Variations in ice thickness and englacial stratigraphy were measured using a custom high-frequency impulse ice-penetrating radar system . Radar data were collected primarily using antennas with a center frequency of 2 MHz (∼80m wavelength in ice), but multiple lines were also collected with 7 MHz antennas (∼20 m wavelength in ice). The radar survey (Fig. ) was designed to sample both the along- and across-flow topography as well as features of interest identified in satellite imagery of surface morphology and velocity. Radar data were acquired at a constant distance interval of ∼15 m using an odometer wheel. Relative positions and orientations between radar traces still vary slightly due to traversing over surface features with variable wavelengths, such as windblown snow or sastrugi.

Radar waveforms were stacked 300–500 times per trace, with the number of stacks determined by driving speeds (time taken to traverse 15 m). Where available, dual-frequency GPS data were used to determine the positions and surface elevations of radar traces. If dual-frequency GPS data were not collected contemporaneously with the radar data, horizontal radar trace positions were interpolated linearly between end points of the profiles where dual-frequency GPS positions were recorded. Surface elevations along radar profiles were interpolated from the Reference Elevation Model of Antarctica REMA; if dual-frequency GPS data were not collected for that profile. We do not correct for surface elevation change as the integrated effect of these signals are small in the area of the survey compared to the range resolution of the impulse radar system. We confine our surface elevation interpretations from radar data to individual profiles to account for the multi-year temporal offset between the time of collection of GPS data and optical stereo-imagery used to construct REMA.

Radar data were processed following the workflow of with some modifications. First, a time correction for antenna spacing was applied to account for transmitter-receiver separation. A bandpass filter between 1–5 or 4–9 MHz was then applied to the 2 and 7 MHz data, respectively. Following geolocation, shorter, spatially adjacent or overlapping profile segments were concatenated to facilitate migration and radar interpretation. Along-track time-wavenumber migration was then applied to more accurately image subsurface reflection positions and slopes. Migration and conversion of travel time to depth assumed an average radar wave speed in ice of 168 m µs-1.

The corrected bed-returned power and depth-averaged and depth-variable englacial attenuation rates for individual traces were calculated for multiple reflectors in each trace (see Appendix  for additional information). Following and , we used Fourier spectral analysis to quantify along-track bed topographic roughness (see Appendix ). These methods assume that the reflection coefficients of internal layers are spatially continuous and that the relative reflectivity between layers is constant for the same transmit pulse settings (these settings include amplitude radiation pattern, transmitted power, and refractive focusing). We chose these power correction methods because they do not depend on assumptions for internal layer or bed depth variability or bed reflection power. As explained in Appendix A, the reflectivities results appear inconclusive, and we therefore do not include these findings in the subsequent analysis.

Due to the absence of continuous layers in englacial radiostratigraphy recorded at Conway Ice Ridge that can be traced and dated to an ice core site, we calculate a depth-age relationship to determine the age of the englacial layers we interpret using analytic models of ice sheet vertical velocity (see Appendix  for additional information). We choose to apply two relatively simple depth-age models due to spatiotemporal uncertainty in input parameters (accumulation rate, basal-melt rate, and ice-flow parameters) and vary these parameters over physically plausible ranges to estimate the age of englacial layers. Accumulation rates across the ridge used in these calculations are shown in Fig. . These models were chosen with the goal of determining the possible range of ages for our deep and shallow bounding layers, rather than providing a precise depth-age scale for this site.

We used velocity data from two sources: in-situ dual-frequency GPS data acquired contemporaneously with the ice-penetrating radar data and satellite SAR imagery. Repeat dual-frequency GPS measurements, acquired at 47 stake locations in November 2001 and November 2003, were used to determine local ice velocity over those two years. Stake positions were determined both seasons using dual-frequency GPS differential carrier phase positioning in a local (east, north, up) coordinate system relative to the base station . Velocity vectors were calculated by transforming local coordinates to geodetic coordinates and then calculating component differences between acquisitions and dividing by the time elapsed between measurements. We also use three SAR velocity mosaics to examine regional ice-velocity changes . These velocity fields were differenced to determine changes in velocity over the last 20 years (Fig. ). Speed change is consistent between products and in the figures and discussion we show the difference in speed between the and .

