Dilatant till facilitates ice-stream flow in northeast Greenland

Earth and Planetary Science Letters 401 (2014) 57–69

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

Dilatant till facilitates ice-stream flow in northeast Greenland Knut Christianson a,b,∗ , Leo E. Peters c , Richard B. Alley c , Sridhar Anandakrishnan c , Robert W. Jacobel b , Kiya L. Riverman c , Atsuhiro Muto c , Benjamin A. Keisling d,b a

Courant Institute of Mathematical Sciences, New York University, New York, NY 10012, USA Physics Department, St. Olaf College, Northfield, MN 55057, USA c Department of Geosciences and Earth and Environmental Systems Institute, Pennsylvania State University, University Park, PA 16802, USA d Department of Geosciences, University of Massachusetts–Amherst, Amherst, MA 01003, USA b

a r t i c l e

i n f o

Article history: Received 17 February 2014 Received in revised form 29 May 2014 Accepted 31 May 2014 Available online xxxx Editor: P. Shearer Keywords: Greenland geophysics glaciology ice streams active-source seismology ice-penetrating radar

a b s t r a c t We present radio-echo sounding (RES), global positioning system (GPS), and active-source seismic data across the central portion of the Northeast Greenland Ice Stream (NEGIS). NEGIS widens downglacier from a small region of high geothermal flux near the ice divide. Our data reveal high-porosity (40+%) water-saturated till lubricating the ice stream. Ice accelerates and thins as it flows into NEGIS, producing marginal troughs in surface topography. These troughs create steep gradients in the subglacial hydropotential that generate parallel “sticky” and “slippery” bands beneath the shear margins. The lowporosity “sticky” sediment bands limit ice entrainment across the margins and thus restrict further widening, producing the long, narrow, and relatively stable ice stream. However, the observed relations among surface elevation, basal water routing, broad sedimentary drape, and till dilatancy suggest that rapid shifts in ice dynamics are possible, including rapid transmission of ocean forcing inland. The source and routing of the subglacial till are unclear, but our data help constrain hypotheses. © 2014 Elsevier B.V. All rights reserved.

1. Introduction Greenland’s ice loss has quadrupled in the past two decades and continues to increase at an accelerating pace (Rignot and Kanagaratnam, 2006; Rignot et al., 2011; Joughin et al., 2012a; Straneo and Heimbach, 2013). This increased ice loss is correlated with the arrival of warm subsurface Irminger Water (IW) or Atlantic Water (AW) in Greenland outlet fjords (Holland et al., 2008; Joughin et al., 2012a; Straneo and Heimbach, 2013). IW or AW is already present at depth at the grounding lines of most Greenland glaciers (Holland et al., 2008; Straneo et al., 2012; Straneo and Heimbach, 2013); in the south the warmth of this water is causing rapid ice loss, but in the north this water cools during transit across the continental shelf, limiting its ability to melt glacier fronts (Straneo et al., 2012). Projected warming of this water by ∼2.0 ◦ C in the next 100 years (Straneo et al., 2012; Yin et al., 2011) is likely to cause sustained circum-Greenlandic ice loss, especially from those marine-terminating glaciers with large inland catchments, such as Jakobshavn Isbræ, Petermann Glacier, and the Northeast Greenland Ice Stream (NEGIS). Of these,

*

Corresponding author at: Courant Institute of Mathematical Sciences, New York University, New York, NY 10012, USA. E-mail address: [email protected] (K. Christianson). http://dx.doi.org/10.1016/j.epsl.2014.05.060 0012-821X/© 2014 Elsevier B.V. All rights reserved.

NEGIS may be especially important because streaming flow penetrates much farther inland, extending 700 km from the coast to nearly the ice-sheet summit dome (Fahnestock et al., 1993; Joughin et al., 2001). Here we investigate controls on the streaming flow of NEGIS and assess the potential for changes in ice dynamics in northeast Greenland. Ice streams and outlet glaciers are the principal paths for rapid ice discharge from ice-sheet interiors to their margins. In most parts of the Greenland Ice Sheet (GrIS), flow from broad inland catchments funnels directly into large outlet glaciers whose position is confined by deep bed troughs, with very little flow directly across the well-developed shear margins (Joughin et al., 2001, 2010). Deformation in deep, warm ice plays an important role in facilitating fast-flow (Phillips et al., 2010), although soft beds may locally contribute (Joughin et al., 2012b). Broad central regions of the ice sheet are frozen to the bed (Oswald and Gogineni, 2008, 2012), and together with bed sills above sea level and widespread crystalline bedrock (Henriksen et al., 2000), this is expected to limit rapid ice discharge to the ocean, with central regions of the ice sheet thinning only by diffusive processes (Pfeffer et al., 2008; Joughin et al., 2012b; Bamber et al., 2013). However, NEGIS differs greatly from this general Greenlandic model, as streaming flow continues to nearly the ice divide, providing a possible mechanism to rapidly transmit ice-marginal forcing far inland.

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Fig. 1. Maps of a) location of NEGIS and ice-core sites in Greenland (black box denotes area shown in (b)–(d)), b) surface ice velocity, c) surface elevation (contour (black) interval is 100 m), and d) bed elevation. White-dashed box denotes area shown in Figs. 2, 5–6. Black-dashed contours (25 m a−1 interval) in (b)–(d) are surface ice speed. Surface elevation is from Howat et al. (2014). Bed elevation is from Bamber et al. (2013). Ice velocity is from Joughin et al. (2010). Background is MODIS imagery (Haran, pers. comm., 2012).

