1. Lineations associated with near-surface diffractors

Radar profiles across four of the lineations (L2, L3, L5 and L6) reveal near-surface diffractors and distorted, often discontinuous deep internal layers (Fig. 3a; Table 1). GPS measurements across these features show broad topographic troughs that are 10–50 m deep over 3–10 km. MODIS and AVHRR images of the Duckfoot (Fig. 1b and c) both show L2 and L3 as bright, narrow lines (which are likely to be due to these troughs) aligned, respectively, parallel and sub-parallel to the most recent flow direction of KIS as indicated by flow stripes in the main body of KIS. Radar data show that the near-surface diffractors are closer to the surface at L2 than L3. Because diffractors here are not distinct, we pick the depth of the deepest continuous internal layer: 30 m for L3 and 18 m for L2 (Fig. 3c and b, respectively). Two sets of near-surface diffractors are detected beneath the lineations on the Goosefoot: one at L5 (22 m below the surface) and another at L6 (17 m below the surface) (Fig. 1a).

Fig. 3. (a) 2 MHz radar profile across L1–L4 along line A (Fig. 1a). The prominent return at approximately −100 m is a glitch in the transmitter. Data in the upper ~ 80 m (covered by a black band) are not resolved by the radar system. The solid grey line outlines the data shown in detail in Figure 4. (b) 100 MHz radar profile across L2. Arrows indicate the depth to the deepest continuous layer. (c) 100 MHz radar profile across L3. Prominent layers have been traced in each of the high-frequency radar profiles

Table 1. Main characteristics of features labeled in Figure 1 and described in text

We interpret these four lineations to be relict ice-stream shear margins now buried under >100 years of accumulation. The numerous near-surface diffractors associated with the lineations are interpreted to be buried crevasses that formed at the surface within a margin of an active ice stream. This interpretation is not new; Reference Retzlaff and BentleyRetzlaff and Bentley (1993) identified L5 as a relict margin of KIS, and Reference Jacobel, Scambos, Nereson and RaymondJacobel and others (2000) identified L3 as a relict margin. The deformed deep internal layer stratigraphy is consistent with conditions in an active ice stream where large cumulative strain is expected. Layer continuity here is lost due to some combination of high strain, causing steeply sloping layers, and signal loss from scattering within the near-surface diffractors. Deeper layers regain continuity about 10 km inside the ice-stream margin towards the middle of the ice stream. The discontinuous layers are probably related to the chaotic zones observed in active shear margins where strain rates are largest and crevassing does not conform to a particular pattern (Reference Echelmeyer and HarrisonEchelmeyer and Harrison, 1999).

Several characteristics of margins are apparent: the presence of near-surface point diffractors, deep discontinuous layers and the coincidence of these margins with broad topographic troughs, which are present across other active and relict ice-stream margins (Reference Whillans and Van der VeenWhillans and others, 1993; Reference Gades, Raymond, Conway and JacobelGades and others, 2000; Reference Conway, Catania, Raymond, Scambos, Engelhardt and GadesConway and others, 2002; Reference Catania, Conway, Gades, Raymond and EngelhardtCatania and others, 2003). One possible explanation for trough formation is that focused strain heating in margins allows for localized melting (Reference Jacobson and RaymondJacobson and Raymond, 1998). Reference Thorsteinsson, Waddington and FletcherThorsteinsson and others (2003) show that ice advected across an ice-stream margin undergoes rapid changes in the stress pattern, such that the direction of maximum compression becomes horizontal. They argue that ice expands vertically as it traverses an ice-stream margin but indicate that recrystallization and crevasse formation may counter this. The troughs that we observe across ice-stream margins probably represent some combination of these processes.

2. Lineations associated with no change in internal layers

L1 is a distinct, narrow lineation parallel to the most recent flow direction of KIS (Table 1). It appears to originate at the junction between the southern edge of Siple Dome and the innermost margin of KIS and fades out about 20 km before the KIS grounding line. Our measurements show that L1 is a low-amplitude (~ 2 m), long-wavelength (~ 2 km) undulation in surface topography, which appears as an alternating light and dark pattern in Figure 1a and b. Both high- and low-frequency radar profiles across L1 show deep disturbed layers and near-surface diffractors beneath the surface undulation.

The depth to near-surface diffractors does not change across L1, and so we do not think this is a relict margin of KIS, as hypothesized by Reference Jacobel, Scambos, Nereson and RaymondJacobel and others (2000). Instead we believe it is a relict flow stripe that originated from within an active shear margin. Reference Merry and WhillansMerry and Whillans (1993), figs 5 and 7) show several examples where flow stripes originate from margins of active ice streams, and Reference Gudmundsson, Raymond and BindschadlerGudmundsson and others (1998) show flow stripes with similar geometry to that of L1 in satellite images.