The Ice, Cloud and land Elevation Satellite-2 (ICESat-2) has now acquired over five years of surface height data at unprecedented spatial resolution and accuracy . We use the recently released Advanced Topographic Laser (ATL) altimeter land-ice products ATL14 (reference elevation) and ATL15 (elevation change relative to reference elevation) (version 1, released 15 December 2021, updated 27 May 2022) to contextualize the changes we observe in surface velocity with changes in elevation observed across Conway Ice Ridge and neighboring van der Veen and Mercer Ice Streams from October 2018 to July 2021 . We interpolate the reference digital elevation model provided in ATL14 onto the coarser grid for the ATL15 height change time series to produce a time series of elevation change relative to the ATL14 reference digital elevation model. Regional subglacial lakes were masked out of the interpolated elevation change products used in our modeling experiments using elevation change in the area of lake documented by .

To understand the mechanisms responsible for evidence of past flow changes interpreted in radar, surface elevation, and surface velocity we use a free-boundary diagnostic ice-flow model . Our method allows us to solve for both the position of past shear margins and past ice velocity near Conway Ice Ridge associated with past changes in ice thickness, boundary velocity, and basal strength that we prescribe to understand the system's sensitivity to changes in flow velocity and surface elevation. We use this model to understand how different configurations of ice-stream thickness, boundary velocity and basal strength affect flow around Conway Ice Ridge and evaluate mechanisms that can explain our interpretations of the radar data. This diagnostic approach allows us to solve for ice-flow speeds without prescribing accumulation and ocean melt-rate forcing, allowing us to evaluate whether different configurations of ice thickness are consistent with evidence of past fast flow we observe with radar data.

The model developed in is a thermomechanical ice-flow model based on the free-boundary formulation of ice-stream flow by based on shelfy-stream approximations proposed by . We solve a depth-averaged free boundary formation of ice flow over a pure plastic bed using Glen's law rheology with n=3 (see Appendix  for detailed discussion of model assumptions). 12∂∂xηH2∂ux∂x+∂uy∂y+∂∂yηH∂ux∂y+∂uy∂x-ρigH∂zs∂x-τcux|u|=0,2∂∂yηH2∂uy∂y+∂ux∂x+∂∂xηH∂ux∂y+∂uy∂x-ρigH∂zs∂y-τcuy|u|=0, where ui is the i component of velocity, ρi is the density of ice, g is gravitational acceleration, H is ice thickness, zs is the surface elevation and τc is basal strength, |⋅| is the L2 vector norm and η is the nonlinear viscosity of ice, defined in Eq. (2) of . We include thermal softening by coupling the 1D thermal model of to solve for the average ice softness at every grid point. Importantly, this approach neglects internal deformation from vertical shear in the ice column. This approximation is reasonable within ice streams and our model can be used to solve for the location of shear margins, but this approximation does not hold well within the ice ridge. Though we show results for the entire domain, model results within the ridge (speeds under ∼30 m yr⁻¹) are likely under estimates of ice speed.

Our model domain is shown in Fig. , and discussed further in Appendix . We select our domain such that the lateral boundaries are defined along modern flowlines in observed surface velocities of the van der Veen and Mercer Ice Streams, and the upstream and downstream boundaries are drawn >10 ice thickness away from our region of interest. Modern surface velocities were used as boundary conditions along the ice-stream boundaries of the model, upstream and downstream boundaries are stress-free, and modern-day surface and bed elevations define ice thickness throughout the domain . We interpolated the basal strength field from estimates of the basal resistance inferred using the Ice-sheet and Sea-level System Model ISSM;. This basal strength field biases our results towards modern flow configurations, so we intentionally limit our investigation to shear margin migration that is possible without significant reorganization of the basal strength distribution.

Using the model framework described above, we conduct three sets of diagnostic model experiments for Conway Ice Ridge in order to understand the system sensitivity to thinning, acceleration of van der Veen and Mercer Ice Streams, and the basal resistance field as a function of the ice thickness change.

For the thinning experiments, we used the surface elevation change observations derived from ALT14 and ALT15, as discussed in Sect. , to linearly extrapolate from modern thickness observations to configurations when the ice streams were on average ∼15 m thinner in increments of ∼3 m. We also include one simulation in this experiment where we extrapolate forward in time and prescribe 3 m of thickening.

For the ice-stream velocity experiment, we change the ice-stream boundary velocity conditions relative to modern velocities . We adjust the ice-stream velocities in increments of 10 % from 90 % to 150 % of modern velocities. These perturbations are only enforced at the along-flow boundaries of the domain, allowing the surface velocity structure to change over interior elements in response to changes in overall force balance.