High geothermal flux in the onset region of NEGIS produces basal water that initiates streaming flow (Fahnestock et al., 2001a). This likely causes the unique lack of a well-developed tributary system to NEGIS; rather than entering through the head of a tributary as in other ice streams, almost all ice entering NEGIS must flow across one of its pronounced shear margins, which persist far into the interior of GrIS (Fahnestock et al., 1993, 2001b; Joughin et al., 2001). Based on shear-margin characteristics and measured flow patterns, Fahnestock et al. (2001b) argued that the ice-stream margins support a significant fraction of the driving stress inland, with efficient lubrication of the central ice stream traceable from the geothermal source (Fahnestock et al., 2001a). Yet, despite occasional local topographic focusing, there is no dis-

tinct, persistent channel under the ice stream (Joughin et al., 2001; Bamber et al., 2013), as is observed for other Greenlandic ice streams (Truffer and Echelmeyer, 2003), suggesting that other mechanisms limit the extent of streaming flow. The downstream outlets of NEGIS sit in overdeepened fjord troughs, similar to other Greenlandic ice streams (Fig. 1d). Although NEGIS is subject to ocean forcing via the proximity of its outlets to AW formation zones in the Greenland Sea, a large bed sill ∼200 km inland of the grounding line (Joughin et al., 2001; Bamber et al., 2013) likely limits retreat via the marine ice-sheet instability (Fig. 1d). However, if warming ocean waters were to drive retreat, thinning would likely propagate up the ice stream beyond the sill, affecting subglacial water flow and also the extent

K. Christianson et al. / Earth and Planetary Science Letters 401 (2014) 57–69

of dilatant till and streaming ice flow, and thus ice-stream mass balance far inland. Hence, among GrIS ice streams, NEGIS may be especially susceptible to oceanic influences. To further investigate these processes, we conducted groundbased radio-echo sounding (RES), active-source seismic, and global positioning system (GPS) surveys in summer 2012, which provide the most direct observations of the basal conditions of NEGIS to date. We note that this fieldwork was part of a coordinated international effort to investigate NEGIS. Here we focus on the influence of the basal processes on ice-stream dynamics. Additional studies concentrate on reconstructing recent climate history from a shallow ice core (Vallelonga et al., 2014) and on examining ice dynamics across northeast Greenland using radar-derived internal stratigraphy (Keisling et al., 2014). Although streaming flow is confined to the central portion of our survey, our data indicate that the entire survey area is underlain by an unconsolidated sediment layer that is at least several meters thick. Combined interpretation of RES and seismic data reveal that water-saturated, dilatant till is confined to the central trunk of NEGIS, which is limited by subglacial hydropotential barriers, suggesting control by water sourcing and routing. However, the broader sedimentary drape across this interior region suggests the potential for rapid shifts in ice dynamics if these hydropotential barriers are overcome. 2. Field measurements and methods 2.1. Field site Our ground-based survey is located ∼140 km from the ice divide (Fig. 1), far enough toward the coast to characterize the ice stream rather than its origin, but far enough upstream to avoid crevassing. The shear margins step outward within our survey grid, as the ice stream flows up the basal topography (Fig. 1d). Ice-speed differences and strain rates across the shear margins in this area are similar to those near the onset of West Antarctic ice streams (Peters et al., 2006; MacGregor et al., 2013). We collected ∼350 line-kilometers of RES data, GPS data, five wide-angle seismic reflection profiles, and five seismic shallow refraction profiles, sampling the central ice stream, shear margins, and areas of slowmoving ice flanking the ice stream. 2.2. RES and GPS data RES data were acquired using a mono-pulse system operating at a center frequency of ∼3 MHz (Welch and Jacobel, 2003; Welch et al., 2009; Christianson et al., 2012). The radar system has separate sleds for a transmitter and receiver, and was towed by a snowmobile at approximately ∼10 kph. 2000 waveforms were stacked per recorded trace. Processing included bandpass filtering (fifth-order Butterworth filter from 1–5 MHz), time correction for antenna spacing, interpolation to standard trace spacing (8 m), two-dimensional time-wavenumber migration (assuming a constant radar-wave speed in ice of 169 m/μs), time correction for spatially variable firn density (Section 3.2.1), and correction for spherical divergence and englacial attenuation (see Appendix A). Using a more-complex velocity model in migration would not retrieve additional information, as essentially all of the radar-wave velocity variation occurs in the upper portion of the ice column obscured by ringing due to the air wave and the direct arrival. Processed RES traces have a vertical precision of ±3 m. Navigational data were collected using dual-frequency GPS receivers, and processed using differential carrier-phase positioning (Chen, 1998); conservatively, horizontal and vertical uncertainties are ±5 cm and ±10 cm, respectively. Surface elevation (Fig. 2), bed elevation (Fig. 5), internal layer topography (Fig. 4), and glaciostatic hydropotential (Fig. 6; Shreve, 1972) were derived from the RES

