Seismic reflection surveys

The acquisition parameters and data processing were the same for both New and Young seismic reflection lines. The seismic source was 300 g of high explosive, placed in 20 in deep holes produced with a small hot-water drill. Shots were Spacod at 120m intervals and were detected by 24 vertically orientated geophones (natural frequency 40Hz) at 10m spacing. This produced single fold data with reflection points Spaced 5 m apart on the bed. Offset from the shot to the first geophone was 30 m. The recording instrument (BISON 9024) sampled ev ery 0.2 ms.

The data were adjusted for slight variations in the shot depths and also for the surface topography, assuming thai the thickness of the firn was constant along each line. A normal move-out correction was applied, which accounts for the different distances from a given shot to each of the 24 geophones. A predictive deconvolution filler was designed for each shot and applied to remove the Surface ghost reflection (this is energy which travels upwards from the shot before bouncing off the surface and going downwards, a short time after the initial downward-travelling shot energy: see Fig. 3). The coefficients and the prediction distance for the dcconvolution filter were designed such that the ice bed reflection was not altered. Finally, the data were migrated to collapse numerous hyperbolas (Reference Smith and OerlerSmith, 1994) and to position reflections correctly.

Fig. 3.

Diagrammatic ray-paths for tee bed reflection (I₁). ice bed ghost (I₁ghost) and first iee multiple (I₂)

Wide-angle and shallow refraction surveys

At a number of sites along both New Line and Young Line, seismic wide-angle and shallow-refraction surveys were conducted to complement the reflection data, as discussed later. The wide-angle experiments recorded data from source receiver offsets ranging from 10m to over 4km. using charge sizes of 750 -2250 g. For the shallow-refraction surveys, small surface sources (either a single detonator or a 50 g charge) were recorded out to source receiver offsets of 440m, to provide information on the velocity and density structure of the tipper 100-120 m of snow and firn.

Ice-bed reflections

The two final processed seismic reflection sections are presented in Figure 4. The arrival labelled I₁(after Reference Crary, Rabinson, Bennett and BoydCrary and others. 1962) is the reflection from the base of the ice. Along both lines, this reflection is usually easy to identify, indicatinga significant acoustic impedance contrast at the ice-bed interface. The acoustic impedance, Z. of a medium is defined as:

(1)

where ρ is its density and V is its seismic velocity If the acoustic impedance of the bed (Z_b) is greater than that of the ice (Z_i). the polarity of the resulting I_ireflection will be normal. Conversely, if Z_b, is less than Z_ithe resulting I₁reflection will be reversed in polarity. The polarity ofthe ice bed reflections was determined by comparisons with direct waves and with reflections from ice-water interfaces, both of which can be predicted. The I₁reflection for all of New Line and for most of young Line is normal polarity, for the eastern end of young Line, the I₁reflection is reversed in polarity.Young Linchad been divided into western and eastern sections on this basis (Fig. 4b).

Fig. 4.

Processed seismic reflection sections for (a) New Line and (b) yong Line, (Note the different horizontal and vertical scales.) I₁is the ice-bed reflection. WA marks the positions of the wide-angle surreys. Box on New Line shows the section enlarged in Figure 5. The first part of a normal polarity reflection is white (unfilled); the first part of a reversed polarity reflation is black (filled).

In a number of places along both New Line and young Line, the normal polarity I₁reflection is followed after 10-20ms by another strong, normal polarity arrival (labelled I₁in Figure 5). It is not obvious whether this later arrival is a reflection from within the bed or else a rellection from another part ofthe ice bed interface, of to one side of the reflection line. A short seismic reflection line was acquired, intersecting New Line and perpendicular to it (that is, parallel to the ice flow direction), The acquisition parameters for this additional line were slightly differeint to those for New and young lines, resulting in a poorer-quality section. However, study of the I₁and I₁′ reflections before and after migration (Fig. 6) shows that, at least at this point on New Line, both these arrivals come from the ice bed interface. One arrival (I₁′) comes from the interface directly below the intersection of the two seismic lines, whereas the other (I₁) comes from a sloping part of the interface, just downstream of the intersection. The similarity between these two arrivals at this point and where they occur elsewhere along New Line suggests that they usually represent two reflections from the ice-bed interface.

