PUBLICATIONS
Journal of Geophysical Research: Earth Surface
RESEARCH ARTICLE
10.1002/2014JF003178
Key Points:
• Cold-based debris-covered glaciers
preserve long-term records of
climate change
• Englacial debris bands reflect
climate oscillations
• Climate changes are observed as arcuate
ridges on debris-covered glaciers
Supporting Information:
• Readme
• Table S1
• Table S2
• Figure S1
• Figure S2
• Figure S3
• Figure S4
• Figure S5
• Figure S6
• Figure S7
• Figure S8
• Figure S9
• Figure S10
• Figure S11
• Figure S12
• Figure S13
Correspondence to:
S. L. Mackay,
smackay@bu.edu
Citation:
Mackay, S. L., D. R. Marchant, J. L. Lamp,
and J. W. Head (2014), Cold-based
debris-covered glaciers: Evaluating their
potential as climate archives through
studies of ground-penetrating radar
and surface morphology, J. Geophys.
Res. Earth Surf., 119, 2505–2540,
doi:10.1002/2014JF003178.
Received 14 APR 2014
Accepted 15 OCT 2014
Accepted article online 22 OCT 2014
Published online 28 NOV 2014
Cold-based debris-covered glaciers: Evaluating
their potential as climate archives through
studies of ground-penetrating radar
and surface morphology
Sean L. Mackay1, David R. Marchant1, Jennifer L. Lamp1, and James W. Head2
1
Department of Earth and Environment, Boston University, Boston, Massachusetts, USA, 2Department of Earth,
Environmental and Planetary Sciences, Brown University, Providence, Rhode Island, USA
Abstract
We describe the morphology and internal structure of Mullins and Friedman Glaciers, two
cold-based, debris-covered alpine glaciers that occur in neighboring valleys in the McMurdo Dry Valleys,
Antarctica. Both glaciers are overlain by a single, dry supraglacial debris layer (8–75 cm thick); each mantling
debris layer is marked with near-identical patterns of arcuate ridges and steps, as well as corresponding
changes in bulk grain size, meter-scale surface topography, and thermal contraction crack polygons. Results
from 24 km of ground-penetrating radar data show that the ice within the uppermost 1–2 km of Mullins and
Friedman Glaciers is essentially free of englacial debris (<1% by volume) but thereafter is interspersed
with bands of englacial debris (each <3 m thick and dipping up glacier) that intersect the ice surface at all
major surface ridges and steps. The similarity in number and pattern of englacial debris bands and
corresponding surface ridges and steps across both glaciers, along with model results and observations
that call for negligible basal entrainment, is best explained by episodic environmental change at valley
headwalls. Our working hypothesis is that layers of englacial debris originate as supraglacial lags that form
in ice-accumulation areas during times of reduced net ice accumulation; following renewed net ice
accumulation, these lags are subsequently buried by snow and ice, flow englacially, and intersect the ice
surface to impart distinctive changes in the texture of supraglacial debris and topographic relief. The
implication is that the englacial structure and surface morphology of these cold-based, debris-covered
glaciers preserves a consistent record of climate and environmental change.
1. Introduction
Debris-covered viscous flow features that are cored with interstitial or massive buried ice are conspicuous,
yet enigmatic landforms in glacial and periglacial environments. Current research aims to understand
their origin(s) [Clark et al., 1998; Humlum, 1996; Martin and Whalley, 1987; Potter et al., 1998], ice-flow
dynamics [Arenson et al., 2002; Haeberli et al., 2006, 1998; Kääb et al., 2003; Kääb and Weber, 2004],
erosive and debris-transport capabilities [Barsch and Jakob, 1998; Haeberli, 1985; Humlum, 2000], and
paleoclimate significance [Clark et al., 1996; Fitzpatrick et al., 1995; Frauenfelder and Kääb, 2000; Humlum,
1998; Kääb and Weber, 2004; Konrad et al., 1999; Steig et al., 1998]. In most cases, the layer of surface
debris that caps these features tends to reduce ablation of underlying ice [Konrad and Humphrey,
2000; Kowalewski et al., 2006, 2011; Lundstrom et al., 1993; Mihalcea et al., 2006; Potter, 1972] and
permits long-term survival of subsurface ice [Clark et al., 1998], possibly for more than 8 Ma
[Sugden et al., 1995b].
The specific origin of ice in many of these features is largely unknown and can vary over time. For example,
ice may originate as segregation ice, from the burial and compaction of snowbanks, from the densification
of snow and firn as for typical glaciers, or from other sources [Burger et al., 1999; Clark et al., 1998; Giardino
et al., 1987; Haeberli and Vonder Mühll, 1996; Johnson, 1984; Potter et al., 1998; Shroder et al., 2000]. Further,
the long duration over which many of these features seem to form [Clark et al., 1998, 1996; Fitzpatrick et al.,
1995; Potter et al., 1998] suggests that ice origins may evolve over time, yielding notable, and perhaps
cyclic changes in the style of debris entrainment, ice-flow dynamics, and overall morphology [Ackert, 1998;
Frauenfelder and Kääb, 2000; Haeberli et al., 1998; Haeberli and Vonder Mühll, 1996; Whalley and Azizi, 1994,
2003; Whalley and Martin, 1992].