3. Lineations associated with syncline layers

Three of the eight lineations examined (L4, L7 and L8) are associated with shallow topographic troughs (~ 1–3 m deep) that span only a few kilometers (Table 1). On the surface, such small-scale topographic changes are not easily discernible but they are visible in satellite images, especially when taken at low Sun angles. These lineations are found at the boundaries between inter-ice-stream ridges and the flat-ice terrains in the mouth of KIS. In the Duckfoot, L4 wraps alongside the southern edge of Siple Dome, diverging away from the paleo-flow direction of KIS towards the grounding line (Fig. 1c). On the south side of KIS in the Goosefoot, L7 also wraps around the western end of Engelhardt Ridge away from the paleo-flow direction of KIS (Fig. 1a). L7 branches away from the edge of the ridge where it merges with L8 and cuts across the Goosefoot towards the grounding line (Fig. 1a).

Internal layers beneath these troughs are warped downwards as much as 200 m, forming a syncline that is unrelated to bed topography (Figs 4–6). The fold amplitude of these synclines increases with depth, and the deepest layers are truncated at the bed. The fold axis of the syncline in the Duckfoot is tilted away from Siple Dome such that the trough of the uppermost isochrone is offset horizontally by ~ 150 m from that of the deepest isochrone (Fig. 4). Assuming that the mechanism causing the downwarp maintained a consistent spatial pattern over time, this offset is probably caused by flow off Siple Dome after the syncline formed. If the present-day surface velocities (0.9 m a⁻¹; Reference Jacobel, Scambos, Nereson and RaymondJacobel and others, 2000) have been constant over the past few hundred years, the processes forming the syncline would have stopped about 170 years ago. Synclines in the Goosefoot (L7 and L8) have vertical fold axes (Fig. 6), implying that the processes causing these synclines may still be active or only recently ceased, although slope-driven flow across L7 might be too low to cause a noticeable tilt in the syncline axis. The most distinct characteristic of L7 is that it marks the location where the depth of deep diffractors changes (Fig. 5). We note that the B-line profile across L7 does not show downwarped layers (Fig. 5), but the other radar profile across L7 (not shown) does.

Fig. 4. 2 MHz radar profile along a section of line A (as shown in the box outlined in Fig. 3a) crossing L4 and the northernmost KIS margin (L3). The prominent return at approximately −100 m elevation is a glitch in the transmitter. Data in the upper ~ 80 m (covered by a black band) are not resolved by the radar system.

Fig. 5. 2 MHz data along line B in the Goosefoot crossing L7 and L8 (Fig. 1a). Note the prominent return at ~ 300 m depth is a glitch in the transmitter. Data in the upper ~ 80 m (covered by a black band) are not resolved by the radar system.

Fig. 6. 5 MHz radar profile along line C in the Goosefoot crossing L5, L7 and L8. Data in the upper ~ 80 m (covered by a black band) are not resolved by the radar system.

Several mechanisms could cause downwarping of internal layers. for example, Reference Vaughan, Corr, Doake and WaddingtonVaughan and others (1999) attribute near-surface downwarped internal layers across the boundary between an inter-ice-stream ridge and a possible relict ice stream to a localized high accumulation. Their data show that the amplitude of the layer syncline increases with depth in the upper 100 m of ice and they attribute this to a 33% increase in accumulation at the slope break between the ridge and the area of flat ice adjacent to it. Several layers across the upper 30 m (~ 350 years) of the Goosefoot and Duckfoot syncline-layer boundaries do not show changes in layer thickness that might suggest an accumulation anomaly here for at least the past few hundred years.

Localized basal melting can also downwarp internal layers over a narrow region. Basal melting can occur through several processes including melt beneath an ice shelf at the grounding line (Reference GillGill, 1973; Reference SmithSmith, 1996; Reference Rignot and JacobsRignot and Jacobs, 2002), melt from a localized high geothermal flux (Fahenstock and others, 2001) and melting due to strain heating within shear margins (Reference Jacobson and RaymondJacobson and Raymond, 1998). We do not think the observed stratigraphy is consistent with that of typical relict ice-stream margins, for several reasons: (1) high-frequency radar data across L4, L7 and L8 do not show near-surface point diffractors, while relict ice streams (e.g. ‘Siple Ice Stream’) as old as 420 ± 60 years (personal communication from B. Smith, 2000) still show evidence of near-surface point diffractors; (2) both L4 and L8 wrap around the edges of inter-ice-stream ridges away from the paleo-flow direction of KIS; (3) deep layers are preserved across these features, indicating that strain rates were not as high as those observed in present-day margins;(4) there is no evidence of flow stripes (indicative of sliding) on the surface in the areas adjacent to the syncline lineations; and (5) radar profiles across known relict and active shear margins do not show localized downwarping to this degree. Rather, the pattern of deeper layers across ice-stream margins appears to be disorganized (Reference Gades, Raymond, Conway and JacobelGades and others, 2000; Reference Catania, Conway, Gades, Raymond and EngelhardtCatania and others, 2003).