Fig. 5.

Enlarged section of ice bed reflection on New Line, where it intersects with the seismic reflection line parallel to ice flow. I₁is the earlier reflection: I₁’ is the later one.

Fig. 6.

Part of the seismic reflection section, parallel to ice flow which intersects. New Line. Upper section is before migration; lower section is afterwards. Ice flow is from right to left.

There are strong similarities between this pattern of two arrivals (I₁and I₁′) on New Line and a similar pattern in places and young Line (below WA in Figure 4b. for example). This might be taken to suggest that, on Young Line too, both the arrivals come from the ice-bed interface. However. young Line is much closer to Tyree Line (12 km away) than to New Line (40 km), and a reflection survey perpendicular to Tyree Line (see Reference SmithSmith, 1996. Fig. 3b) showed a very smooth, level bed. very different to that in Figure 6. Hence, although it is likely that the I₁′ reflections on Young Line do come from the ice-bed interface, this conclusion cannot yet be proved.

In the following analysis the I₁reflections are studied in detail. The I₁′ reflection has not been used in the analysis because of possible corruption by energy immediately following I₁. Even where the I₁′ arrival is actually the one which is directly beneath the seismic line (as in Fig. 6). this is not considered to compromise the analysis, as the reflection points on the bed are close to each other and show reflections of similar strengths.

Ice thickness and bed elevation

Ice thickness was calculated using a seismic velocity determined in the following way. Reduction of the shallow-refraction seismic experiments, following the method ol Reference Kirchner, Bentley, Bentley and HayesKirchner and Bently (1990). gave a mean seismic velocity of 2839 ms⁻¹for the snow and firn in the upper 100 in of the ice column (Reference Smith and DoakeSmith and Doake, 1994) A mean seismic velocity for the rest of the ice column of 3841 m s⁻¹was calculated Iruni Kohnert’s (1974 temperature velocity relationship. This calculation assumed a constant temperature in the upper half of the ice column (−27° C; Reference Jenkins and DoakeJenkins and Doake. 1991; personal communication from R. M Frolich. 1996), followed by a lineariucrease in temperature to the pressure-melting point (−1.5°C) at the base. The seismic velocity in the ice column is accurate to ±15⁻¹; ice thicknesses are accurate to ±7 in. Surface elevation above sea level (based on the WGS72 geoid model) was determined by Doppler satellite surveying (accuracy ±5m) at selected locations, combined with standard optical survey techniques. Ice thickness on New Line ranges from 2335 m a1 the western end of the line to 2036m close to the middle. Bed elevation is lowest (−1839m) and highest (−1540m) at lhese two points, respectively (the equal surface elevation at diese two points is coincidental). The ice-bed interface of Young Line is smoother. Ice thickness is greatest (2190m) at the western end and least (2121m) close to the eastern end. Corresponding bed elevations at these points are −1880 and −1798 m, respectively.

Reflection coefficient at the ice-bed interface

The energy E₁of an ice bed interface reflection I₁; (the energy of a wavelet being the sum of the squared amplitude values) is given by (Reference RoethlisbergerRoethlisberger 1972):

(2)

where E₀is the initial energy of the seismic wavelet at the source. R is the reflection coefficient at the ice bed interface, ice thickness is h. and nais the attenuation within the ice. Record lengths Iiir the wide-angle surveys were sufficient to record the first ice multiple reflection (I₂Figs 3 and 7). Assuming that the reflection coefficient at the snow air interface is −1, Equation 2can be used to give the change in energy between the I₁and I₂reflections (Reference RoethlisbergerRoethlisberger, 1972):

(3)

where E₂is the wavelet energy of the I₂. reflection. Wavelet amplitudes (and hence, energies) on the wide-angle data were measured by identifying the high-amplitude central lobe and the adjacent lobe on each side. Zero-crossings at the start and end of these three lobes were used to delineate the wavelet.