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and GPS data (see Section 3.2.1), interpolated to a 10-m grid using the nearest-neighbor algorithm with a search radius of 2.5 km, and then smoothed using a two-dimensional Gaussian filter with a diameter of 2 km. These parameters were chosen to eliminate obvious gridding artifacts. We do not use mass-conservation gridding (Morlighem et al., 2011, 2013, 2014) due to poorly-constrained, locally highly variable accumulation and basal-melt rates (Keisling et al., 2014), which would probably result in dubious apparent massbalance assumptions, and large errors in flow direction (Morlighem et al., 2014). We note that because our quantitative interpretations are restricted to length scales of a few hundred meters or more, they are relatively insensitive to gridding procedure. Projection of data in all figures is polar stereographic relative to the WGS84 ellipsoid with its central meridian at 45◦ W and standard parallel at 70◦ N. 2.3. Wide-angle seismic reflection data Five wide-angle seismic reflection profiles were collected across NEGIS (Fig. 2) for seismic amplitude variation with offset (AVO) analysis of the ice-bottom reflection (Aki and Richards, 2002; Peters et al., 2007, 2008; Peters, 2009). At each seismic site (labeled A–E in Fig. 2), a walk-away experiment was performed, using 48 vertical-component geophones (28 Hz center frequency) at a 20-m spacing to record energy from one-kilogram explosive charges placed 15–25 m below the surface. Sources and geophones were positioned to sample the same 480-m section of the bed over continuous source-to-receiver offsets of 0 to 4800 m (0–∼45◦ incidence angle), to better constrain the elastic properties (compressional-wave (P-wave) velocity, shear-wave (S-wave) velocity, and density) of the basal material. Minimal bandpass, dip, and FK-filtering were applied to the raw data to remove surface wave interference while preserving the true phase and amplitude of the ice-bottom reflection. Amplitude and uncertainty analysis follows Peters et al. (2007) and Peters (2009) (see Appendix B). 2.4. Seismic shallow refraction surveys Five seismic shallow refraction surveys were conducted across NEGIS to characterize the structure of the firn for accurate imaging and amplitude analysis. The locations of these surveys are approximately coincident with the five wide-angle seismic reflection sites (labeled A–E in Fig. 2). The profiles were collected using 48 vertical-component geophones at a 20-m spacing. For each shot, a 250 g explosive charge was placed ∼30 cm below the surface, with detonations at 5 m, 10 m, 15 m, and 20 m from the first geophone, providing continuous seismic sampling at a 5-m spacing from 5 m to 960 m source-to-receiver offset. This dense sourceto-receiver offset spacing was employed to model detailed seismic velocity–depth profiles of the firn and upper ice column, following the approach of Kirchner and Bentley (1990). We then applied the seismic velocity–density relationship of Kohnen (1972) to derive density–depth profiles through the firn at each site (Fig. 3). 3. Results 3.1. Surface topography In our survey area, we find a broad surface depression (∼10 m amplitude and ∼20 km wavelength) associated with the area of streaming ice flow (Figs. 2, 4a–c). This is a persistent feature along the entire length of the ice stream (Fig. 1c) (Fahnestock et al., 2001b; Joughin et al., 2001; Howat et al., 2014). Although elevation generally decreases from the ice-sheet summit towards the coast, there are numerous localized along-flow elevation increases within the ice stream (Figs. 1c, 2). These elevation anomalies are

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Fig. 2. Surface elevation in vicinity of ground-based NEGIS survey (white box in Fig. 1) from kinematic GPS data. Profile I–II is shown in Fig. 4. The five seismic sites where both an analysis of the firn column and amplitude analysis of the ice-bottom reflection were conducted are labeled A–E. Black contours are surface elevation (5 m interval). Dashed-white lines mark shear margins (areas with longitudinal strain rate greater than 0.002 a−1 ). Elevation outside of the ground-based grid is from Howat et al. (2014). Background is MODIS imagery (Haran, pers. comm., 2012).

likely due to ice flow over local bedrock highs or “sticky spots” (Joughin et al., 2001). The resulting topography is modified by trapping of drift snow, as shown by the accumulation variability in a shallow ice core retrieved near the center of the ice stream (site C in Figs. 2, 5), roughly 7 km downstream of the high point of one such elevation anomaly (Vallelonga et al., 2014). There are also well-defined, narrower surface troughs along both shear margins (Figs. 1c, 2, 4c) (Fahnestock et al., 2001b; Joughin et al., 2001). These troughs are ∼10 m deeper than the central portion of the ice stream and ∼20 m deeper than the flanking slow-moving ice (Figs. 2, 4c). Because of the absence of a well-defined tributary system, streaming flow in NEGIS must be supplied by ice that passes through a shear margin. Lateral drag on the ice is reduced rapidly after it flows through the shear margins due to downstream widening (Fig. 1). This results in a strong velocity increase as the ice flows through the shear margin, causing vertically compressive strain to balance longitudinal extension. The resulting surface-elevation low affects the basal hydropotential, which feeds back on the character of the surface troughs through its effect on basal lubrication. However, because surface topography can affect local snow accumulation rate, and both accumulation rate and strain rate can affect firn densification, careful consideration of firn density is required to properly calculate hydropotential. 3.2. Shear margins 3.2.1. Enhanced firn densification Based on surface strain rates, Fahnestock et al. (2001b) calculated that the marginal troughs would be more than 200 m deep; from this, they inferred that the accumulation rate in the troughs is enhanced to values of 1 m a−1 or more, versus an average accumulation rate of ∼0.11 m a−1 over the last 400 years in the center of the ice stream (determined from a firn core;

Vallelonga et al., 2014), due to trapping of drift snow in the troughs. Higher accumulation favors lower-density firn by burying that firn more rapidly. However, as summarized by Alley and Bentley (1988), anomalously high stresses from ice flow accelerate densification of high-density firn. At low density (≤550 kg m−3 ), densification is dominated by linear-viscous boundary sliding (Alley, 1987), which is not affected by ice-flow stresses. At higher densities (>550 kg m−3 ), densification is primarily by power-law creep, which is accelerated by ice-flow stresses. Physical understanding thus leads to the expectation that near-surface firn in the shearmargin troughs has anomalously low density, but that deeper firn might transition to higher density if the ice-flow stresses are sufficiently high. Our seismic shallow refraction data demonstrate that this occurs in our study region (Fig. 3). The density–depth profiles of the shear margin sites (B and D) possess lower densification rates at densities below ∼550 kg m−3 and transition to enhanced densification above ∼550 kg m−3 when compared to the non-shearmargin sites (A, C, and E), highlighting the importance of powerlaw-creep processes. We note that the seismically-derived density–depth profiles at site C overestimate the observed densities (Vallelonga et al., 2014) by up to ∼10% (Fig. 3). This may arise from the limited calibration data set (Kohnen, 1972), which is constrained by only four co-located seismic and density observations through the firn in Antarctica and Greenland. Another possibility is local heterogeneities exist within the firn, as the in situ firn densities of Vallelonga et al. (2014) and our seismically-derived densities for site C are separated by ∼2 km. Regardless, all seismic shallow refraction data were collected and processed in the same manner; thus observed variations in these seismically-derived densities reflect true site-to-site variability in firn densification, particularly for densities greater than 550 kg m3 .