Variable geothermal heat sources could potentially produce the observed melt signature. The syncline-layer boundaries appear to be associated with the edges of the inter-ice-stream ridges which may represent geologic as well as topographic boundaries. Geomagnetic data are not available for this region to determine whether local anomalous heat sources are a possibility (Reference Sweeney, Finn, Blankenship, Bell and BehrendtSweeney and others, 1999). Because of this, we are not able to rule out the possibility that these syncline-layer boundaries result from a geothermal anomaly. However, this mechanism gives no explanation for the deep line diffractors seen in adjacent terrains. It is possible that the diffractors and the synclines are unrelated, but we search for a hypothesis that can simply explain all observations.

Our preferred hypothesis is that the melting occurred near a grounding line when the region was once floating. Grounding-line melt occurs at the back of a sub-ice-shelf cavity through the circulation of high-salinity shelf-water (HSSW) which is produced when salt is rejected during the formation of sea ice (Reference MacAyeal and JacobsMacAyeal, 1985). HSSW is delivered along the seabed to the grounding line and can remain up to 0.5 °C warmer than the in situ freezing point of water beneath the ice-shelf base (Reference GillGill, 1973). HSSW can induce basal melting, but strong stratification often prevents its contact with the ice-shelf base. Reference MacAyealMacAyeal (1984) shows that large tidal currents, developed in areas with a small water column thickness (such as at grounding lines), can stir the water column and promote localized basal melting. The pattern of grounding-line melt is not clear; some modelling predicts that melt would be spread over wide (~ 100 km) areas downstream of the grounding line (Reference Jenkins and DoakeJenkins and Doake, 1991), although limited observational evidence indicates that it could be more focused (>10 km) near the grounding line (Reference SmithSmith, 1996).

1. Ridge terrain

Inter-ice-stream ridges (e.g. Siple Dome and Engelhardt Ridge in Fig. 7) rise several hundred meters above the surrounding ice streams, but have a smooth surface topography over scales of hundreds of meters (Fig. 1). Surface velocities are slow, <8 m a⁻¹ (Reference Whillans and Van der VeenWhillans and Van der Veen, 1993), and radar profiles generally reveal continuous internal layers with undulations that appear as smoothed versions of the bed topography (Reference Nereson, Raymond, Jacobel and WaddingtonNereson and others, 2000).

Fig. 7. Lineations and terrains of lower KIS and WIS mapped from Figure 1. Flat-ice terrains are light grey, inter-ice stream ridges are dark grey, and ice-stream terrains are medium-grey. The modern grounding line is a thick dotted line, and the hypothesized relict grounding lines are thin dotted lines. Active ice-stream margins are shown as thick black lines while relict margins are thin black lines.

2. Ice-stream terrain

Reference RoseRose (1979) deduced that Ice Stream C (here called KIS) had stagnated, based on a lack of visible surface crevasses compared to other ice streams. Satellite images confirm this observation but also show that the surface topography of KIS is complex and more undulating than the surrounding interice-stream ridges and flat-ice terrains (Fig. 1). Typical slopes of ice-stream terrain are ~ 10⁻³ and although KIS is now slow-moving (0.1–10 m a⁻¹), numerous flow stripes still visible on the surface and buried marginal crevasses provide evidence of fast flow in the past. Typical internal structure within ice-stream terrain includes near-surface diffractors and deeper layers that are often continuous but highly distorted due to the large cumulative strain. Relict and active ice-stream terrains are indicated in Figure 7.

3. Flat-ice terrain

Flat-ice terrain is a distinctive, previously undefined, terrain type that lies between the inter-ice-stream ridges and the main body of relict KIS. Surface slope (10⁻⁴) and velocities (<0.3 m a⁻¹) are small (Reference Jacobel, Scambos, Nereson and RaymondJacobel and others, 2000) and the surface topography across these regions is smoother than ice-stream terrain but rougher than ridge terrain (Fig. 1).