Reference SmithSmith (1997)presented a compilation of attenuation data from a number of.sites. Attenualion is related to the ice temperature. Estimating the mean ice-column temperature on Rutford Icc Stream to be −20°C suggests that Ote correct attenuation value there is 0.21 x 10⁻³m⁻¹(see Reference SmithSmith, 1997. Table 1) . An error in the estimated mean ice temperature (it could be warmer than −20°C but is unlikely to be colder) would lead to an underestimate of t he attenuation value. However, a mean ice-column temperature up to 5°C wanner than that estimated (the error is unlikely to be greater than this) would lead to the magnitude of the calculated reflection coefficients being mo low by less than 10%. Reference SmithSmith (1997)also show cd that, even with unrealistically high and low attentuation estimates, the interpretation of reflection coefficients and derived acoustic impedance values remains unaffected.

Table 1.

Mean values of acoustic impedance in the different sections of the bed

Equation 3is strictly, only valid fora reflection perpendicular to a plane interface, so only data from geophones relatively close to the shots were used (maximum offset 460m), corresponding to a maximum incident angle of around 6° . At each wide-angle site two sites on New Line and one on Young Line; Fig. 4), E₁/E₂was measured from 16 separate seismic traces (8 traces from each of two shots), and a mean value calculated. Using Equation 3, mean values of R¹were then calculated for each of the wide-angle sites. From these, R was determined, the sign being given by the observed polarity of the I₁reflection.

Knowing the reflection coefficient at the positions of the wide-angle sites, the way it changes along the reflection line can be calculated using the variation in reflection strengths along the line. Foreaeh shot on the seismic reflection lines, the I₁, wavelet on all 24 channels was averaged and its energy (E₁) measured. Knowing R at a wide-angle site allows a value of E₀to be calculated from Equation 2, using a value for E₁measured from the reflection line at that point. It is necessary to assume that the initial energy is the same !breach shot on the reflection line. Although this is unlikely to be strictly correct, as the coupling of the shots with the firn will have varied to some extent, it is reasonable to assume that, for the majority of shots, the initial energy imparted to the firn will be similar. The consistency of the strenght of both the I₁reflection and the surface ghost (before it was removed) between adjacent shots supports this assumption. Rcan then be calculated for each shot on the seismic reflection lines. Clalculaled reflection coefficient values of the ice-bed interface along New Line and Young Line are given in Figure 8.

Fig. 8.

Reflection coefficient (solid lines ) of the ice-bed interface along (a) New Line and (b) young line. Dashed lines are the bed elevation profiles.

An estimate of the accuracy oi the derived reflection coefficient values in Figure 8can he obtained from the two wide-angle sites on New lane. At each site, 16 traces were used to calculate the reflection coefficient, giving a total of 32 separate estimates of the reflection coefficient along the bed of New Line, from which a mean value was calculated. The resulting estimate of the rnss error in the reflection coefficients was . A similar error value was calculated for young Line, from 16 traces at the single wide-angle site.

Acoustic impedance of the bed: calculation and errors

The reflecion coefficient is relaled to the acoustic impedances of the ice and die bed (Z_iand Z_b. respectively) by:

(4)

Using the reflection coefficient data presented in Figure 8, the acoustic impedance of the bed material. Z_b, can be calculated from Equation(4) if the acoustic impedance in the iee Zi_b, is known. Reference Atre and BentleyAtre and Bentley (1993)discussed the estimation of Z_iin some detail, with particular reference to the Stream B, concluding that a value of (3.33 ± 0.004) x 10⁶kg ms⁻²s⁻¹was the best estimate. Basal ice temperatures on Rutford Ice Stream and Ice Stream B are expected to be similar and the seismic reflections from the beds of the two ice streems were of similar frequencies (200-300 Hz; Atrc and Bentley, 1993). so Atre and Bentley’s value for the acoustic impedance in the ice,Z_i, is used here. Ii is expected to vary little laterally (Reference Atre and BentleyAtre and Bentley 1993). The acoustic impedance of the bed along New Line and Young Line can thus be calculated (Fig. 9a and b.) The resulting values are similar to those which are typical of soft, water-saturated sediments.