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In an effort to elucidate the climatic significance of some long-lived viscous flow features in Antarctica, we
initiated in 1995 a suite of detailed geomorphological and geophysical studies of buried ice in interior regions
of the McMurdo Dry Valleys (MDV) [Kowalewski et al., 2012, 2006, 2011; Marchant and Head, 2007; Marchant
et al., 2002; Sugden et al., 1995b; Swanger and Marchant, 2007; Swanger et al., 2010]. Our initial results
demonstrated that buried ice can survive for millions of years [Marchant et al., 2002; Sugden et al., 1995b], that
sublimation and subsequent vapor diffusion through capping debris can be extremely slow on the order
of <101 mm a1 [Kowalewski et al., 2012, 2006, 2011; Marchant et al., 2002], and that the origin of buried
ice deposits varies with local microclimate conditions [Marchant and Head, 2007; Marchant et al., 2013;
Swanger et al., 2010]. Here we focus on the origin and internal ice-flow dynamics of two viscous flow features,
previously mapped as cold-based debris-covered alpine glaciers: the Mullins valley and Friedman valley
debris-covered glaciers (hereafter informally termed Mullins Glacier and Friedman Glacier). We selected
Mullins and Friedman Glaciers for detailed study because (1) initial studies suggest that they are cored by
glacier ice and thus represent the polar, end-member class of cold-based, debris-covered glaciers [Kowalewski
et al., 2011]; (2) their proximity allows for direct analyses and comparison of ice-flow dynamics, debris
entrainment, and landscape modification under shared climate conditions [Kowalewski et al., 2011; Shean
et al., 2007; Shean and Marchant, 2010]; (3) they have been cited as terrestrial analogs of similar-appearing
viscous flow features on Mars [Head et al., 2010; Levy et al., 2006; Marchant and Head, 2007; Marchant et al.,
2010]; and (4) preliminary geophysical surveys suggest that the uppermost portion of Mullins Glacier,
at least, contains a nonuniform distribution of englacial debris that may reflect environmental change
[Shean and Marchant, 2010].
2. Background
2.1. Setting
Debris-covered glaciers in the MDV are relatively uncommon and appear restricted to inland regions where ice
accumulates in the lee of steep cliffs, most notably beneath cliffs of Ferrar Dolerite. Mullins Glacier (77.89°S,
150.58°W, 1557 mean sea level (msl)) originates along rock cliffs of Ferrar Dolerite and Beacon Supergroup
sandstones at the headwall of Mullins Valley, ~2100 m elevation. It flows along the southwest side of Vestal
Ridge for ~3.5 km and then bends eastward into upper Beacon Valley, one of the major valleys that nearly bisect
the Quartermain Mountains (Figures 1 and 2a). After traveling ~8 km, Mullins Glacier merges imperceptibly at
1350 m elevation with stagnant, Miocene-aged buried glacier ice derived from an ancient, southward flowing
incursion of Taylor Glacier into central Beacon Valley [Kowalewski et al., 2011; Sugden et al., 1995b] (Figure 1).
(Unless otherwise noted, reported distances along the glacier centerline are measured from the
valley headwall).
Friedman Glacier (77.90°S, 160.51°W, 1540 msl), situated just to the west of Mullins glacier, likewise originates
along steep cliffs of Ferrar Dolerite and Beacon Supergroup sandstones, flows northward from ~2050 m
elevation alongside the westside of Rector Ridge (Figure 2b), and appears to terminate at a lobate front at
1500 m elevation and 3.8 km distant in upper Beacon Valley; buried ice from Friedman Glacier may, in fact,
extend beyond this point [e.g., Kowalewski et al., 2011], but if so it is unreachable in hand-dug excavations and
unrecognizable in ground-penetrating radar (GPR) data (see below). A small, unnamed debris-covered glacier
(referred to herein as “DCG_03”) flows for 0.5 km alongside uppermost Friedman Glacier.
Data from interferometric synthetic aperture radar (InSAR) show that both glaciers display peak horizontal surface
velocities near their respective valley headwalls of ~40 mm a1 but thereafter decelerate to near 0 mm a1
(within measurement error of ~ ±.002 m a1) at ~3.6 km (Friedman) and 4.8 km (Mullins) [Rignot et al., 2002].
2.2. Ice Accumulation and Ablation
Ice accumulation is derived from a combination of windblown snow transported off the East Antarctic Ice
Sheet (which is ultimately trapped in the lee of the steep cliffs of Ferrar Dolerite and Beacon sandstone that
nearly encircle both Mullins and Friedman valleys) and from direct precipitation along valley headwalls
(Figure 2). The relative contribution of each source is unknown, but the values have likely varied over time
[Marchant and Head, 2007]; modern-day regional snowfall is <50 mm water equivalent a1 [Fountain et al.,
2009]. A few isolated clasts of Ferrar dolerite, and to a lesser extent Beacon sandstone, dot the surface of
the ice-accumulation area during summer months; some of these clasts lie along the distal end of linear
MACKAY ET AL.
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Figure 1. Part of the Landsat Image Mosaic of Antarctica showing the southern portion the McMurdo Dry Valleys (MDV).
2
Centered at ~77°35′S, the MDV encompasses ~4000 km of mostly ice-free terrain between the East Antarctic Polar
Plateau (left) and the western Ross Sea (right). The white box in the lower left highlights the location of Mullins and
Friedman valleys, both tributaries to Beacon Valley in the Quartermain Mountains and depicts the geographic location for
Figures 3 and 4 and Figure S1 in the supporting information. The inset plots the location of the MDV (small blue box) on a
sketch map of the Antarctic continent.