Finally, we evaluate the sensitivity of ice flow around Conway Ice Ridge to assumptions for basal strength, accounting for influences on pore pressure and sliding at the ice-bed interface as the grounding zone advances from the ridge. Much of the bed of Conway Ice Ridge is situated below sea level, and thus, parts of Conway Ice Ridge and the surrounding ice streams are within 100 m of the flotation thickness. When the basal hydrology system is well connected to the ocean, changes in effective pressure on basal sediments vary with the height above flotation, which can change significantly as we force changes in ice thickness. For till-covered grounding zones, basal strength varies linearly with changes in effective pressure . We thus run two cases of basal strength dependence on ice thickness: one where basal strength remains independent of ice thickness changes and one where we adjust the basal strength linearly with changes in height above flotation 2HAF=H+ρwρizb3τc=τISSMHAFH0AF where HAF is height above flotation, H is ice thickness, zb is bed elevation, ρw,ρi are the densities of water and ice, respectively, τISSM is the basal strength inferred using ISSM, and H0AF is the currently observed height above flotation. A summary of the model runs can be found in table 1. More information about the model as it was applied in this study is included in Appendix D.

We connect our model results back to observations by interpreting two main features: the onset of fast ice motion where surface velocities exceed 30myr-1, which we treat as a proxy for shear-margin location, and locations where surface stresses from horizontal strain are high enough to potentially initiate crevasses. We identify the stress conditions that can initiate surface crevasses using a von Mises stress threshold of 200kPa, the regional estimate for the Siple Coast by , and well within the observed estimates of near-surface tensile strength determined from satellite imagery 90–320 kPa . To calculate modeled von Mises stress, we compute the modeled effective strain from our modeled velocities and use Glen's rheology to solve for the von Mises stress, assuming the entire column is pure ice: 4σv=3A-1/nϵE˙1/n where n=3. Importantly, we use a rate factor A=1.2×10-25s-1Pa-3 corresponding to -20 °C ice, which is approximately the mean annual surface temperature for this region. Our stress estimation does not account for internal heating or frictional heating at the bed, but since we are concerned with near-surface crevassing, it is appropriate to use this cold temperature for near-surface ice in our estimation of potential locations where surface crevassing may occur. Using these two model outputs, shear-margin location and the von Mises stress, we evaluate whether different configurations of the ice thickness, ice-stream velocity, and basal strength are consistent with our interpretations of the radar data.

Using the twenty years of satellite observations of surface velocity and elevation in this area, we map recent changes in velocity and ice thickness of Conway Ice Ridge. We calculated ice thickness changes from differences in vertical velocity of the surface from repeat GPS measurements following . Vertical velocities from repeat GNSS and accumulation stake measurements over the two-year period indicate that the average rate of thickening for the ridge is +1 cm yr⁻¹. Our result is consistent with previous calculations of the flux divergence using ice thickness measurements and the horizontal surface velocity fields obtained from satellite data that suggest the ridge centers are near balance . Three-year averages of gridded elevation change observations from ICESat-2 satellite laser altimetry indicate similar rates of thickening (Fig. ) and coincide spatially with changes in surface velocity (Fig. ). Where laser altimetry reveals ongoing thickening, the ice velocities have also slowed; where the ice-sheet thickness is unchanged, the flow velocities of the ridge are also unchanged over the last twenty years (Fig. ).

The distributed horizontal surface velocity and surface elevation fields also reveal the modern structure of Conway Ice Ridge that informed the designation of regions we use to present and discuss our radar and modeling results. Velocities derived from GPS and satellite imagery reveal that Conway Ice Ridge is currently bisected by two regions of faster flow (Fig. ). These small inter-ridge tributaries drain much of the ice from the interior of the ridge to Mercer Ice Stream and Whillans Ice Stream (Fig. ). Near the confluence of Mercer and Whillans Ice Streams, there is also a promontory where ice flows slowly from Mercer Ice Stream towards Whillans Ice Stream.

These three features (the inter-ridge regions of faster flow and the promontory where ice flows from Mercer towards Whillans) separate the ridge into four distinct regions that generally coincide with changes in the surface velocity and elevation of the ridge (Fig. ) and surface roughness revealed in SAR backscatter and passive microwave imagery . We refer to the promontory near the confluence of Whillans and Mercer Ice Streams as region I. Here, ice on the seaward tip of the ridge moves more slowly than the surrounding Mercer and Whillans Ice Streams and a tributary internal to the ridge that transports ice from Mercer to Whillans Ice Stream. On the northeastern ridge immediately adjacent to van der Veen Ice Stream, the surface is smooth in both SAR and MODIS imagery and qualitatively looks similar to the currently inactive Siple Ice Stream Fig. ;. We refer to this region as region II. To the west of region II is a small internal tributary that separates region II from the southwestern ridge, which we refer to as region III. In this region, near the confluence of Mercer and Whillans Ice Stream, the surface is generally smooth, but is interrupted by several depressions that run orthogonal to the surface elevation gradient. South of this region is the inter-ridge tributary that flows into Mercer Ice Stream that separates region III from the southern ridge, which we refer to as region IV. Here, the surface of the the ridge is rough, and average flow velocity is higher than any of the other slow flowing regions. These four regions, the promontory near the confluence of Whillans and Mercer Ice Streams (region I), the northeastern ridge (region II), the western ridge (region III), and the southwestern ridge (region IV), are marked in Fig. .

