compactness over a given area. The model incorporates equations of continuity and momentum. The momentum equation is based on a balance between air-ice stress, water–ice stress, Coriolis force, and internal ice resistance. Since our interest is in short-range small-scale prediction, as a first step we have neglected sources and sinks terms in the continuity equation and assumed ice to be a film of Newtonian highly-viscous fluid. The solution of a simple equilibrium formulation is used as initial conditions.

Surface winds over the area of interest are obtained from an available model developed on the basis of similarity theory. The compactness model has been applied for a few real cases in summer 1975 over the Beaufort Sea area. The preliminary results of the model are encouraging. The main advantage of this model is that it can run on mini-computers available at most forecasting centres.

Since the winter of 1972–73, NOAA visible and thermal images of North America have been available for the study of ice on the Great Lakes. The shallowest of the Great Lakes is Lake Erie, on which ice occurs from December usually through April. The NOAA series of polar-orbiting satellites collect two thermal and one visible image per day; the Geostationary Observational Environmental Satellites (GOES) collect visible and infrared images every 30 min. NASA’s Landsat satellite has an 18 d revisit cycle (a 9 d cycle when two are operating), with superb ground resolution (79 m). Satellite images from all three satellites were used in this study.

Ice formation, movement, and break-up in Lake Erie are different in each sub-basin, Western, Central, and Eastern, owing to the variation in depth of water, geomorphology, and meteorologic factors.

Ice movement after initial break-up in 1974–77 resulted in the accumulation of ice at the entrance of the Niagara River at the eastern end of the lake. In 1973, satellite images show the unusual development of an ice dam across the entire width of Lake Erie and the “ice plug” at the mouth of the Niagara did not form.

The effect of wind direction on break-up patterns is seen to be critical to any attempt at forecasting date of complete ice melt as well as day-to-day distribution of ice. Satellite observation of ice in Lake Erie—and in all large lakes—provides information regularly, routinely, and synoptically from a point in space. These data are useful for the study of ice formation, movement, and break-up, and they provide an unrivaled and unprecedented data set of historical records.

A stability equation is derived for a model glacier of initially uniform thickness and of infinite extent transverse to the primary flow flowing without slip down an inclined plane. A stress-dependent power-law viscosity is wholly incorporated into the equations of motion. Stability of the glacier is tested against long-wavelength surface perturbations. Results for this initial formulation indicate that the glacier is stable against infinitesimal amplitude surface perturbations, although for certain variations of model parameters, the decay-rate of the disturbance becomes very slow, approaching neutral stability. Results are presented in terms of decay-rate magnitudes over a large range of perturbation wavelengths for many model glaciers in which bed slope, ice thickness, and ice rheology parameters are varied. For all models, the maximum decay-rate of the perturbation occurs at disturbance wavelengths of roughly three to six times the glacier thickness. Infinite-wavelength perturbations are found to be only neutrally stable. Long-wavelength perturbations propagate at a faster rate down-glacier than do the intermediate- or shorter-wavelength ones which tend to remain fixed on the glacier surface andride down-glacier with the primary flow as they decay.

We here describe the process by which a complex model set of equations and boundary conditions may be rationally reduced to a simpler and more manageable set by the. processes of non-dimensionalization and asymptotic approximation. Such a reduced model (derived elsewhere) is then presented for an incompressible, two-dimensional ice flow. It consists of two coupled equations for the stream function and enthalpy variable, together with a complex set of boundary conditions.

The important dimensionless parameters which arise are given, and various limiting values of these are commented on. Nye’s (1960) equation for kinematic waves may be reproduced, and a non-linear analysis of this reveals that disturbances reach the glacier snout in finite time, and are uniformly bounded there: in the particular case considered here, one can also show that the temperature field is stable.

It is shown that the effect of introducing a (realistic) sliding law which is continuously dependent on the temperature has a major effect on the bedrock temperature profile.

Lastly we consider seasonal waves using a kinematic wave equation based on a plausible form of the sliding law when cavitation is present. The main observed features are qualitatively reproduced.

The elongation of air bubbles discovered in a wide area of a glacier surface covered by supraglacial debris does not coincide with the direction of flow. When ice samples including elongated bubbles were subjected to simple-shear experiments, the elongated bubbles deformed passively with ice and their final orientation was a good indicator of the strain induced in the ice. Based on these experimental results, the strain and the velocity field of the glacier were deduced from the distribution pattern of the elongated bubbles. The results agreed with the measured flow velocity. The bubble foliation pattern could also be explained in terms of the passive deformation.

It appears that the well-known theory describing flow of glaciers and ice sheets over undulations is defective with regard to the precise form of the field equations and boundary conditions to be applied. In particular, when surface-wave phenomena are to be described the formulation of Budd does not seem to be applicable.

A side-scan sonar survey was conducted near the ice front of Riiser-Larsenisen during the 1976–77 Norwegian Antarctic Research Expedition to investigate the underwater ice-shelf morphology. The survey covered several shorter, widely-spaced sections, for a total of about 10 km along the ice front. Only a smaller part of the sections showed a smooth, straight ice front. Most of the observed ice fronts were very rough at the scale of decametres. Projections at around l00 m water depth commonly extended 50–200 m seaward of the above-water ice shelf, and had widths of similar dimensions.

Plastic deformation of ice shelves in the region along the grounding line where they bend with the tide can dissipate enough tidal energy to make them the most important single source contribution to tidal friction (Doake, 1978). Some of the effects are to slow the rotation of the Earth and to cause the moon to retreat from the Earth.

Systematic radio echo-sounding during three seasons since 1971–72 has produced data on the configuration of the ice sheet in East Antarctica. In the sector extending inland from southern Victoria Land, the ice sheet exhibits a large ridge which drives ice towards David Glacier in the north and Mulock and Byrd Glaciers to the south. Within 100 km of the McMurdo dry-valley region soundings along ten sub-parallel lines (c. 10 km apart) provides detail on ice surface and flow patterns at the ridge tip. A small surface dome lies just inland of Taylor Glacier. The surface drops by 100 m or more before rising to join the major ridge in East Antarctica.