Figure 2. (a) Oblique aerial view of Mullins valley (view to the south) showing the accumulation area and a portion of the
ablation area (Regions 1 and 2) of Mullins Glacier. (b) Oblique aerial view of Friedman valley (view also to the south)
showing the accumulation area and a portion of the ablation area (Regions 1–3) of Friedman Glacier. Bedrock cliffs are
Ferrar Dolerite (dark brown) and Beacon sandstones (light brown). (c) Large clast of Ferrar Dolerite (>35 cm in diameter)
within the uppermost ablation area of Mullins Glacier (Region 1). The clast is perched on an ice pedestal at the center of a
~15 cm deep moat-like depression. (d) Similar clast as observed in Figure 2c, but with small gravel-and-cobble-sized clasts
at the base of the central rock, and within the surrounding moat. (e) Dolerite clasts, <10 cm in diameter, embedded in ice
within the uppermost ablation area of Mullins Glacier (Region 1); the clasts are covered with a very thin (<1 cm) layer of
refrozen meltwater.
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Figure 3. Geomorphic sketch map showing the Mullins and Friedman Glaciers and surrounding regions. Both glaciers
grade from small ice-accumulation areas at valley headwalls (very light blue) and transition down valley to exposed
glacier ice (uppermost ablation area, Region 1, light blue) and finally to fully debris-covered ice in Regions 2–4 (various
shades of grey). DCG_3 highlights a portion of the undifferentiated debris-covered glacier in uppermost Friedman Valley.
Stippled areas of increased polygon maturation highlight abrupt local increases in trough depth fringing polygon centers
[see also Levy et al., 2006].
tracks in the snow, which can be traced upslope, tens-to-hundreds of meters, toward points of local
detachment along isolated rock outcrops.
The transition to net ice ablation on both glaciers is irregular and spatially complex. In general, the transition
appears to coincide with a subtle break in ice-surface slope at ~1600 m elevation. However, isolated blue ice
ablation areas occur up to 2000 m elevation on both glaciers and attest to complex ablation patterns, most
likely influenced by minor variations in solar radiation (shadowing, slope/aspect), and seasonal winds.
The ablation areas of Mullins and Friedman Glaciers both show abrupt transitions from exposed glacier ice
(e.g., ice dotted with scattered and isolated rocks, and thin patchy gravels) to uniform coverage and burial
beneath thin supraglacial debris (Figure 3). This transition coincides with the first major arcuate ridge on both
glaciers and lies at ~1.5 km for Mullins Glacier and ~1.7 km for Friedman Glacier. Thereafter, supraglacial
debris on both glaciers thickens down valley, increasing from ~7 cm to ~45 cm on Friedman glacier and from
~10 cm to a maximum of ~75 cm on Mullins Glacier (see section 4.4). As discussed in Kowalewski et al. [2011],
the supraglacial debris is most likely sourced from rockfall at valley heads; additional input from valley
sidewalls is unlikely because (1) large topographic depressions exist at the lateral margins of both Mullins
[Kowalewski et al., 2011] and Friedman Glaciers, and (2) valley sidewalls are typically dry and do not appear to
collect wind-blown snow in sufficient quantities to initiate detachment and rockfall (Figure 2) [Augustinus and
Selby, 1990; Kowalewski et al., 2011]. Additional input from wind-blown sediment is likely minor. We base
this assertion on the general paucity of wind-blown sediment in the Dry Valleys [e.g., Lancaster, 2002] and
on the absence of suitable ice-free source areas to the south and southwest of Mullins and Friedman Valleys
(e.g., in the direction of the strongest and most frequent winds in the region [Marchant et al., 2013]).
2.3. Cryoturbation
The cold and dry conditions limit active layer cryoturbation. However, for low-albedo dolerite rocks (albedo of
0.07 [Campbell and Claridge, 2006]), intense solar radiation may warm surfaces above 0°C, even though
atmospheric temperatures remain well below 0°C (see also section 4.1). Wherever these warmed rocks come
in direct contact with exposed glacier ice, as occurs in the uppermost ablation areas of both Mullins and
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Friedman Glaciers, sublimation increases and minor melting may ensue; this meltwater may flow down glacier,
in very small channels <2 mm deep, and ultimately refreeze in a thin layer of superposed ice just inward of
the first topographic ridge on both glaciers (Figure 3; e.g., the frozen pond of Shean and Marchant [2010]; see
also sections 4.4.1 and 4.4.2). Because the 0°C isotherm does not penetrate more than a few centimeters in
the exposed rocks and debris [Kowalewski et al., 2006, 2011], glacier ice buried beneath at least ~10 cm of debris
is perennially dry and remains frozen year round (e.g., Kowalewski et al. [2011]; see also section 4.1).
Given these environmental conditions, ablation beneath supraglacial debris ≥10 cm thick is accommodated
solely by sublimation. Calculated rates of ice sublimation beneath Mullins debris are on the order of 0.01 to
0.001 mm of ice loss a1 [Kowalewski et al., 2011] and soil gravimetric water content is ≤3% [Campbell and
Claridge, 2006]; taken together, these factors imply negligible saturated active-layer cryoturbation, e.g.,
negligible lateral and vertical displacement of rocks and debris from alternate freezing and thawing of
ice. This assertion is supported by (1) the finding of coherent signals in repeat synthetic aperture radar
interferometry, which suggest negligible independent movement of surface boulders (at least over the 3 year
measurement period, 1996–1999) [Rignot et al., 2002], and (2) the preservation of millimeter-scale
microstratigraphy and in situ volcanic ash fall [Marchant et al., 1996, 2002, 2013], which suggests insignificant
sediment sorting and involution over considerable time frames.