Ice thickness varies from 800–1200 m on Conway Ice Ridge (Fig. b). The thickest ice is on the northeastern side of the ridge in region II, towards the currently active shear margin of van der Veen Ice Stream, where we image half of a trough that is ∼300 m deeper than the surrounding basal topography (km 60 to km 70 in Fig. ). The current van der Veen shear-margin position does not coincide with the boundaries of this trough. The interior edge of this trough sits ∼10 km towards the ridge center from the current shear margin.

The interior tributaries where ice flows rapidly between region II and region III and between region III and region IV are underlain by wide (∼20 km) basal troughs that are ∼300 deeper than the surrounding topography (km 30 to km 50 in Fig. ). These troughs become shallower as the ice flows downstream into van der Veen and Mercer Ice Streams (Fig. ). Interior ridge tributaries appear to be unique to Conway Ice Ridge compared to other inter-stream ice-ridges along the Siple coast where the bed topography is smooth outside of the regions of fast streaming flow . Ice surface velocities are faster in these troughs. In the margins of these tributaries and across the promontory, we observe changes in the shorter wavelength bed roughness. In particular, along the profile near the confluence of Mercer and van der Veen Ice Streams (Fig. ), the transition from slow (<5 m yr⁻¹) to faster (>5 m yr⁻¹) flow speeds are co-located with rougher bed topography (Fig. a–c). These rough shorter wavelength features of the bed are also associated with visible depressions in SAR and MODIS surface imagery (Figs. , 5).

In profiles that cross slow flowing regions of Conway Ice Ridge (regions III and IV), bed conformal layers are generally present through the entire ice column (e.g. km 0–30 of the profile in Fig. c and km 0–30 of Fig. f). However, in locations where there is modern fast flow (>5 m yr⁻¹), for instance within the troughs that bisect the ridge (i.e. km 30–50 of profile in Fig. c and km 35–45 in Fig. f), observed layers are not bed or surface conformal. The abrupt change in layering with depth suggests a change in the strain history of the ice. There are two prominent observations of disrupted stratigraphy where ice flow is currently slow (<5 m yr⁻¹), indicating that stratigraphy may document past, more rapid ice-flow conditions in these locations.

The first observation of disrupted layers occurs on the eastern side of region II, where ice flows from Conway Ice Ridge into van der Veen Ice Stream shown in Fig. 6c km 50–72, Fig. f km 50–65. Here, smooth surface conformal stratigraphy is present in the top, younger portion of the ice column, but layers are deformed in the lower, older portion of the column. Following , we determined the timing of the changes in stratigraphy by tracking the deepest undisturbed layer above the disrupted stratigraphy. We then date the disturbed layer using a model for the vertical ice advection through the column and the apparent depth of the feature Fig.  and Appendix E;.

Our results indicate that layers deposited in the last 1000 years are continuous and surface conformal, whereas layers deposited between ∼1000 and ∼3250 years before present are disrupted so that they are not surface or bed conformal but are generally traceable, and deeper layers older than ∼3250 year before present are only continuous for short intervals and are not easily traceable. These disturbed layers in the upstream profile (km 50–65 in Fig. e–g) are consistent with past streaming flow and suggest the van der Veen margin migrated northward (narrowing). In the more downstream profiles, the same signal of outward migration of the margin (km 50–65 in Fig. b–d) suggests the ice stream also narrowed downstream; however, the disturbed layering ∼10 km closest to the present day margin is shallower suggesting it slowed later (∼1500 to 950 years before present). The shallower, and thus younger, disturbances closer to the margin in downstream profiles indicate that the stagnation and shear-margin migration started from upstream and that the van der Veen shear margin migrated northward 3250 years before present by 36 km prior to migrating another ∼10 km to the current position at a slower rate ∼1000 years before present.