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where ρd and ρi = 917 kg m−3 are the densities of the firn and ice, respectively, and i = 3.2 is the dielectric permittivity of ice. We neglect the imaginary part (  = σ /0 ω ; σ is conductivity; ω is frequency; 0 is permittivity of free space) of the dielectric constant ( =   − i   ) when calculating radar-wave speed in ice, because it is generally two orders of magnitude smaller than the real part (  ; permittivity) √ (Eisen et al., 2002). We then derive velocity profiles using v = c /   , where we assume that the medium is ice (i = 3.2) below 90 m, and then use these velocity profiles to convert travel time to ice thickness. Although this column of variable density firn is quite thin (∼70–90 m), due to the density variation (∼350–900 kg m−3 ), the effect on ice thickness calculations is significant.

Fig. 3. Seismically-derived firn density profiles from sites A–E (Figs. 2, 5–6). Dashed line at site C (green) denotes direct density measurement at the NEGIS shallow ice-core site. Note that firn density contrast between the shear margins (dotted lines) and other sites (solid lines). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Most calculations of ice depth from RES data assume a constant radar-wave speed in ice (typically 169 m μs−1 ), with either no firn correction or a firn correction taken to be constant beneath a sufficiently small area. However, Dutrieux et al. (2013) demonstrated that spatial variability in firn-air content on scales of only a few tens of kilometers occurred on the Pine Island Glacier ice shelf, likely arising from the complex ice dynamics there; such spatial variations in firn structure can alter the overburden pressure of the overlying firn/ice column by up to tens of meters. Our density data (Fig. 3) also indicate that using a constant regional firn correction may underestimate the radar-derived ice thickness in the NEGIS shear margins by as much as ∼10 m, with a substantial effect on hydropotential calculations. To account for the variable firn densification rate, we first assume that the firn densification is linearly related to the longitudinal strain rate. Following Van der Veen (1999), Price et al. (2002), and MacGregor et al. (2013), and using InSAR-derived velocities (Joughin et al., 2010), we calculate longitudinal strain rate (Fig. 4a) as:

ε˙ xx =

∂u ¯ u¯ s | · uˆ  = ∇| ∂x

(1)

where |u¯ s | is the surface ice speed and uˆ  is the surface ice velocity unit vector in the along-flow direction. We then linearly interpolate density along our radar profiles using longitudinal strain rates and our five measured density–depth profiles (Fig. 4b). From the density–depth profiles, we use a Looyenga dielectric mixing model (Looyenga, 1965; Macheret et al., 1993) to calculate the firn dielectric permittivity (d ) profiles, or:

d = 1 +

 ρd   1/3  −1 ρi i

(2)

3.2.2. RES internal layer folds Radar reflections within the ice are primarily isochrones from fallout of individual volcanic eruptions. In profiles orthogonal to flow, these layers show high-amplitude folds (several hundred meters) confined to both shear margins (Fig. 4f). In the southeast margin, these folds are clearly visible in all layers through the upper ∼1.5 km of ice, with only a weak correlation to underlying bed topography. The folds in the northwest margin are more complex, but generally persist through the entire ice column, have highest amplitude near the bed, and, in some places, coincide with undulations in bed topography. The greater fold amplitude downward, consistent behavior through the whole thickness imaged, and lack of tight correspondence to bed topography indicate that the folds arise primarily from variations in basal lubrication (e.g., Whillans and Johnsen, 1983; Christianson et al., 2013). Ice flowing into NEGIS across the shear margins experiences along-flow compression and vertical extension from increased basal shear stress in the poorly lubricated shear-margin bands; the ice then encounters well-lubricated streaming-ice regions where reduced basal shear stress causes along-flow extension and vertical thinning. This pattern is present along the entire shear-margin, likely persistent in both time and space (Keisling et al., 2014). Upglacier of our survey grid, the northwestern shear margin steps outward (Fig. 1) (Fahnestock et al., 2001b; Joughin et al., 2001). The additional folds in the ice stream near km 15–18 in Fig. 4f are approximately on the flowline from this step, and we suspect that the folds and the step share a common cause; modeling based on additional detailed surveys extending upglacier would be required to test this hypothesis, and to learn whether these folds record a fully steady flow history. RES internal and basal layer reflection strength may be affected by transmission losses or off-nadir reflections that are not recorded by the receiver (Holschuh et al., 2014). We see strong, consistent internal reflections everywhere in the survey area except in regions of steeply dipping layers in the shear margins, where both the internal layers and the bed echo appear anomalously weak (especially in the southeast shear margin), suggesting that internal losses are influencing the bed reflection strength. To test this hypothesis for basal reflection power (BRP) and internal reflection power (IRP), we follow Gades et al. (2000) and calculate the power returned within a time window t 2 − t 1 as:

Pr ≡

1 2(t 2 − t 1 + 1)

t2 

A n2

(3)

n=t 1

where A n is echo amplitude. For BRP, the time window is centered on the bed echo. To calculate IRP, we use a time window from t 1 = 1 μs to t 2 = tb − 1 μs (tb is the arrival time of the basal reflection) to avoid effects of the direct arrival and basal reflection. In Figs. 4d–e, after excluding migration edge effects, we can see that raw BRP and IRP decrease together only in the

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Fig. 4. Profile I–II in Figs. 2, 5–6. a) Ice speed (Joughin et al., 2010) and longitudinal strain rate, b) interpolated firn density based on longitudinal strain rate and seismic shallow refraction data, c) surface and bed elevation from kinematic GPS and RES data, d) internal reflection power, e) hydropotential and basal reflectivity (colors), and f) radargram. Ice flow is into the page.