Radar-detected stratigraphy across these areas shows layers that are continuous and lightly disturbed in the upper half of the ice, but discontinuous, with numerous diffractors, in the lower half of the ice (Fig. 6). These diffractors extend throughout the flat-ice terrains from their boundary with the inter-ice-stream ridges at the location of L4 and L8 to several kilometers (up to 10 km in Goosefoot) into the relict shear margins of KIS. The diffractors occur in distinct sets with uniform heights. In the Duckfoot there is one set ~ 150 m above the bed (Fig. 4), while in the Goosefoot there are two sets of diffractors: one ~ 200 m above the bed and another ~ 250 m above the bed (Figs 6 and 5). The transition between the two sets of diffractors in the Goosefoot occurs at L7.

Although the pattern of layer distortion above these diffractors appears to be random in the Duckfoot, it is partly organized in the Goosefoot. Figure 6 shows that the upper layers in the Goosefoot undulate with a wavelength of ~ 1 km; peaks in the layers roughly align with the tops of diffractors below them. Other profiles in the Goosefoot (Fig. 5) show flat layers overlying diffractors, implying that this organized pattern may not be widespread.

Grid and star patterns of radar data from across deep diffractors in the Goosefoot and Duckfoot reveal more information about diffractor geometry and orientation. Radar profiles across diffractors in directions perpendicular to L4 and L8 show diffraction hyperbolas, and migration of radar data along the C-line profile (Fig. 6) reduces these hyperbolas to point sources. In contrast, radar profiles in directions parallel to L4 and L8 show a line diffractor at the same depth as the top of the hyperbolas. This pattern is indicative of a source of contrasting permittivity along a line oriented parallel to L4 and L8.

We are not certain of the origin of the deep diffractors in these terrains. We rule out the possibility that they are buried surface crevasses because at such depths crevasses would be closed, reducing the permittivity contrast necessary to produce a diffractor. Reference Clarke, Liu, Lord and BentleyClarke and others (2000) found similar deep diffractors in the small inter-ice-stream ridge between the two limbs of Whillans Ice Stream (the Unicorn). They proposed several possible causes for the diffractors including the presence of basal crevasses, englacial morainal debris and a residual zone of temperate ice produced within the high-strain-rate environment of an active shear margin. Unlike the Clarke diffractors, diffractors in these flat-ice terrains occur up to 20 km away from relict shear margins and so we discount that they originated as temperate ice zones within shear margins.

Debris-rich ice can be accreted to the base of the ice column as seen beneath parts of KIS (personal communication from H. Engelhardt, 2001). Reference Kamb, Alley and BindschadlerKamb (2001) calculated freeze-on at the UpC camp of 0.0045 m a⁻¹ from borehole data, and rates of up to 0.1 m a⁻¹ have been reported beneath Matanuska Glacier, Alaska, USA (Reference Lawson, Strasser, Evenson, Alley, Larson and ArconeLawson and others, 1998). Even using this high rate, accretion of 250–300 m of basal ice (the height of the point diffractors above the bed) would take at least 2500 years. Although we do not rule out basal freeze-on, we suspect that other processes dominate.

Similar diffractor features have been detected in airborneradar profiles over ice shelves (Reference Jezek, Bentley and CloughJezek and others, 1979; Reference Jezek and BentleyJezek and Bentley, 1983). These are usually interpreted to be basal crevasses formed in floating ice. Tops of basal crevasses are likely to produce a strong reflection, especially if pockets of brine are trapped within them prior to crevasse closure. for a free-floating ice shelf of thickness H, Reference WeertmanWeertman (1980) calculated that a single bottom crevasse could penetrate up to ¶H/4. In the case of multiple basal crevasses, penetration height is expected to be lower (~ H/2) because the stress causing crevasse formation is spread over multiple crevasse tips (Reference Van der VeenVan der Veen, 1998). The remarkably constant height of these diffractors above the bed across large regions of these terrains might suggest that large areas were susceptible to a similar stress environment. The height of basal crevasses can be increased through tidal flexing of the ice shelf. Diffractor orientation supports the idea that, if these features are basal crevasses, they formed from increased tensile stress as ice flowed from the ridges into floating and/or faster-moving ice.

If these diffractors were caused by basal crevasses, an implication is that these terrains were once floating, which is consistent with the hypothesis that the syncline features described above formed through basal melting at relict grounding lines (marking the outermost boundary of these terrains). However, the ice in these terrains is presently grounded (Fig. 2b); in order for these regions to be floating they would have to have been 50–100 m thinner or sea level must have been much higher than at present (Fig. 2).