Fig. 9.

Acoustic impedance of the bed along (a). New Line, (b) young Line and (e)Tyree Line (modifiedfrom Reference SmithSmith, 1997). The bands labelled “Till” (shaded) and “Ice” (white) and the ranges of porosity values are taken from Reference Atre and BentleyAtre and Bentley (1993). vertical bar shows the acoustic impedance of the bed at one site on Ice Stream B (Reference Blankenship, Bentley, Ronney and AlleyBlankenship and others. 1987b). The satellite image with the line positions and corresponding acoustic impedance plots is the same as that presented in Figure 2 but has been rotated 90° clockwise.

The main sources of error in ihe Calculated acoustic impedance values are the likely errors in lhe reflection coefficient and the assumed seismic velocities within the ice. These combine to give an rms error in the acoustic impedance of the bed of (where Z_b, > Z_i), though the error ran occasionally be more than this where the magnitude of the reflection coefficient is high. The calculated reflection coefficients may also include an unknown systematic error resulting Iroin an underestimate ol the attenuation value within the ice. This is unlikely to be greater than 10% and. as it would lead to the magnitude of reflection coefficients being underestimated, would imply that bed acoustic impedance values differ from those in the ice by even greater amounts than are shown in Figure 9. This means that the interpretation which follows is valid even where the rms error is high and even if there is an unknown systematic error arising from the estimated attenuation value.The overall conclusions arising from the interpretation can therefore be treated with confidence.

Relationship between acoustic impedance and sediment porosity

Figure 9also includes likely acoustic impedance values for ice and for water-saturated till. The range of values for till has been taken from data compiled by Reference Atre and BentleyAtre and Bentley (1993)and includes both deforming and non-deforming (lodged) till. Reference Atre and BentleyAtre and Bentley (1993)show how the acoustic impedance of likely till material varies for porosity values between 0.3 and 0.45 (see Reference Atre and BentleyAtre and Bentley. 1993. Fig. 12) This range of porosities is expected to be particularly relevant to subglacia] sediments. Lodged basal till has a porosity of ≤0.3, whereas shear deformation of saturated sediments causes dilation and an increase in porosity to around 0.4 (Boulion and Dent, 1974: Boulion and others. 1974; Boullon and Paul. 1976). Porosiiy can thus be used to distinguish between deforming and lodged subglacial till (e.g. Reference Alley, Blankensee, Bentley and RooneyAlley and others, 1986. Reference Alley, Blankenship, Bentley and Rooney1987). Corresponding porosity ranges, taken from Reference Atre and BentleyAtre and Bentley (1993), are included in Figure 9. Reference SmithSmith (1997)proposed that subglacial sediment porosity can be estimated from acoustic impedance. Although this is not a well-constrained relationship it is good enough to distinguish between porosity values of 0.3 and 0.4. Hence, the acoustic impedance can be used to distinguish between lodged till and till which is dilated and deforming.

New Line

The acoustic impedance of the bed along New Fine (Fig. 9a) is always greater than that of the ice and is mostly greater than or close to the top of the range for till. Implied porosily is mostly 0.3 or less. Il is therefore unlikely that this malerial is deforming to any significant degree. Ice-stream Ilow may be influenced more by basal sliding than by subglacial deformation at this site. Figure 10’(modified from Reference SmithSmith. 1997. Fig. 7) presents a compilation of measured values of velocity and density for saturated, porous sediments and poorly lithified Sedimentary rocks. Figure 10also includes a line of constant acoustic impedance determined from the average value along New Line (Table 1). Where this line intersects the region of values for sediments and rocks shows that the bed of New Line is not only at the upper end of the main range of values for sediments but also at the bottom end of the range for sedimentary rocks. Considering the soft sediment identified elsewhere beneath Rutford Ice Stream (Reference SmithSmith. 1997) and also beneath other ice streams (Reference Blankenship, Bentley, Rooney and AlleyBlankenship and others. 1986; Reference Engelhardt, Humphrey, Kamb and FahnestockEngelhardt and others. 1990: Reference Clarke and EchchelmeyerClarke and Echelmeyer, 1996), the bed material on New line is Likely to be lodged till rather than poorly lithified sedimentary rock. The data in Figure 10suggest that the average scismic velocity in the top few metres of the bed material is around 1750-2100 m⁻¹.