An important corollary that arises from these observations is that the surface morphology and grain-size
characteristics of supraglacial debris on both Mullins and Friedman Glaciers, as well as the concentration and
grain-size characteristics of shallow englacial debris (e.g., debris observed at and just below the buried ice
surface), reflect the intrinsic properties of the glaciers themselves, rather than postdepositional modification
and alteration via saturated, active layer cryoturbation (see also Kowalewski et al. [2011]). None of the clasts
examined on or within Mullins and Friedman Glaciers exhibit glacial striations, facets, or polish.
2.4. Sensitivity to Environmental Change
For a typical clean-ice alpine glacier, surface albedo decreases and atmospheric temperature increases downglacier, resulting in increased ablation towards the toe. Mullins and Friedman Glaciers adhere to this general
increasing ablation trend with distance down-glacier, but only in their uppermost portions. Once supraglacial
debris thickens beyond a critical threshold, ablation is reduced [e.g., Ostrem, 1959]. For Mullins and Friedman
Glaciers, ablation is highest in the uppermost portion of the glacier and then decreases dramatically beneath
a thickening cover of supraglacial debris in the down-ice direction (see also section 4.4; [Kowalewski et al.,
2011]). This spatial pattern of ablation is characteristic of many debris-covered glaciers [Ackert, 1998; Konrad
et al., 1999; Scherler et al., 2011]. For Mullins Glacier, the ablation rate decreases abruptly by over 2 orders of
magnitude from a measured rate of ~6 cm a1 for mostly exposed ice in the uppermost ablation area (see
section 4.3) to a rate of ~0.5 mm a1 across the first major topographic ridge (arcuate surface discontinuity
(ASD)-1, see below and Figure 1), where surface debris abruptly thickens from ≤2 cm to 7–10 cm [Kowalewski
et al., 2011]. The net effect is a distinct increase in sensitivity to climate change being focused on a relatively
small area of mostly exposed ice directly down valley from the present accumulation area [Konrad et al.,
1999]. Thus, any perturbation in environmental conditions that reduces net accumulation (ameliorating
atmospheric temperatures, wind patterns less favorable to snow deposition, and/or overall precipitation/ice
accumulation reduction) will result in disproportionate ice loss in the exposed uppermost ablation area,
yielding low-elevation, “spoon-shaped” ablation hollows that merge down valley with elevated remnants of
protected (debris-covered) glacier ice [e.g., Marchant et al., 2013]. The development of such spoon-shaped
hollows has been discussed for debris-covered ice flows outside Antarctica [Ackert, 1998; Johnson and
Lacasse, 1988; Potter, 1972; Whalley, 1979], and these features have been implicated as indicators of local
environmental change. Within the McMurdo Dry Valleys, the features are described as “beheaded debriscovered glaciers” [Marchant and Head, 2007], one of which occurs in association with remnant, debriscovered glacier ice along the western margin of upper Beacon Valley [e.g., Marchant et al., 2013, Figure 7].
3. Methods
3.1. Environmental Data Collection
3.1.1. Micrometeorological Automatic Weather Stations
Micrometeorological conditions were monitored for up to 6 years using automatic weather stations (AWS) at
seven locations on Mullins and Friedman Glaciers (see Figure S1 in the supporting information). The suite of
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environmental sensors (all Onset Computer Corporation “smart sensors”) varied among AWS (see Table 1)
but included a combination of (1) soil temperature and soil moisture sensors typically arrayed in vertical
profiles down to the buried ice surface; (2) downwelling shortwave solar radiation; (3) surface barometric
pressure; (4) air temperature and relative humidity (RH) at ground level (5 cm) and mast height (2 m); and (5)
mast height (2 m) wind speed and direction (for individual sensor specifications including range and
accuracy, see Kowalewski et al. [2006]). Each station sampled environmental conditions every 3 min and
logged averages to HOBO® microstation U21 data loggers [e.g., Kowalewski et al., 2006]. Temporal coverage
varies between 1 and 8 years, depending on installation date and occasional sensor failures (Table 1).
3.1.2. Borehole Temperature Monitoring
In November 2008, we recorded downhole temperatures for borehole MCI-08-01 (see section 3.3) located
~1.8 km from the headwall of Mullins Valley (77.89438 S, 160.58324°E) at 5, 10, 15, and 20 m depths. A second
thermistor string was deployed at borehole MCI-09-03 (4.6 km from Mullins headwall, 77.8787°S, 160.53614°E)
in December 2009 to monitor temperatures at 10 m, 20 m, and 30 m depths. Both arrays were deployed for a
minimum of 5 days duration.
3.1.3. Ablation-Stake Array
In 2010 and 2011 we installed an array of ablation stakes in the uppermost ablation areas of Mullins and Friedman
Glaciers: 14 stakes for Mullins and 5 stakes for Friedman (Figure S1). Each stake was constructed from 3.175 cm
diameter PVC pipe (capped and filled with snow) and installed via drilling a 3.175 cm diameter hole ~70 cm down
into glacier ice. The heights of the stakes above the glacier surface have been measured annually since, with
measurements referenced to the top of a board (15 cm × 20 cm) placed over the adjacent ice surface; the latter
minimizes errors associated with localized ice loss from minor wind scour immediately adjacent to stakes.
