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Large glacial lakes, including Glacial Lake Washburn, were present in the McMurdo Dry Valleys, Antarctica, during the Last Glacial Maximum (LGM) despite a colder and drier climate. To address the mechanism capable of generating enough meltwater to sustain these large lakes, a conceptual model was developed based on the warming potential of infrequent contemporary föhn winds. The model suggests that föhn winds were capable of generating enough meltwater to sustain large glacial lakes during the LGM by increasing degree days above freezing (DDAF) and prolonging the melt season. A present-day relationship between infrequent summer föhn winds and DDAF was established. It is assumed that the Taylor Dome ice core record represents large-scale palaeoclimatic variations for the McMurdo Dry Valleys region. This analysis suggests that because of the warming influence of the more frequent föhn winds, summer DDAF in the McMurdo Dry Valleys during the LGM were equivalent to present-day values, but this enhanced summer signal is not preserved in the annually averaged ice core temperature record.
A new deep ice-core drilling site has been identified in north Greenland at 75.12° N, 42.30° W, 316 km north-northwest (NNW) of the GRIР drill site on the summit of the ice sheet. The ice thickness here is 3085 m; the surface elevation is 2919 m.The North GRIP (NGRIP) site is identified so that ice of Eemian age (115–130 ka BP,calendar years before present) is located as far above bedrock as possible and so the thickness of the Eemian layer is as great as possible. An ice-flow model, similar to the one used to date the GRIP ice core, is used to simulate the flow along the NNW-trending ice ridge. Surface and bedrock elevations, surface accumulation-rate distribution and radio-echo sounding along the ridge have been used as model input.The surface accumulation rate drops from 0.23 m fee equivalent year−1 at GRIP to 0.19 m ice equivalent year−1 50 km from GRIP. Over the following 300km the accumulation is relatively constant, before it starts decreasing again further north. Ice thicknesses up to 3250 m bring the temperature of the basal ice up to the pressure-melting point 100–250 km from GRIP. The NGRIP site islocated 316 km from GRIP in a region where the bedrock is smooth and the accumulation rate is 0.19 m ice equivalent year−1. The modeled basal ice here has always been a few degrees below the pressure-melting point. Internal radio-echo sounding horizons can be traced between the GRIP and NGRIP sites, allowing us to date the ice down to 2300 m depth (52 ka BP). An ice-flow model predicts that the Eemian-age ice will be located in the depth range 2710–2800 m, which is 285 m above the bedrock. This is 120 m further above the bedrock, and the thickness of the Eemian layer of ice is 20 m thicker, than at the GRIP ice-core site.
10m firn temperatures are commonly used on the Antarctic plateau to estimate mean annual air temperatures. 10m firn temperatures measured at Taylor Dome (also referred to as McMurdo Dome in the literature), Antarctica, are influenced by a factor other than altitude and latitude that varies systematically across Taylor Dome. Some inter-related factors possibly contributing to the modern temperature variability are differences in sensible heat from warm or cold air masses, differences in wind strength and source region, differences in temperature inversion strength and differences in cloudiness. Our preliminary data are compatible with spatially variable katabatic winds that could control the winter temperature inversion strength to provide a large part of the signal. This has implications for paleoclimate studies.(1) Variations of the stable isotopes δ18O and δD from ice cores are a proxy for paleotemperature. The isotope thermometer is calibrated by comparing local isotope ratios with corresponding measured temperatures. In order to derive a useful isotope-temperature calibration, we must understand the processes that control the modern spatial variability of temperature. (2) In order to quantify past changes in local climate, we must understand processes that influence local spatial variability. If those processes differed in the past, ice-core climate reconstruction would be affected in two ways: through alteration of the geochemical record and through alteration of deep ice and firn temperatures.
Time scales for adjustment in the length of a glacier to changing climate may be described in terms of a relatively short time scale, TS and a longer memory time, Tm. The memory Tm represents the time scale needed for asymptotic approach of the glacier to a steady state following a climate change event. Tm is determined by simple continuity considerations concerning the total volume change that must occur to reach steady state and the balance rate that drives the change. We show that Tm is relatively independent of the size of the climate change or to details of how ice flow is related to the geometry of the glacier. The time scale TS represents the time between a climate event and the occurrence of substantial changes in the glacier length. We show that, in contrast to Tm, Ts is highly dependent on the size of the climate change and on details of ice dynamics. This dependence is investigated by several ice-flow models including a simple one in which ice transport is determined by local thickness and slope, as in the analysis of kinematic waves, and a finite element representation that fully includes longitudinal stress gradients. The ice-flow models are subjected to mass-balance perturbations of varying size — from small, for which linearization approximations are valid, to large, for which linearization breaks down.