shear margin, with the effect strongest in the southeastern margin; otherwise, IRP is relatively constant. We thus consider it most likely that features of the shear margin are reducing the power detected at the receiver. We calculate that strain heating is minor (1 ◦ C) and unlikely to significantly increase englacial attenuation (MacGregor et al., 2013), because flow through the margins is relatively rapid. Crystal-fabric orientation has a relatively weak effect on power returned in radar data (∼1–5 dB) at single frequencies in comparison to the large changes (up to 20 dB) we are investigating here (Eisen et al., 2007; Matsuoka et al., 2012); we thus discount fabric changes in the shear margins. The effects of roughness are similarly minor (<2 dB) at these frequencies (MacGregor et al., 2013). We consider it most likely that the low amplitude of the basal reflector in these zones is due to incoherent backscatter from the steeply dipping layers (Figs. 4d–f), which leaves less energy to propagate deeper into the ice, rather than an actual weakening of the electromagnetic contrast at the bed. We do not believe that we can make a quantitatively accurate correction for the lost power, so we mark these areas in Figs. 4 and 6, and interpret the reflection strength in other areas.

3.3. Bed characterization 3.3.1. RES bed characterization Unlike most ice streams (Bingham et al., 2012; Anandakrishnan et al., 1998; Bell et al., 1998; Peters et al., 2006; Winsborrow et al., 2010), the fastest ice flow of NEGIS is over a region of high and rough basal topography (Figs. 4–5). Surface topography along central regions of the ice stream does respond to basal topography, but the shear-margin surface depressions have much less basal topographic control. Although the northwest shear margin is loosely associated with a broad, shallow basal trough, we do not identify any controlling relationship. A narrow basal trough is present on the southeast side of the ice stream, but is oriented ∼15◦ off-axis from the current shear margin (Figs. 4–5). The ice stream location is more closely correlated with hydropotential (Fig. 6). Subglacial water flow is generally directed toward the coast along the ice stream, with some lateral spreading as the ice stream widens. Basal water routing along the shear margins is complex owing to influences of surface and bed topography, but the surface troughs generally overlie hydropotential lows. RES basal reflectivity varies across the ice stream (Figs. 4e, 6), but there are important discernible patterns. Scattering from internal layers complicates interpretation of basal reflectivity in the

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Fig. 5. Bed elevation in vicinity of ground-based NEGIS survey (white box in Fig. 1) from kinematic GPS and RES data. Profile I–II is shown in Fig. 4. Black contours are bed elevation (50 m interval). Dashed-white lines mark shear margins (areas with longitudinal strain rate greater than 0.002 a−1 ). Background bed elevation is from Bamber et al. (2013).

Fig. 6. Hydropotential and RES basal reflectivity (colored profiles) in the vicinity of the ground-based survey (white box in Fig. 1) from kinematic GPS and RES data. Profile I–II is shown in Fig. 4. White solid contours are hydropotential (5 kPa interval). Dashed-white lines mark shear margins (areas of longitudinal strain rate greater than 0.002 a−1 ). White box denotes area of high backscatter of steeply dipping internal layers.

shear margins. As noted in Section 3.2.2 and Figs. 4 and 6, areas of weak internal reflections are associated with anomalously low bed amplitude, likely because of power loss caused by steeply dipping layers. If so, the bed reflectivity in these areas is higher than indicated. Then, across most of the survey area we find stronger reflections where convergent flow and low hydropotential gradients indicate that there is accumulation of water, such as certain

portions of the shear margin complex and the central portion of the ice stream (Fig. 6). Our survey is in a region where Fahnestock et al. (2001a) and Keisling et al. (2014) estimated basal-melt rates of up to several centimeters per year. The consistently strong, bright radar reflections beneath the fast-flowing ice of NEGIS point to a thawed bed beneath the main trunk of NEGIS. The weaker radar reflections in

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Fig. 7. Cross plot of basal reflectivity and hydropotential. Black line is a second-order polynomial fit. Line becomes dashed when slope goes to zero, which indicates that hydropotential (i.e., water pooling) is no longer the primary control on reflectivity.

widespread regions outside the ice stream and in bands along the margins (Fig. 6) are consistent with a thawed bed possessing less free water (Peters et al., 2005), although we cannot entirely exclude the possibility that the bed outside the ice stream is locally frozen (Oswald and Gogineni, 2008). Within the ice stream, RES lines along ice flow may show slightly higher reflectivity than across-flow, consistent with flowinduced streamlining of the bed (Fig. 6); additional data are needed to confirm this trend. Reflectivity tends to be higher in regions of lower hydropotential gradient and near hydropotential minima, with lower reflectivity in regions of steeper gradient and across hydropotential maxima, consistent with the expectation that subglacial water accumulation increases reflectivity. In turn, this suggests that regions of low reflectivity are hydrologically controlled sticky spots. A cross plot of reflectivity and hydropotential supports this conclusion (Fig. 7); high reflectivity is strongly correlated with low hydropotential (i.e., water pooling). However, interpretation of hydropotential generally indicates the direction of water flow (i.e., down gradient), not the presence of water itself, although water will tend to collect in hydropotential minima. Thus, this relationship weakens or disappears at high hydropotential, perhaps because water consistently flows away from these areas and other factors, such as the nature of the basal materials, influence reflectivity; additional data and statistical tests would be necessary to test this further. However, a few deviations from this trend suggest additional complexity. Of particular interest is the banded reflectivity structure along the shear margins (Fig. 6). The steep hydropotential gradients driving water away from the center of the ice stream into the marginal troughs are regions of low reflectivity, with bands of high reflectivity in or just outboard of the trough axes. Interpreting high reflectivity as evidence of abundant water promoting lubrication by sliding or till deformation, these data suggest that the surface troughs are controlling hydropotential that in turn is controlling lubrication that feeds back to influence the surface troughs. The slight outboard offset of the highest-reflectivity zones from the estimated hydropotential minima can be explained by the estimated errors in our calculations from the effect of incoherent backscattering off steep internal layers, and from uncertainty in estimating firn-density anomalies. 3.3.2. Seismic AVO analysis Our seismic analysis of the subglacial environment was concentrated at 5 sites: two outside the ice stream (A and E), two