3.2. Geomorphic Mapping and Sedimentological Analyses
Morphometric analyses relied on field mapping and study of remotely sensed data. The latter included
analyses of high-resolution aerial photographs (U.S. Geological Survey (USGS) TMA3079V0296 and
TMA3080V0276, November 1993), multispectral and panchromatic satellite imagery (1.5 m and 0.41 m
resolution, respectively, GeoEye01, January 2004), and airborne LiDAR data (2 m resolution) [Schenk et al.,
2004]. We collected detailed topographic measurements and sediment samples at 75 sampling localities
along both glaciers; samples included supraglacial debris and subjacent glacier ice (Figure S1). Each ~2 kg
sample (approximately ~1 L) of supraglacial debris was sieved at 16 mm in the field; the >16 mm fraction was
analyzed for lithology, shape, and weathering characteristics [e.g., Swanger et al., 2011], and the smaller
fraction was stored in plastic Whirl-Pack® bags for standard grain-size analyses at Boston University [e.g.,
Kowalewski et al., 2011]. Unlike earlier sampling strategies that were designed to provide spatial coverage and
detect general trends in grain-size characteristics and polygon morphometry along Mullins and Friedman
Glaciers [e.g., Kowalewski et al., 2011; Levy et al., 2006], our strategy here focused on obtaining numerous,
closely spaced samples across mapped surface ridges and arcuate steps.
At the base of most soil excavations, we quarried 10–40 cm down into buried glacier ice to collect samples for
δ18O and δD analyses, and total dissolved solids (TDS). We recorded the visible concentration of englacial
debris, if present. Ice samples were double-bagged in Whirl-Pak® bags and kept frozen at 20°C until
analyzed. Isotopic analyses were conducted on an isotope ratio mass spectrometer at the Boston University
Stable Isotope Laboratory. Analyses for δ18O were accomplished via CO2 equilibration; deuterium analyses
were accomplished via pyrolysis using a GV Instruments Ltd. ChromeHD™ system. TDS were estimated from
electrical conductivity (EC) measurements.
3.3. Ice Core Drilling
We collected 21 shallow ice cores along the central flowline of Mullins glacier (Figure S1). Seven relatively
deep cores (15–30 m) were extracted using an electrically driven Koci Drill designed and manufactured by Ice
Core Drilling Services for use in both clean and dirty ice glaciers [Green et al., 2007]. Fourteen shallow cores,
each ≤ ~5 m, were obtained using Snow, Ice, and Permafrost Research Establishment and Kovacs Mark II hand
corers; core recovery at all sites was >95%. Drill sites were cleared of supraglacial debris (if present),
photographed, and georeferenced using high-resolution GPS. To protect the pristine environment, no
drilling fluids were used. Drill heads on the KOCI drill were changed from ice cutters to rock-drilling heads
whenever isolated rocks and/or thick (>10 cm wide) sand wedges were encountered [Green et al., 2007;
Kowalewski et al., 2011]. Maximum penetration using the KOCI drill was ~30 m (at MCI-09-03); the primary
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Ice surface temperature
h
Debris surface temperature
RH
Wind
Air temperature
Location
g
F_Met_01
M_Met_01
M_Surf_01
7.5 (21.9)
0.8 (16.8)
13.4 (26)
194
8.2 (21.8)
5.3 (19.8)
10.8 (24)
12.5
77.89445°S
160.58001°E
1543
2.4
2
NA
9
12/10–01/13
M_Surf_02
25
12/06–01/13
M_Met_02
50
12/10–01/13
M_Surf_04
77.8804°S
77.87588°S
160.53431°E
160.5379°E
1376
1342
4.2
4.8
3
4
10 (22)
9 (18)
13 (25)
3.1
44 (52)
12/063/11, 12/
NA
08–01/13
4 (5)
45 (47)
8.6 (22.5) 5.4 (21.0)
1.8(16.0)
2.0 (16.2)
14.2(26.4) 11.8 (25.6)
45
195
10.6 (22.9) 9.5 (21.0)
10 (20.0)
9.9(22.3)
9.2 (20.8) 10.1 (19.8)
11.4(23.6) 9.9 (21.6) 10.4 (19.8)
0
0
0
77.89154°S
160.57093°E
1523
2.8
2
11 (23)
7 (18)
15 (26)
0.8
48 (57)
NA
18
12/04–01/13
M_Surf_03
Mullins
The observation period for each station varied due to staggered initial installation dates and occasional data logger failure. The time periods listed here refer to the period for the entire station.
All statistical valuations account for only the total time the station was in service; no specific adjustments are made to incorporate the seasonality of any outages. The observation intervals for the
wind measurements are listed separately to reflect their higher volatility.
d
The total time the sensor recorded temperatures ≥0°C divided by the number of years in the observation record.
e
Anemometers sensors failed at a higher rate than all other sensors.
f
Because the anemometers failed during high-magnitude wind events, this value likely underestimates the true maximum velocity.
g
Measurements from 1 cm under surface of the debris layer.
h
Measurements at the base of the debris layer at the interface with massive ice (except for M_Met_01 which has no debris layer).
(in parentheses) observations. NA, not applicable.