Fig. 7.

Example of data from one of the wide-angle surveys showing the I₁and I₂reflections for a single shot. A low-pass filler has been used to remove frequencies greater than 300 Hz from the traces. showing the I₂wavelet The amplitude of these traces has been increased slightly relative to the I₁refection, for clarity.

Fig. 10.

Published values of density and seismic velocity for sediments and rocks, compared with Iinrs of constant acoustic impedance determined for New Line and the various sections of young and Tyree Lines. Mndified from Reference SmithSmith (1997). Triangles are from saturated fresh-water sediments and from marine sediments, corrected to fresh-icalcr values (following Reference Blankenship, Bentley, Ronney and AlleyBlankenship and Others, l987b) where necessary. Sources are. Reference Nafe, Drake and HillNafe and Drake (1963). Reference MorganMorgan (1969)and Reference HamiltonHamilton (1970). Open circles are from poorly lithified Cenozoic rocks from drillholes around McMurdo Sound, Ross Sea (Reference Barrett and FroggattBarrett and Froggatt. 1978). Vertical bar is the velocity and density for the bed of Ice Stream B (Reference Blankenship, Bentley, Ronney and AlleyBlankenship and others. 1987b). Shaded areas show the lower ends of typical ranges for sedimentan rocks (e.g. Reference Gardner, Gardner and GregoryGardner and others, 1974).

Young Line

The bed of Young Line has two distinct sections (Fig. 9b): a western section with relatively high acoustic impedance, and an eastern section with a much lower acoustic impc-rlanee, mostly less than that for ice anil at the lower end of the range for till, implied porosity for the higher-impedance, western section is 0.3 or less, and the average acoustic impedance (Table 1) is very similar to that for New lane, as shown by the curve of the constant acoustic impedance labelled “Young Line (W)” Figure 10. This suggests a similar interpretation for the western section to that for New Line, namelv, a lied of lodged till with an average seismic velocity of around 1750-2100 m⁻¹in the top few metres of the bed. Acoustic impedance for the bed on the eastern section olAbimg Line is much less than that in the west (see the line labelled “Young Line (E)” in Figure 10), Implied porosity is around 0.4 or above. Such high porosity is believed to be possible only during sediment deformation (Reference Alley, Blankenship, Bentley and RooneyAlley and others. 1986), so the eastern section ol Young Line is interpreted as deforming subglacial sediment. T his interpretation is supported by the fact that the average acoustic impedance for the eastern part of Young Line is even less than that reported for the till beneath Ice Stream B (Fig. 10: Table 1), which is believed to be deforming (Reference Blankenship, Bentley, Ronney and AlleyBlankenship and others, 1987b). Seismic velocity in the bed material of the eastern section appears to be around 15001700 ms⁻¹.

Variations in bed characteristics between the three seismic reflection lines

Reference SmithSmith (1997)presented a detailed analysis of Tyree Line (Fig. 2), which can be divided into four sections. A-D (Fig. 9c). Two sections. A and C. have a low acoustic impedance which is believed to indicate subglacial deformation. These two sections are separated by another section, B, with a much higher acoustic impedance, except for a short section of low impedance, referred mas “the bump”. In section B, ice flow is believed to be influenced more by basal sliding than by subglacial deformation (Reference SmithSmith, 1997). The bump is believed to be similar to the two sections of subglacial deformation (sections A and C). A fourth section, D, has an acoustic impedance between that for section B and that for the other sections, but is still believed to exhibit subglacial deformation. lines of constant acoustic impedance for sections B and D and for the mean of A, C and the bump are included in Figure 10.

