Linear models of the relationships between meteorological observations and the flow of river Tungnaá at the western margin of glacier Vatnajökull were investigated by means of spectral analysis and estimation of the impulse response. Most of the variation of Tungnaá is confined to the lowest frequencies and the diurnal variations. The temperature has most effect on the rapid variations around 1 cycle/day whereas the largest coherences with the precipitation are in the lowest frequencies. The wind explains over 20% of the variations in the frequency range from 0–1 cycle/day, but this is partly due to its coherence with the precipitation. The time lag between changes in the temperature and the river is about 2 h, but the time lag between precipitation and the river is longer. Analysis of longer records of daily observations from Ϸjórsá shows that the coherence of the run-off and temperature increases at frequencies too low to be estimated from these data. At frequencies over 1 cycle/day most of the observed variations of the river cannot be explained by means of a linear relationship with the meteorological series.

Water flowing in tubular channels inside a glacier produces frictional heat, which causes melting of the ice walls. However the channels also have a tendency to close under the overburden pressure. Using the equilibrium equation that at every cross-section as much ice is melted as flows in, differential equations are given for steady flow in horizontal, inclined and vertical channels at variable depth and for variable discharge, ice properties and channel roughness. It is shown that the pressure decreases with increasing discharge, which proves that water must flow in main arteries. The same argument is used to show that certain glacier lakes above long flat valley glaciers must form in times of low discharge and empty when the discharge is high, i.e. when the water head in the subglacial drainage system drops below the lake level. Under the conditions of the model an ice mass of uniform thickness does not float, i.e. there is no water layer at the bottom, when the bed is inclined in the down-hill direction, but it can float on a horizontal bed if the exponent n of the law for the ice creep is small. It is further shown that basal streams (bottom conduits) and lateral streams at the hydraulic grade line (gradient conduits) can coexist. Time-dependent flow, local topography, ice motion, and sediment load are not accounted for in the theory, although they may strongly influence the actual course of the water. Computations have been carried out for the Gornergletscher where the bed topography is known and where some data are available on subglacial water pressure.

Mass-balance quantities at specific points on a glacier as defined in [IHD] (1970) relate either to annual maxima or minima in ice mass at that point (the stratigraphic system), or to values at the beginning and end of a hydrologic year (the annual or fixed-date system). Most quantities measured in the field relate to summer surfaces, which correspond to the annual minima at the measurement points. When stratigraphic system point values are integrated over a whole glacier, the result may be meaningless because annual maxima and minima and summer surfaces may form at different times at different places.

The combined system utilizes several kinds of data to derive meaningful area-average results that can be directly related to other hydrologic and meteorologic information. Measurements to summer surfaces at certain specific times, including the beginning and end of a hydrologic year, are added together with proper recognition of the types of material involved: old firn and ice, snow and superimposed ice of the year under study, new firn formed during that year, and late snow deposited toward the end of the year. Other “balance increment” terms relate values at the beginning and end of a hydrologic year to corresponding area-average balance minima. As a result, two types of “net balance” and many other terms are given precise meaning for a glacier as a whole. The scheme is sufficiently versatile to be used on any glacier, although the terms relating to summer surfaces are not defined on a glacier in which ablation or accumulation is continuous throughout a year.

The proposed method is to use a radio-echo sounder located on the surface of the ice sheet and to record the shape of the pulses returned from the rock bed. If the instrument is moved horizontally, the shape of the returning pulse will change, large differences occurring in horizontal distances of the order of the wavelength of the radiation, say 5 m. The spatial pattern of this variation, being largely determined by the details of the reflecting rock bed, is expected to remain fixed in position while the ice moves. It thus forms a fixed reference against which ice movement can be measured at places far removed from fixed landmarks. Analogue experiments in the laboratory, using ultrasonic waves in air in place of radio waves in ice, suggest that it should be easy to detect a movement of λ/10, 0.5 m, and, with further refinement, much smaller movements. A field test is in preparation.

The closure of water-filled glacier bore holes is considered and it is concluded that freezing due to conduction to the surrounding ice is usually the dominant process. On this basis englacial temperatures are obtained from a bore hole near the equilibrium line of Blue Glacier, U.S.A., where the ice thickness is about 125 m. Temperatures range from −0.03° C near the surface to −0.13° C at a depth of 105 m, with an estimated uncertainty of 0.02 or 0.03 deg. On the average the temperature is about 0.05 deg colder than the equilibrium temperature of ice and pure water. It is shown that at this temperature small amounts of water-soluble impurities play an important role in the thermal behavior of the ice. This leads to a new definition of temperate ice in terms of its effective heat capacity. The effective heat capacity of the Blue Glacier ice is apparently much larger than that of pure ice at the same temperature.