Fig. 8. Seismic amplitude variation with offset (AVO) analysis. a) AVO gather at site D. b) Modeled reflectivities for water (blue), dilatant till (green), dewatered till (red), consolidated sediments (black), and lithified sediments to crystalline basement (yellow) are given; the solid lines mark the bounds for each potential subglacial bed type. The circles, with error bars, give the calculated reflectivities of the ice-bottom seismic reflection from sites C (green), D (red), and E (black) (Figs. 2, 5–6). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

within the shear margins (B and D), and one in the trunk of NEGIS (C) (Figs. 2, 5–6). These sites show symmetric but spatially variable basal conditions (Fig. 8b; Table 1). A sedimentary drape is imaged at each seismic site, with a weak reflection from the base of the layer observed between the ice-bottom reflection and its “ghost” reflection (since the explosive charges are detonated 15–25 m below the surface, the “ghost” reflection arises from energy that first travels up to the surface before propagating down to the ice bottom and back to the surface, arriving ∼20–25 ms after the ice-bottom reflection in this instance (Fig. 8a)). This basal layer thickness varies from 11 ± 3 m in the central portion of the ice stream (site C) to between 6 ± 1 m and 9 ± 2 m at each of the four sites off the main trunk of NEGIS (A, B, D, and E). As summarized by Peters et al. (2007) (also see Blankenship et al. (1987) and Anandakrishnan (2003) for further details), ice overlying a weak, high-porosity (dilatant) till layer with water pressure very close to the ice overburden pressure (see characteristics for site C in Table 1 and Fig. 8b) produces a negative reflection coefficient at small offset angles, switching to a positive reflectivity at the larger incident angles we sample. This behavior is distinctive of saturated, high-porosity, presumably deforming sediment, and is observed in our data only at site C, within the main trunk

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Table 1 Modeled and best-fit elastic properties of the ice and potential subglacial bed scenarios for upper NEGIS. Bed lithology

P-wave velocity (m/s)

S-wave velocity (m/s)

Density (kg/m3 )

Ice Lithified sediment/crystalline basement Consolidated sediments Unconsolidated sediments/lodged till Dilatant till Water Site A Site Bd Site Ce Site Df Site Eg

3840a 3000–6200 2000–2600 1700–1900 1600–1800 1498c 1900 ± 200 1750 ± 150 1700 ± 250 1700 ± 200 2100 ± 200

1930b 1200–3400 1000–1400 900–1200 100–500 0 800 ± 100 900 ± 200 250 ± 150 600 ± 200 1100 ± 100

917c 2200–2800 1600–1900 1600–1800 1600–1800 1000c 2100 ± 200 1900 ± 200 1800 ± 100 1750 ± 300 2300 ± 200

a

Average P-wave velocity from sites A–E. Assuming a Poisson’s ratio of 0.33. c CRC Handbook of Chemistry and Physics (Lide, 1995). d We were unable to detect an ice-bottom multiple to definitely constrain absolute reflectivity. However, the negative polarity of the ice-bottom reflectivity at all incidence angles and the slightly negative slope of the reflectivity curve, suggest a till layer that is more compacted or “lodged” than at site D. e The site C reflectivity observations are shown in Fig. 8. f The site D reflectivity observations are shown in Fig. 8. g The site E reflectivity observations are shown in Fig. 8. b

of NEGIS. Following the approach of Blankenship et al. (1987) to constrain porosity n and effective pressure  P , we find best-fit values of n = 0.35 (range between n = 0.29 and n = 0.45) and  P < 100 kPa; combined with the bright RES reflectivity beneath the trunk of NEGIS, these two geophysical datasets point to a wet, dilatant till layer present where streaming ice flow is observed in our study region. The reflectivity curves for the remaining seismic sites point to more-competent, dewatered sediments outside of the main trunk of NEGIS (Fig. 8b; Table 1). When such a sediment layer is “lodged”, but not greatly compacted, the reflection coefficient is negative at all observed incidence angles, as seen for sites B and D, which are in the inboard parts of the shear margins in the bands of steep hydropotential gradient and low basal reflectivity. In comparison to site B, the lowered S-wave velocity at site D suggests higher porosity, water-saturated sediment, but less than at central site C. However, the weak nature of the sediment matrix at sites B and D gives rise to the possibility of dilating the sediment layer if sufficient basal water were rerouted to these marginal regions. With further compaction and dewatering, the reflection coefficient becomes positive for all incidence angles, as observed at sites A and E, indicating that a consolidated sedimentary layer exists outside of the fast-flowing ice of inland NEGIS. Best-fit curves to the data include enough uncertainty that we cannot completely exclude the possibility that the till at site C has lower porosity than the deforming tills of Whillans Ice Stream, West Antarctica (Blankenship et al., 1987), as Alley et al. (1987) suggest that a till porosity >30% is necessary for a till layer to become dilatant and potentially deformable. However, the distinct switch in the polarity of the basal reflectivity at site C is sufficiently diagnostic (Anandakrishnan, 2003) that we consider it highly likely that we are seeing a very soft, high-porosity, highwater-pressure, and thus likely actively deforming till beneath the main trunk of NEGIS, nondeforming till in the dewatered parts of the shear margins, and a more-compacted till drape outboard of the shear margin. Note that at shear-margin site B, we were unable to detect an ice-bottom multiple to definitely constrain absolute reflectivity. However, the negative polarity of the ice-bottom reflectivity at all incidence angles and the slightly negative slope of the reflectivity curve suggest a till layer that is more compacted or “lodged” than at site D (e.g., Dow et al., 2013). The seismic results presented are average conditions for a 480-m window of the subglacial environment at each of the five wide-angle sites. Given the peak frequency of the data analyzed