The pattern of acoustic impedance is very similar on on line (Fig. 9b) and on that part of Tyree Line which spans section B and the bump (Fig. 9c). The highest points on both Lines show low acoustic impedance values. On Tyree Line this point (the bump) is interpreted as a mound of deforming sediments, tentatively called a subglacial drumlin (Reference SmithSmith. 1997), resting ona harder substrate. Although the corresponding eastern section of Young Line does not appear to the directly upstream of the bump (sec satellite image in Figure 9), it is clearly possible that these two could be a continuous sub-ice feature. The western section of Young Line has a similar acoustic impedance to section B of Tyree Line, as well as to all of New Line (Fig. 10). Hence, although Young Line is relatively short (3.6km) the fact that it is situated close to Tyree Line suggests that the pattern of acoustic impedance at the bed beyond both ends of Young Line could be similar to that seen on Tyree Line. The bed in the region of both these lines (Young and Tyree) therefore has areas where significant subglacial deformation is occurring and other areas where it is not.

The similarities between Young Line and Tyree Line suggest that the boundaries between the areas of different bed character are aligned roughly in the direction of ice flow. Assuming that the highest point on Young Line and the bump on Tyree Line are parts of a straight, continuous feature, then this Ieaiure is orientated approximately 10° to the ice-flow direction. A second estimate of the orientation of bed features, relative to the flow direction, can be made from the hyperbolas which remain on lhe seismic sections (Fig. 4; Reference SmithSmith, 1997, Fig. 3). The fact that a number of hyperbolas remain on the seismic sections after migration could be due to features which are not perpendicular to the seismic line. Re-migration of the sections using a progressively higher velocity value can be used to estimate what the difference in orientation might he. The migration velocity (V_m) required to collapse a hyperbola is related to the angle between the seismic line and the strike of a featues(α) by:

In this way, it was determined that some of the features at the bed could be orientated up to 13 to the direction of the seismic line. Reference Atre and BentleyAtre and Bentlcy (1994)found similar orientations of bed features (up to around 20°) on the ice plain of Ice Stream B.

The pattern of acoustic impedance on New Line is clearly different to that on both Tyree Line and Young Line. Along most of New Line, there is no evidence for pervasive subglacial sediment deformation. Il is therefore similar to only the relatively high-impedance parts of Tyree Line (section B) and Young Line (western section). Further downstream, at Tyree and Young Lines, sediment deformation occurs across a significant proportion of the ice-stream width (almost three-quarters of the width of Tyree Line). Reference SmithSmith’s (1997)analysis of Tyree Line concluded that significant variations in bed characteristics occurred across the width of the ice stream. The comparison between Tyree Line, Young Line and few Line shows that significant along-flow variations in bed characteristics also exist on RutIiird Ice Stream.

Comparison with ice-bed reflections on Ice Streams B and C

Reference Atre and BentleyAtre and Bentley (1993, 1991)found that seismic reflections from the bed of Ice Streams B and C (Fig. 1) varied between normal and reversed polarity.They also found that, qualitatively, the reversed-polarily reflections ranged from strong to weak, whereas the normal-polarity ones, beneath Ice Stream 1B at least, were weak. Reference Atre and BentleyAtre and Bentley (1993)concluded that reflection polarity alone cannot distinguish between a dilated, deforming bed and an undilated. non-deformity one. In contrast to Ice Stream B, the strength of both normal- and reversed-polarity reflections from the bed of Rutford Ice Stream varies from strong to weak. The quantitative analysis of reflection strength shows thai the acoustic impedance of some parts of the bed of Rutford Ice Stream is too high for the bed malerial to be deforming. Combining the reflection polarity analysis with a quantitative amplitude analysis therefore allows die distinction between deforming and non-deforming beds to be he made.