Debris thickness (cm)
0
0
9
c
Observation period (MM/YY) 12/10–5/11, 12/ 11/09–1/10, 11/ 12/10–01/13
11–4/12
10–01/13
Latitude
77.91553°S
77.9096°S
77.90226°S
Longitude
160.52429°E
160.59789°E
160.58998°E
Elevation above sea level (m)
1613
1642
1605
Distance from headwall (km)
1.0
0.9
1.5
Ablation area region
1
1
2
Mean (°C)
14 (18)
14 (19)
Mean daily maximum (°C)
11 (15)
11 (16)
Mean daily minimum (°C)
17 (22)
16 (22)
-1 d
Time ≥ 0°C (h a )
0
0
Mean (%)
54 (52)
53 (48)
e
Wind data period
12/10–5/11, 12/ 11/09–1/10, 11/
NA
11–4/12
10–01/13
-1
3 (4)
5 (4)
Mean (m s )
-1 f
17 (26)
16 (23)
Maximum (m s )
Mean (°C)
NA
NA
8.7 (22.6)
Mean daily maximum (°C)
NA
NA
0.5 (17.6)
Mean daily minimum (°C)
NA
NA
14.4 (26.3)
-1 d
NA
NA
145
Time ≥ 0°C (h a )
Mean (°C)
12.2 (24.0) 9.9 (22.8)
Mean daily maximum (°C)
7.0 (20.1)
7.0 (20.8)
Mean daily minimum (°C)
15.6 (26.7) 12.4 (24.6)
-1 d
Time ≥ 0°C (h a )
0
17
b
Friedman
a
No data indicated by “-.“ Statistical values are listed for both summer-only (defined as D-J-F) and annual
b
Both M_Met_01 and M_Met_02 are located in Region 1 in areas devoid of any surface debris layer.
c
Surface
Atmosphere
Physical
a
Table 1. Micrometeorological Observations
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Figure 4. Location map for transects imaged with ground-penetrating radar (GPR). Antenna center frequencies are shown
by line color and pattern; dotted lines show transects collected using multiple antenna frequencies. Transects acquired
with 200 MHz and 400 MHz antennas were collected in distance mode. Transects using 80 MHz antennas were collected in
point mode unless otherwise noted. Due to the great length of the longitudinal transects Am–Am′ and Af–Af′ (5.1 km and
3.2 km, respectively), data collection occurred over a period of several weeks, and with different depth ranges configured.
The background is a hill-shaded relief map from airborne LiDAR (2 m resolution [Schenk et al., 2004]). Reference distances
down valley from the headwall datum (referred to throughout the text) are indicated in the grey boxes. DM, Distance Mode.
obstacle in drilling deeper was frictional heat (increasing the probability of drill entrapment downhole via the
refreezing of meltwater) generated by cutting through rock and ice-rich sediment. The core data were used
to help calibrate GPR velocity (see section 4.5) and ground-truth mapped englacial reflectors.
3.4. Ground-Penetrating Radar
We used a Geophysical Survey Systems, Inc. (GSSI) Subsurface Interface Radar 3000 ground-penetrating radar
(GPR) system at multiple antenna frequencies (80 MHz, 200 MHz, and 400 MHz) to collect over ~24 km of
common offset GPR data from Mullins and Friedman Glaciers (Figure 4).
Although radar investigations of ice-and-rock mixtures typically use low antenna center frequencies in the
range of 15–50 MHz [Berthling et al., 2000; Degenhardt, 2009; Degenhardt and Giardino, 2003; Hausmann et al.,
2007; Isaksen et al., 2000; Monnier et al., 2008], we follow Shean and Marchant [2010] and De Pascale et al.
[2008] in selecting higher-frequency antennas in the range of 80–400 MHz that are capable of achieving both
high spatial resolution and sufficient depth penetration. Due to the excellent transmission properties of the
relatively clean and cold glacier ice, the 200 MHz and, in specific locations, the 400 MHz were successful in
reaching depths of ~100 m. However, to achieve these depths, the 200 MHz and 400 MHz antennas required
close physical contact with the buried ice surface, and thus, considerable effort was required to remove
surface debris in order to expose buried glacier ice (or at least get to within a few centimeters of its surface).
By contrast, the 80 MHz antenna was capable of achieving penetration of 100+ m without the necessity of
removing surface debris. This capacity greatly enhanced our ability to maintain minimal environmental
disturbance, enabling preservation of the unique morphology and stratigraphy of surface debris along GPR
transects. Transects were collected along both longitudinal (down glacier) and transverse paths (Figure 4).
3.4.1. Survey Execution
These surveys were conducted in either distance mode, where a terrain-calibrated survey wheel triggered
individual scans as a function of distance, or point mode where several scans were stacked at each prescribed
MACKAY ET AL.
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distance step. All long (>400 m) transects were collected using the unshielded Multiple Low Frequency (MLF)
antenna in common offset mode configured for a center frequency of 80 MHz. Shorter GPR profiles collected
with the 200 MHz and 400 MHz antennas were conducted on naturally exposed glacier ice, over glacier ice
covered with snow, or where acceptable connection with the buried ice surface could be achieved with
minimal removal of supraglacial debris. For the latter, all clasts > ~20 cm were cleared along the transect line,
permitting close connection between the antenna base and the buried ice surface; typically a vertical
separation between antenna and ice of <10 cm was achieved. For additional details, see Table S1 in the
supporting information.
A Trimble 5700 Global Positioning System (GPS) receiver with a Zephyr geodetic antenna was used to collect
high-resolution GPS (12 min minimum residence time) at each survey end point and continuous topographic
GPS points along each survey line. After differential location correction utilizing data from a base station
located ~7 km away on University Peak (77.86259°S, 160.75983°W; 2195.3 m), all position and elevation data
displayed millimeter- to centimeter-scale accuracy. Points were imported into ArcGIS 10.0, and continuous
elevation profiles were extracted from an airborne LiDAR digital elevation model (DEM) for Beacon Valley
[Schenk et al., 2004]. Differences between the measured GPS and DEM elevations are minimal
(mean = 0.33 m, 1σ = 0.39 m).