(120–150 Hz), the quarter-wavelength criterion of Ricker (1953) states that the subglacial bed must be at least 3 m thick for conclusive identification of a basal layer. We identified a seismic reflector arriving between the ice bottom and its ghost at each site, with thickness (calculated using our estimated velocities) greater than this quarter-wavelength criterion, and this is the layer we are discussing. To test the possibility of thin-layer effects on seismic amplitudes, we followed the approach of Booth et al. (2012) for each of the five seismic sites. A three-layer model was tested at each site, consisting of: i) ice, ii) water or dilatant till, and iii) lodged or consolidated sediment; the elastic properties for each of these potential layers are given in Table 1. For sites A, B, D, and E, a simple two-layer model (i.e., ice over meters-thick consolidated sediments for sites A and E; ice over meters-thick lodged till at sites B and D) provides the best fit to the data. Data from site C, however, allow both ice overlying 6 m of dilatant till, and ice over a thin layer (3 m) of water or dilated till over a layer of dilated till or somewhat more consolidated sediment. All permissible fits include material (water or dilated till) that would lubricate efficiently. The suggestion of vertical complexity is consistent with observations beneath Antarctic ice streams (Kamb, 2001). Additional data are likely to reveal more details of the basal conditions, and reduce the uncertainties here. 4. Discussion 4.1. Ice-dynamic feedbacks Unlike NEGIS, “normal” ice streams are constrained by topographic troughs or rift basins filled with soft sediments (Bingham et al., 2012; Anandakrishnan et al., 1998; Bell et al., 1998; Peters et al., 2006; Winsborrow et al., 2010). Generally, ice, lubricating water, and till are all funneled into the ice-stream channel from a broad inland catchment. Increase in velocity along-flow is relatively gradual, unless there is flow over a prominent bedrock headscarp, so there is no tendency for localized ice thinning producing a trough in basal hydropotential that would block water ingress. Little ice passes through the shear margins that develop along the sides of the ice-stream trunk, and this ice tends to feed a slow-moving boundary layer. Thus, this ice does not experience the strong stretching along-flow that would favor development of a surface trough affecting basal hydropotential and thus lubrication (Alley and Whillans, 1991; Anandakrishnan et al., 1998; Joughin et al., 2002; Rignot et al., 2002). Furthermore, even if such

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K. Christianson et al. / Earth and Planetary Science Letters 401 (2014) 57–69

Fig. 9. Cartoon schematics of a) effects of ice drawdown on subglacial till and water routing, and ice-stream extent, b) possible role of high geothermal flux and erosion of geothermally-altered rock in initiating streaming ice flow, and c) possible influence of past ice-sheet margin location on the onset of fast flow.

marginal troughs developed, inland flow could still enter the head of the ice stream without crossing the shear-margin trough. Hence, coastal thinning can directly and quickly draw down inland ice through most ice streams. Although the near-coastal part of NEGIS is similar to other ice streams, the features we imaged do not arise from inland extension of the coastal influence. Instead, our data, coupled with earlier studies (Fahnestock et al., 1993, 2001a, 2001b; Joughin et al., 2001), motivate the following extended hypothesis. NEGIS starts near the ice divide due to high geothermal flux there, from a feature with similarities to Yellowstone in size and heat flow (Fahnestock et al., 2001a). This region of hot bedrock produces meltwater that flows subglacially toward the coast down the hydropotential gradient, lubricating flow of the ice above. The ice stream widens downflow, likely due to lubricating water spreading as more is generated from the heat of sliding and flows around localized bed obstacles, which implies that ice-stream extent is uniquely limited by interactions between basal sliding speed and bed topography. Faster flow causes the ice stream to have a lower surface than the surrounding ice, which then tends to flow into the stream. But, the nearly point-source origin of the ice stream means that it lacks a well-developed tributary system, so all ice entering streaming flow must pass through a shear margin. Downstream widening reduces the lateral drag on ice as it moves from the shear

margin into the central trunk of NEGIS, allowing a strong velocity increase along-flow that favors surface thinning and thus creation of a local low in surface elevation. The resulting low in basal hydropotential focuses lubricating water, leaving bands of poorly lubricated bed as part of the shear-margin complex, and these poorly lubricated bands restrict ice and water inflow to the ice stream (Fig. 9a). Therefore, inland ice does not form a broad catchment tightly coupled to marginal changes. However, if there were large and rapid drawdown of the central part of the ice stream in response to coastal forcing, it would influence the shear-margin complex, and might even allow easier inflow and development of a broader catchment; testing of this hypothesis with modeling seems warranted (Fig. 9a). 4.2. Subglacial till The presence of till across our survey region helps explain the ice stream, as the extra water from the geothermal source favors high water pressures that allow dilation and till deformation without low-pressure Rothlisberger channels (e.g., Boulton, 1979; Walder and Fowler, 1994; Tulaczyk et al., 2000). Furthermore, because till acts as a fault gouge that limits erosion (e.g., Cuffey and Alley, 1996), it helps explain why this fast-flowing ice stream has not eroded a major subglacial channel.