The following behavior may be identified. Increasing the size of a mass-balance rate change causes a more rapid initial response of a glacier terminus, which tends to shorten Ts. Longitudinal stress gradients damp local variations in velocity and thereby slow the propagation and diffusion of kinematic waves and retard the response of the terminus, which tends to lengthen Ts. Longitudinal stress gradients transmit forces to the terminus region and influence the terminus motion without the necessity of redistributing mass from the glacier length into the terminus zone, which tends to shorten Ts. These various results indicate that accurate modeling of the short term responses of glaciers to climate change requires fairly sophisticated ice-flow models, However, for purposes of tracking glacier lengths (or areas) over time scales considerably great than Ts, fairly simple ice-flow models may suffice.
A hydrodynamic model of interface stability in a stratified fluid is reviewed. The model predicts that irregularities on the boundaries of a stiff layer, embedded in a soft matrix, are unstable in pure shearing flow, when compression is normal to the layer. Perturbations on such a layer can grow to form symmetric pinch-and-swell structures called boudins. The model predicts initial perturbation growth rates on the boundaries of an interglacial period ice layer. We find that, beneath an ice divide, irregularities on the Sangamon layer boundaries will not kinematically decay, as the layer thins. Finite-element modelling is used to determine the strain history of Sangamon ice beneath the divide at Summit, Greenland. That history suggests boundary irregularities have grown, relative to layer thickness, at least 26 fold over the past 90000 years. The result may be severe distortion or severing of the layer. Core holes penetrating the layer may recover anomalously thick or thin columns of ice resulting in erroneous environmental and climatic interpretations. Radio echo-sounding may be useful in searching for zones of boudinage, which should be avoided when coring. Initial perturbations might arise from mass-balance spatial variations or from transient flow fields.
Scheuchzer’s dilatation theory and Altmann’s rigid sliding theory were the first glacier-flow theories to receive serious scientific attention. When Agassiz began a research program at Unteraargletscher in 1839, he held several incorrect notions about glacier flow. Forbes understood the difficulties with the existing theories, and in the early 1840s he and Agassiz acquired motion, temperature, and structural data that were incompatible with the dilation and sliding theories but were suggestive of flow analogous to that of a viscous fluid. How an apparently brittle rigid solid like ice could flow became the central paradox requiring explanation. Some of the most able physicists of the mid-nineteenth century went through contortions in their largely misguided efforts to explain the viscous behaviour in terms of the known physics of rigid solids. Personality and speculation played a far larger part in their debates than we see in scientific discussions today.
Wave ogives arise in a solution of the continuity equation by the method of characteristics. Steady ice flow is assumed. Ice velocity, channel width, and mass-balance functions combine to form a wave-excitation potential that yields the forcing function for wave ogives. This linear-systems formulation extends the ogive theory of Nye. Convolution of the temporal cumulative mass balance and spatial forcing functions gives the total wave pattern below an ice fall. Many ice falls do not generate ogives because the wave amplitude is modulated by a factor related to ice-fall length. The wave ogives at Austerdalsbreen, Norway, are due almost entirely to ice acceleration at the top of the ice-fall, i.e. the same zone that King and Lewis showed was responsible for forming Forbes bands.
A 1984 strain net on the Snowdome of Blue Glacier showed that the surface slope is a good estimator of ice flow direction and divide location. Topographic maps from 1939, 1952, 1957, 1979, and 1984 show that the flow divide migrates within a zone up to 350 m wide, in response to changes in east-west gradient in snowfall. This zone encloses 6% of the Blue Glacier accumulation area and up to 10% of the year-end residual snow. An ongoing 28-year mass-balance study has used an extreme, westerly divide, giving systematically high net balance estimates. The correct catchment area, for a given balance year calculation, depends on the future migration sequence of the ice divide, with a time constant of about 30 years.
Climate studies using ice cores require knowledge of the ice deformation at a detailed level, obtainable only by integrated surveying and flow modelling. Field programs should consider model abilities and requirements at the planning stage. Strain and topographic surveys should enclose the flowlines to all boreholes and extend beyond. Only then is it possible to (1) calculate representative slopes at the drill sites and (2) use simple boundary conditions at locations where they do not affect the calculated flow near the holes. Mass conservation models, which may include a parameterized velocity field, estimate the imbalance between integrated accumulation and ice discharge. Momentum conservation models find the actual velocity field, and can reveal a more detailed flow history, but require detailed survey information for boundary conditions. A mass conservation model suggested that flow near core sites at Agassiz Ice Cap, Ellesmere Island, had been steady for more than 3000 years; however, a momentum conservation model showed that either the present transverse strain rate is much smaller than required by the mass conservation model, or the ice is much stiffer than accepted values. It also revealed transients in the flow and microclimate οf which the impact on the derived climate still needs to be assessed by integrated modelling and surveying.
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