Reference Atre and BentleyAtre and Beniley (1993)suggested thai some of the variations in acoustic impedance at the bed of Ice Streams B and C could be due to relatively minor changes in die porosity and/or sediment type of a dilatant, deforming bed. This may also explain some of the variations seen within the low-impedance parts of the bed of Rutford Ice Stream (the difference between sections C and D on Tyree Line, for example). However, such minor changes are insufficient Io explain the high-impedance parts of the bed of Rutford Ice Stream. These can be explained by relatively large changes in the porosity and the degree of deformation of the subglacial sediments.

Sediment supply

The areas of subglacial deformation are evidence dial there is a supply of at least some deformable sediment beneath Rutford Ice Stream. It is possible that this represents the total amount of sediment which Ls available. At the other extreme, it is possible that the whole of the ice-stream bed represents a source of sediment which is available to be mobilised (if not already mobile) under suitable conditions. This latter condition is considered more likely on the basis of the acoustic impedance of die bed material (Figs 9 and 10; Table 1). All of New line and parts of both Tyree and Young Lines hav e relatively high acoustic impedance values. However, these values are still low relative to other rock types, even poorly lithified ones. There is probably very little difference between sediments classified as lodged till and those classified as weak sedimentary rocks, in terms of how easily they can be eroded, mobilised anil deformed. The implication is therefore that the ice-stream bed beneath all of the seismic lines would be relatively easily deformed, given suitable conditions. It seems likely, though not certain, that lack of available sediment is not the reason for the differences in bed characteristics.

Water supply

There are a number of reasons why a lack of water at the ice-stream bed is considered very unlikely. The fast flow of Rulford Ice Stream and low basal shear stress indicate that the ice is not frozen tothe bed and that the base must be close to the pressure-melting point. Sources of heal at the bed include geothermal heat flow. strain-heating due to deformation within the ice (or the bed) and friction as the ice slides over the bed. It is likely that the ice at the bed is melting and supplying water lo whatever subglacial drainage system may be present. This is supported by the fact that water has been found at the beds of other ice streams (Reference Engelhardt, Humphrey, Kamb and FahnestockEngelhardt and others, 1990; Reference Iken, Echelmeyer, Harrison and FunkIken and others, 1993).

Effective pressure

It follows that the differences between the bed beneath New Line and that beneath Tyree and Young Lines are probably related to the amounts and distribution of subglacial water and the pressure of this water at the ice bed interface. For deformation to occur, the effective pressure must be low, but not zero. As deformation occurs along much of Tyree Line, the effective pressure there must he low. If the bed material behaves as a Mohr-Caiulomb material (Reference Boulton and HindmarshBoulton and Hindmarsh, 1987), the effective pressure must be of a similar order of magnitude to the basal shear stress (10-40kPa at Tyree Line: Reference Frolich, Mantripp, Vaughan and DoakeFrolich and others, 1987) for deformation to occur.

Beneath New Line there is no evidence for deformation. However, the seismic data cannot show whether this is because the effective pressure is too low or to high. If the effective pressure is zero the ice can lift off the bed, removing any coupling across the ice bed interface and precluding deformation. If, however, the effective pressure is too high, some shearing of the bed may occur but it is not expected to be pervasive and will not be associated with dilation or a significant increase in porosiiy. Either of those two possibilities would be consistent with the interpretation of the seismic data, but the high basal shear stress at New Line (around 100 kPa: Reference Frolich, Mantripp, Vaughan and DoakeFrolich and others. 1987: personal communication from R. M. Frolich, 1996) suggests that the effective pressure there is relatively high and decreases downstream. This decrease could be associated with the accumulation of meltwater and/or changes in the basal drainage system. Alternatively, effective pressure at New Line could be relatively low. with the basal shear stress distributed over localised sticky spots, particularly as bed topography in the region of New Line is relatively rough (Fig. 4a). Basal water pressure may also be very variable on a relatively small scale (Reference Murray and ClarkeMurray and Clarke. 1995).There is still considerable uncertainty in the pattern of effective pressure at the ice-stream bed.