3.4.2. Vertical and Horizontal Resolution
The theoretical vertical resolution of a radar signal is calculated using vertical resolution = λ/4 = (V/F)/4
[Reynolds, 1997], where λ is the center wavelength of the antenna (m), V is the radar wave velocity of the
medium (m s1), and F is the center frequency of the antenna (MHz). However, the actual vertical resolution
will likely be significantly lower due to the progressive filtering of higher-frequency elements from the
waveform [Hubbard and Glasser, 2005; Trabant, 1984], which may be particularly noticeable for debriscovered ice bodies [Berthling et al., 2000]. For this reason, we calculate a conservative estimate of vertical
resolution = λ/3 for each of the antennas, yielding 0.15 m, 0.3 m, 0.75 m, for the 400 MHz, 200 MHz, and
80 MHz antennas, respectively, using a velocity of 0.168 m ns1 (see section 4.5 for velocity determination).
We follow the simplified approach of Welch et al. [1998] and use λ/2 as an approximate minimum size for
general horizontal object resolution in migrated data. This approach results in working horizontal resolutions
of 0.2 m, 0.4 m, and 1.1 m for the 400 MHz, 200 MHz, and 80 MHz antennas, respectively.
3.4.3. Postprocessing
GSSI Radan 6.7 and Radan 7.0 softwares were used in postprocessing GPR data. Processing steps common to
almost all survey lines included the following: (1) horizontally appending individual data files comprising
multipart survey lines; (2) static position correction to bring air-wave return to the top of the scan time
window; (3) distance normalization using GPS and/or survey wheel data; (4) application of a finite infinite
response filter with a 50/110 MHz, 120/220 MHz, or 340/480 MHz band-pass boxcar filter for the 80 MHz,
200 MHz, and 400 MHz survey data, respectively; and (5) gain adjustment. Additional processing steps,
performed on a subset of survey lines, included (6) background removal, (7) migration, (8) Hilbert transform,
and (9) topographic surface correction. Background removal is a spatial filtering process to remove constant
artificial horizontal banding across the radargram that is caused by ringing between the source and receiver
antennas and which may mask subtle “real” radar features. All of the 80 MHz survey data displayed mild to
severe banding in the range 0– ~ 200 ns, possibly caused by the relatively close spacing of the source and
receiver antennas producing partial signal oversaturation [Hauck and Kneisel, 2008]. For these data, we
applied a background removal spatial filter of length 71 scans, over the range window 21 ns to 250 ns. We
then performed a simple two-dimensional Kirchoff migration assuming a mean velocity of 0.168 m ns1 (see
section 4.5) to all survey data. Although migration at this velocity produced excellent hyperbola collapse
of point reflectors in most profiles, we note the probable distortion of steeply dipping and complex
continuous reflectors that may arise from the migration of radar data acquired in topographically complex
environments [Lehmann and Green, 2000]. The large step size (0.7 m to 1.4 m) used in the 80 MHz surveys
(relative to the signal wavelength) further reduced the efficacy of applying traditional migration. For this
reason, data have been displayed in both migrated and unmigrated forms. For many of the migrated profiles,
we performed an additional Hilbert magnitude transform [e.g., Yilmaz, 2001] to isolate the energy envelopes
of individual scans [Arcone, 1996]. Finally, both the migrated and unmigrated data were topographically
corrected based on 2 m elevation data extracted from a LiDAR DEM [Schenk et al., 2004] and confirmed with
postprocessed kinematic GPS elevation data (section 3.4.3).
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4. Results: Field Data
4.1. Automatic Weather Stations
Measured atmospheric conditions within the
uppermost ablation areas of both Mullins and
Friedman Glaciers (Region 1; Figure 3) are nearly
identical and show persistent air temperatures well
below 0°C (average of 19°C) and low-average
values of relative humidity (RH) near ~53% (Table 1).
At the lowest elevation meteorological station
(Figure S1), M_Met_02, located 4.2 km down valley
from the headwall on Mullins Glacier (1376 m
elevation), near-surface air temperatures (measured
5 cm above the ground surface) exceeded the
melting point for an average of only 3.1 h each year.
Temperatures at the buried ice surface at this site
(at 25 cm depth) and elsewhere were consistently
<0°C. Complete meteorological data are provided
in Table 1.
Figure 5. Measured borehole temperatures, mean annual
atmospheric temperatures, and mean annual soil surface
temperatures at, and near, selected borehole locations. The
middle portion of the figure displays borehole temperatures;
all data for MCI-08-01 are plotted in grey; MCI-09-03 data in
black. Borehole data were collected between 17 November
2008 and 26 November 2008 (MCI-08-01) and between 1
December 2009 and 6 December 2009 (MCI-09-03). The
mean annual atmospheric temperature and mean annual
surface temperature of supraglacial debris (compiled from
data collected at nearby meteorological stations <50 m
away) are plotted in the upper portion of the image (The
representative meteorological data for MCI-08-01 come
from met station M_Surf_02; the representative data for
MCI-09-3 come from met station M_Met_02; see Figure S1
for locations). Stippled areas show the thickness of supraglacial debris at each representative met station.