The second observation of disrupted layering is in region I and in the interior of the fast-flowing tributary separating regions II and III. In these locations, hyperbolic reflectors at depth are consistent with the presence of buried crevasses Fig. c–e;. At the northwestern tip of the promontory (region I), stratigraphy is disturbed below 100 m depth and smooth and surface conformal above 100 m depth (younger than ∼750 years before present) on the seaward 20 km of the profile (Fig. ). More than 20 km inland from the confluence tip in region I, stratigraphy is smooth and bed conformal throughout the full ice column. The strong direct radar wave arrival prevents interpretation of stratigraphy in the upper ∼100 m of the ice column, so our ability to comment on changes in the recent past (∼400 years) is limited in the interior regions of the ridge.

In the tributary between region II and region III crevasses are imaged at 200–300 m depths (Fig. ). Using continuous layers traced above these hyperbolic reflectors, the presence of buried crevasses across the survey indicates that the surface stresses were previously sufficiently high to promote surface fracture . In the area of the tributary between regions II and III, depth-age calculations (see Appendix C) suggest that the buried crevasses would have been exposed at the surface ∼3000 years before present or earlier in this region. The abrupt transition between continuous, surface conformal layers at shallower depths (<300 m) and buried crevasse features suggest that ice flow slowed rather abruptly ∼3000 years ago. The smooth stratigraphy to the shallowest observable depths suggests that flow has remained relatively slow through the tributary to the present. At the tip of the ridge at the confluence of the ice streams, buried crevasses are present at ∼200 m depth below the surface for the first 10–15 km inland from the confluence tip (Fig. ). These depths correspond to an age of ∼1000–750 years before present, suggesting that ice slowed later here.

To better constrain the physical processes that might have led to the ice-dynamic changes suggested by our radar data, we use an ice-flow model to understand how the velocity field at Conway Ice Ridge responds to changes in ice surface elevation, basal strength, and the velocity of the neighboring van der Veen and Mercer Ice Streams. We use this model to identify how ice thickness change, ice-stream velocity change, and changes in basal strength could produce past flow conditions consistent with those inferred from internal stratigraphy. Specifically, we compare our model results against observations of disturbed layering implying historic shear-margin location, and buried crevasses implying historically elevated surface stress conditions.

We show the spatial changes in ice thickness in Figs. 3, derived from ICESat-2 ATL14/15 measurements used to test three different velocity responses to changes in ice thickness and boundary flow conditions. In these cases, we prescribe (i) thinning, (ii) thinning with 1.5× speed up of Mercer and van der Veen Ice Streams, and (iii) thinning, speed up, and weakening of the basal strength field depending on the height above flotation. The surface elevation fields produce the ice speeds shown in Fig.  (see Fig.  for all model experiments). The black dashed line represents the 30myr-1 contour, and the gray dashed line is the 30myr-1 contour of the present day case plotted for comparison. Figure e–g shows the modeled surface stress, with a green dotted line and the abrupt change from blue to green in color scale corresponds to 200kPa, the von Mises threshold we associated with surface crevasse formation.

We find that across most of the domain the shear margin defined by the 30 m yr⁻¹ contour remains stable in response to changes in thinning, with the notable exception of the promontory region. Here, our simulations show significant migration of the Mercer and van der Veen Shear Margins in response to ∼6 m of prescribed thinning (Fig. ). In these cases where we consider only changes in ice thickness, we do not detect notable change in the location of the 200kPa surface stress contour, which is our proxy for likely crevasse formation. This is true even for the case when we prescribe 15 m of thinning (Figs. b–d, ).

To consider the impact of changing regional ice velocities on Conway Ice Ridge, we alter our boundary conditions to mimic a change in the ice-stream velocities that bound the ice ridge. We find that changes to ice-stream velocity do not impact shear-margin location as much as changes in surface height, but do impact the surface shear stresses near shear margins and across the promontory, as shown in Fig. . Though speed up of the ice streams does increase surface stresses, stresses across the promontory remain below 200kPa our proxy for likely crevasse formation (Fig. f).

Finally, to consider the impact of changing basal strength with changing overburden pressure, we consider the scenarios above, but with basal strength now varying as a function of changes in height above flotation (HAF). This basal strength scenario considers the case of a perfect connection between the subglacial network and the ocean, resulting in a linear relation between basal strength and height above flotation. Compared to the cases where we hold the basal strength constant, the HAF dependent basal strength scenarios exhibit more migration at the promontory shear margin, as well as higher surface stresses at the promontory shear margin for all the prescribed thinning cases with increased ice velocity (Figs. g, ). The cases of 130 %–150 % velocity are the only cases that surpass the 200kPa von Mises stress threshold for crevasses formation in the region of observed buried crevasses.