A network of passages situated along three-grain intersections enables water to percolate through temperate glacier ice. The deformability of the ice allows the passages to expand and contract in response to changes in pressure, and melting of the passage walls by heat generated by viscous dissipation and carried by above-freezing water causes the larger passages gradually to increase in size at the expense of the smaller ones. Thus, the behavior of the passages is primarily the result of three basic characteristics: (1) the capacity of the system continually adjusts, though not instantly, to fluctuations in the supply of melt water; (2) the direction of movement of the water is determined mainly by the ambient pressure in the ice, which in turn is governed primarily by the slope of the ice surface and secondarily by the local topography of the glacier bed; and, most important, (3) the network of passages tends in time to become arborescent, with a superglacial part much like an ordinary river system in a karst region, an englacial part comprised of tree-like systems of passages penetrating the ice from bed to surface, and a subglacial part consisting of tunnels in the ice carrying water and sediment along the glacier bed. These characteristics indicate that a sheet-like basal water layer under a glacier would normally be unstable, the stable form being tunnels; and they explain, among other things, why ice-marginal melt-water streams and lakes are so common, why eskers, which are generally considered to have formed in subglacial passages, trend in the general direction of ice flow with a tendency to follow valley floors and to cross divides at their lowest points, why they are typically discontinuous where they cross ridge crests, why they sometimes contain fragments from bedrock outcrops near the esker but not actually crossed by it, and why they seem to be formed mostly during the later stages of glaciation.

Ten-meter temperature measurements show that Athabasca Glacier is temperate in the accumulation area but not in the ablation area. An important factor in determining whether all the ice will reach a temperature of 0° C by the end of summer is how much of the layer of ice, cooled during the previous winter, is removed by ablation. However, calculations show that, even when allowance is made for ablation, not enough heat is conducted into the ice to bring it all to the melting point. As recent work suggests that ice at 0° C is permeable to water, latent heat released by refreezing of percolating melt water is considered; it appears likely that this process is an insignificant heat source in the ablation area. Available data show that the penetration of solar radiation can probably also be neglected. The question of how widespread temperate glaciers are is discussed; it is predicted that in most, if not all, glaciers there is a region below the equilibrium line where, because ablation is low, the glacier is not strictly temperate.

Cylindrical samples of ice with 0.0 to 0.35 volume fraction fine sand were tested in unconfined uniaxial compression at stresses between 5.3 and 6.4 bar and at temperatures between −7.4 and −9.4° C. Secondary creep rates were obtained from the slope of the total strain vs. time curve and were normalized to 5.6 bar and −9.1° C. Creep rates in ice with low sand concentrations were in some cases higher and in other cases lower than in clean ice. However at higher sand concentrations the creep rate decreases exponentially with increasing volume fraction sand. The latter results are in general agreement with theories developed to explain dispersion hardening of metals, and suggest that each sand grain is surrounded by a tangled network of secondary dislocations which impede passage of primary glide dislocations.

As the cause of a surge we assume a sudden improvement of the sliding conditions at the glacier bed. This improvement has the same effect on the behaviour of the glacier as an equivalent variation of the annual balance. We are therefore able to apply Nye’s (1963) theory of the advance and retreat of glaciers to our problem. A simple glacier model discussed in this theory shows as a consequence of an improvement of the sliding conditions, an additional discharge of ice, which can be separated in two phases. The first phase of strong motion yields a transport of ice from the upper to the lower part of the glacier. The ice essentially remains within the previous limits of the glacier. The sequence of motion during this phase corresponds to a surge. The second phase of motion is essentially slower. During this phase the ice accumulated in the lower part of the glacier flows beyond the previous limits, till the whole surface level is lowered.

The vast Bering piedmont glacier, which has large folds in the medial moraines in its terminal lobe, recently experienced two surges with a combined ice displacement of as much as 13 km. Vertical aerial photographs taken before and after the surges disclose the direction and magnitude of ice flow in various parts of the piedmont lobe. The ice moved toward the terminus and expanded in a normal, radial pattern with no evidence of unusual shearing that would result in the formation of large folds. Many surging glaciers display repeated lateral displacements in their medial moraines which result from periodic surging of the main glacier past non-surging tributaries. Moraines of Bering Glacier display small periodic irregularities of this nature. The large “accordion” folds in the moraines in the piedmont lobe are judged to be due to the combined effects of compressive flow and lateral or transverse expansion of these previously formed irregularities. The initially small pre-existing perturbations in the moraines are simply spread laterally and shortened radially into large folds as the ice spreads out.

A very large debris band composed of repeatedly folded medial moraines extends across the center of the Bering Glacier lobe. These remarkable folds are thought to result from the deformation of surge-related irregularities in medial moraines as they pass through the zone of intensive shear near the glacier’s margin.

Analysis of glaciologial data indicates that grounding of the Ward Hunt Ice Shelf and its conversion into an ice rise was primarily the result of local thickening of a floating ice shelf and the availability of a very gently sloping sea floor on which the ice shelf came to rest. Application of heat conduction theory to a series of thermal profiles through the Ward Hunt Ice Shelf, and the Ward Hunt, Camp Creek and Cape Discovery ice rises shows that present heat flow in this area of northern Ellesmere Island is more than twice normal, and that the outer and intermediate parts of the Ward Hunt ice rise grounded 250–350 years ago, during a cycle of climatic deterioration. Development and localization of ice rises along northern Ellesmere Island are strongly influenced by topography, and all ice rises we have studied seem to have formed within the past 1 600 years, possibly with major growth in the interval from 1 000 to 150 years ago.