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However, this does not explain the anomalous presence of till in north-central Greenland, which was generally thought to be underlain by crystalline bedrock (Henriksen et al., 2000). We suggest two hypotheses of processes that may have worked together, without excluding others. First, if the geothermal source is a Yellowstone-type caldera (Fahnestock et al., 2001a) or in other ways consists of geothermally altered rocks, it may be especially easily eroded to provide till (e.g., Geyer and Bindeman, 2011). Regions of high geothermal flux extend outside the head of the ice stream (Fahnestock et al., 2001a), and may have sourced till adjacent to the ice stream in our study area (Figs. 1, 9b). Another possibility is linked to the observation that large ice sheets often scour central regions and deposit extensive till sheets under marginal regions (e.g., Clark and Pollard, 1998). If the icesheet margin retreated sufficiently close to this region during peak past warmth, it may have favored deposition of till eroded from farther upglacier, and some of that till may still exist. If the ice sheet retreated entirely past our survey site, fluvial or aeolian deposition or soil formation and permafrost processes could have provided unconsolidated materials that allowed till formation following ice re-advance (Fig. 9c). Our data do not allow us to test these or other hypotheses; however, important insights would be gained from further mapping of geology and till extent in the upper regions of NEGIS. Downstream of our survey area, the boundaries of NEGIS generally correspond to a subtle basal trough (Fig. 1d), though it does not coincide exactly with the margins of streaming flow. Because till reduces ice-stream erosivity, even this subtle trough may be consistent with existence of a rift basin or other geological feature localizing erosion and thus ice-stream flow. We note, however, that the limited basal topographic expression of the ice stream also may indicate that it is a relatively young feature. 5. Conclusions To understand the importance of NEGIS to the past and future mass balance of GrIS, additional field and modeling studies are necessary. Here we suggest that downstream portions of NEGIS are sensitive to ocean forcing, and that upstream portions may be especially sensitive to subglacial water routing and feedbacks between ice thickness and subglacial erosion that influence hydropotential. Sufficiently large thinning along the ice stream from coastal forcing or general ice-sheet drawdown due to surface melt could remove the surface troughs and their effects, allowing NEGIS to tap a larger catchment area, especially if forcing can be rapidly transmitted from coastal to inland zones. Such thinning may already be occurring (Khan et al., 2014). This possibility, plus a host of questions such as the sources and transport of till, and whether the ice stream can have persisted over long times without eroding a trough or if it is in the process of eroding one, motivate additional observations and modeling, especially given the high likelihood of ocean warming at the terminus in the next century. Studies of crustal structure in northeast Greenland would be particularly revealing, as they may indicate sediment source and distribution, and would add greatly to our knowledge of the past and future behavior of GrIS, ice-stream evolution, and ice-sheet/crustal interactions. Acknowledgements The U.S. National Science Foundation (grant OPP-0424589) funded this work. UNAVCO provided GPS base station data. CH2MHILL Polar Services, the New York Air National Guard, Kenn Borek Air, the Alfred Wegener Institute for Polar and Marine Research, and the North Greenland Eemain Ice Drilling project provided logistical support. We thank Bernd Kulessa and an anonymous reviewer for comments that improved this

67

manuscript. GMT, QGIS, and the NWRC and Columbia colormaps were used to create figures (licensed under creative commons attribution share-alike 3.0 (http://creativecommons.org/licenses/ by-sa/3.0/deed.en) from sources http://commons.wikipedia.org/ wiki/File:North_Rhine-Westphalia_Topography_01.svg and http://fr. wikipedia.org/wiki/Fichier:Columbia_Mapa_Relieve.svg, respectively). Appendix A. RES basal reflectivity The radar returned power (echo intensity) P r is a function of radar system characteristics (S), geometric spreading (G), and medium and reflector characteristics (I ). In a decibel scale, this is written as:

[ P r ]dB = [ S ]dB + [ I ]dB − [G ]dB

(4)

where [ I ]dB = [ R ]dB − [ L ]dB − [ B ]dB is a function of the reflectivity of the target interface (R), integrated dielectric attenuation along the propagation path (L), and birefringence (B) (Bogorodsky et al., 1985; Matsuoka et al., 2010; Matsuoka, 2011). We neglect birefringence from crystal-fabric orientation alignment because of its small effect (∼1–5 dB) on P r relative to the large changes expected in R (Eisen et al., 2007; Matsuoka et al., 2012). For a ground√ based system, geometric spreading is simply [G ]dB = 2[ z/   ]dB  for a depth (z) and mean relative permittivity ( ). Thus, following Matsuoka (2011), the geometrically corrected returned power from gc the bed ( P bed ) is:



gc

P bed



dB

= [ P bed ]dB + [G bed ]dB ∝ [ R bed ]dB − [ L bed ]dB

where [ L bed ]dB = 2z N , N  =

(5)

(104 log10 e)σ √ is the depth-averaged  c 0



attenuation rate, σ is the conductivity of the medium, 0 is the permittivity of free space, and c is the speed of light in vacuum. Thus we can calculate N  from the slope of a linear fit c of [ P bed ]dB as a function of z over a region of varying ice thickness that have the same R (Jacobel et al., 2009). We apply this method to 3479 traces with especially bright echo intensities in the north-central portion of the ice stream. This yields an attenuation rate of N  = 8.0 ± 0.3 dB/km, which is in agreement with other radar studies of northeast Greenland (Oswald and Gogineni, 2008). Adding this correction yields the corrected returned power c from the bed ([ P bed ]dB = [ R ]dB + [ S ]dB ). As we do not have knowledge of relevant system parameters to calculate S, but can safely assume it is constant for this system (Christianson et al., 2012), c we scale [ P bed ]dB such that the central trunk of the ice stream has reflectivity consistent with an ice/groundwater till interface (−6 to −15 dB) and the brightest reflection is consistent with that of an ice/groundwater interface (−2 dB) (Peters et al., 2005; Anandakrishnan et al., 2007). Appendix B. Seismic AVO analysis The amplitude of a seismic reflection A R at a given offset x can be defined as (Aki and Richards, 2002):

A R (x) = A 0 G (x) R (x)e−ar (x) ,

(6)

where A 0 is the source amplitude, G (x) is the geometrical spreading correction, R (x) is the reflection coefficient, a is the seismic attenuation in ice, and r (x) is the length of the seismic ray path. Our goal here is to constrain R (x), as each R (x) curve is unique for a given set of elastic properties (P-wave velocity, S-wave velocity, density) (Fig. 8b; Table 1). We use the amplitude relationship between the ice-bottom reflection and its first multiple to constrain A 0 (Peters et al., 2008), and spectral analysis of the ice-bottom reflection to determine a through the ice column

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