In his now classic book L’dy Arktiki [Arctic ice], Zubov discussed the melting of sea ice during the Arctic summer by thermal interaction with the surrounding water and derived an expression which indicates that the proportion of open water increases exponentially with time until total ice-free conditions result. His equation predicts that the time required for complete decay of the ice cover after initial break-up is greater than one month and more likely as long as two months for representative values of incident shortwave radiation and initial ice thickness upon break-up. It is unlikely that above-freezing temperatures persist for this length of time.

To explain the observed complete disintegration of the annual ice cover in many sheltered areas of the Arctic, a modified model of the thermal decay process has been introduced. This model takes into account the influence of radiation absorbed by the ice which was not included in the Zubov formulation. Considerable reduction in the time required for complete decay, generally by about a factor of 2 if an albedo of 0.4 is assumed for the ice surface, is obtained.

In 1970 water pressure was measured in several moulins on the White Glacier. Pressure variations in some moulin channels extended over the full measuring range of the instruments (0–1 and 0–2 bar above atmospheric pressure), even at depths of less than 50 m below the surface. Measurements at different depths showed that total pressure variations were sometimes greater than this. The pressure data are compared with variations in the surface velocity of the glacier.

Under-ice sonar profiles and surface laser profiles of the Arctic pack ice have been analyzed using power-spectrum techniques to extract significant spectral peaks corresponding to spatial periodicies in the ice. The analysis suggests that, for a section of ice sampled by two intersecting under-ice profiles, the ridges are not randomly oriented. Moreover, the lineation or directionality of the ridges may be approximately determined from the two intersecting profiles. Also the spectra from surface profiles of multi-year ice and from surface profiles of first-year ice are of a much different nature, thus suggesting a technique for determining ice types from laser profiles.

The present surge of the Donjek Glacier on the north-east side of the St Elias Mountains, first noticed in 1969, is producing a number of morphological effects adjacent to the glacier in the terminus area. Although the effects of the surge are minimized by the lobate form of the glacier terminus, several types of push structure, erosional forms and certain drainage changes are being produced. These forms are similar to older forms close to, or on, the Neoglacial maximum moraine. It is considered that the similarities suggest that surges may have occurred throughout most of the Neoglacial period.

The rate of snow accumulation at “Byrd” station, Antarctica, was measured by various methods. Surface measurements yield a mean accumulation of 117 kg m−2 year−1 for the time interval 1962–70, within a 10 m radius of the station. The distribution of fission products with depth indicates a rate of 67 ± 4 kg m−2 year−1 for the period 1955–68 and of 100 ± 10 kg m−2 year−1 for the period 1965–68. The 210Pb method yields a 60 year average of 110 ± 10 kg m−2 year−1.

Multiple scattering of the solar flux in snow and “bubbly” ice can account for the variable albedo, the non-specular reflection, the non-exponential flux decrease near the surface, and the large upward flux within the medium. The scattering problem has been formulated and solved exactly, assuming isotropic scattering in a plane-parallel, semi-infinite, grey medium. The solution shows a non-exponential flux decrease near the surface and an exponential decrease deep in the medium. For such a medium, the albedo will increase with decreasing solar altitude in a manner which agrees to within one per cent of observed snow albedos in the Antarctic.

A study of the movement of Rusty Glacier was undertaken and continued through four summers because it is believed to be a surging glacier in the last stages of the inactive phase preceding a surge. The entire glacier is very slow moving, essentially motionless in the lower third and most rapid in an area well above the firn line. Unusually steep flow-line emergence angles and higher than average longitudinal compression rates in the lower–middle part of the glacier indicate gradual thickening of the ice above the stagnant lower tongue. There is no clear correlation between local variations in flow rates and surface or bottom topography. The glacier is mostly colder than 0° C to the bottom, and in the one known area of 0° C bottom temperature, flow rates are not greater than elsewhere. Although the glacier is everywhere very thin, maximum flow rates seem clearly related only to variations in ice thickness. The nearby Trapridge Glacier is also a surging glacier and exhibits a strikingly similar flow pattern.

Detailed field mapping in three massifs of the central Andes of Peru indicates that there are four glacial moraine stages. Historical evidence and correlation with Patagonia and South Georgia suggest that the three youngest stages relate to Neoglacial re-advances which culminated before 4000 b.p., between a.d. 1750 and 1800 and during the late nineteenth century. The oldest moraine stage may be of late Wisconsin/Weichselian age. The absence of older moraines suggests that the Peruvian Andes were not high enough earlier in the Pleistocene to support larger glaciers.

The existing theory of kinematic waves on glaciers does not account for self-induced instability of the flow, or for surges which are not the result of external disturbances. Surges do however have much in common with kinematic waves. Unstable behaviour of glaciers can be explained by a straightforward modification of kinematic wave theory. In this modification the relation between the ice flow and the ice depth is not the same when the ice is accelerating as when it is slowing down. A similar effect has previously been shown to explain observed instabilities in the flow of highway traffic, a phenomenon which is otherwise adequately described by a simple kinematic wave theory.