4.2. Borehole Temperatures
Borehole temperatures show a general increase
with depth, reaching 23.2°C and 21.5°C near the
base of MCI-08-01 (at 20 m depth) and MCI-09-03 (at
25 m depth), respectively (Figure 5). The 15 m depth
borehole temperature within both temperature
strings is very close to the mean annual
atmospheric temperature at both core sites. This is
in agreement with the standard approximation of
10–15 m depth for the maximum penetration depth
of the seasonal thermal wave [Paterson, 1994; Wade,
1945] and for a thermal system dominated by
conduction only; it also supports the assumption
that there are no englacial heat sources or sinks
including melting/freezing or strain. The nonlinear
concave downward shape of both borehole profiles at the shallowest depths most likely reflects the lingering
influence of the seasonal wave at the time of sampling.
4.3. Ablation Measurements
Repeat measurements of the ablation-stake arrays on both glaciers show that ablation of exposed glacier ice
in the uppermost ablation areas on Mullins and Friedman Glaciers average 5.56 and 6.86 cm a1, respectively
(see Table S2 in the supporting information). Mullins stake 01, the only stake located on the secondary frozen
meltwater pond (Figures 2, 3, and S1) [Shean and Marchant, 2010] (see also section 4.4.1.3), recorded a net
increase of 2.5 cm between 2011 and 2013 (Table S2).
4.4. Ground Penetrating Radar, Ice Cores, and Surface Geomorphology: Organization of Data
Data from ground penetrating radar (GPR), shallow ice cores/ soil excavations, and surface morphology reveal
striking similarities across Mullins and Friedman Glaciers. These similarities, especially those observed in GPR,
provide an informal basis for partitioning the ablation areas of Mullins and Friedman Glaciers into four
distinct regions. Region 1 includes the uppermost ablation areas, whereas Regions 2–4 represent further
subdivisions across the remaining, fully debris-covered portions of each glacier (Table 2 and Figure 3). Below,
we present data from GPR transects, shallow ice cores and hand-dug ice excavations, and geomorphologic
analyses of surface features for each region. For consistency, we provide data for Mullins Glacier first, followed
by Friedman Glacier. Section 4.6 provides a short synthesis of key interpretations for all regions.
MACKAY ET AL.
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MACKAY ET AL.
Start elevation (m)
End elevation (m)
Horizontal length (m)
Average slope (°)
2
Area (km )
f
Proportion of active flow (%)
3
7g
Ice volume (m ) × 10
-1 h
Average annual ablation rate (cm a )
-1 i
Velocity (mm a )
j
Supraglacial debris coverage (%)
Supraglacial debris thick. (cm) (minimum,
maximum, (average))
k
Patterned ground coverage (%)
l
Average polygon diameter (m)
Average polygon trough depth (cm)
m
Near surface englacial debris concentration (%)
Headwall start elevation (m)
Ablation area start (GPR start) elevation (m)
GPR end elevation (m)
Massive ice end elevation (m)
Full glacier horizontal length (m)
GPR coverage horizontal length (m)
Average slope (not including HW) (°)
2b
Nadir area, active flow only (km )
3c
Ice volume (not including HW) (m )
0
NA
NA
<1
0
NA
NA
-
e
1652
1611
824
3
0.408
12.0
4.5–6.8
5.56
5
0, 2, (NA)
Region 1
2247
1642
644
38
0.730
21.5
<1
NA
d
ACA
a
e
55
3
4, 69, (20)
<1, 35, (0.8)
1611
1406
2600
4.5
1.618
47.7
5.2–7.7
0.052
3, 40, (14.3)
100
3, 21, (11)
Region 2
e
100
8.2
25, 125, (74)
2, 25, (2)
1406
1344
918
3.7
0.635
18.7
1.4–2.2
0.013
9, 0.5, (2.9)
100
15, 60, (31)
Region 3
2247
1652
1341
?
> 6000
4420
4.32
3.39
8
1.0–1.6 × 10
Mullins Glacier
100
22.0
40, 120, (110)
15, 50 (25)
1344
N/A
1.5
0.006
~0
100
35, 70, (48)
e
Region 4
0
NA
NA
-
2227
1667
942
35
0.724
27.9
<1
NA
d
ACA
e
0
NA
NA
<1
1667
1597
865
4.5
0.387
14.9
3.1–4.6
6.86
14, 26, (19.1)
15
0, 3, (NA)
Region 1
85
4.9
20, 68, (35)
<1, 35, (0.5)
1597
1484
1440
2.5
0.973
37.5
4.5–6.8
0.050
3, 38, (22.1)
100
6, 29, (12)
e
Region 2
e
1484
1417
695
4
0.509
19.6
0.3–0.5
0.020
4, 0.5, (3.2)
100
15, 57, (22)
Region 3
100
9.8
30, 170, (150)
1.5, 20, (2)
2227
1667
1400
1413
3952
2900
4.24
2.59
8
0.8–1.2 × 10
Friedman Glacier
e
100
21.2
15, 30, (20)
NA
1417
N/A
0
NA
NA
~0
100
69, 71, (70)
Region 4
data indicated by “-.“
The total nadir surface area of the actively flowing portion of each glacier includes the headwall and accumulation zones. The actively flowing portion of each glacier, as determined from Rignot
et al. [2002], includes the portion up to 5.0 km and 3.8 km on Mullins and Friedman Glaciers, respectively.
c
Ice volume was determined by the integration of the local cross-sectional area along the desired distance. The cross-sectional area was estimated using the local width at the glacier surface and
the estimated local depth from the GPR (see section 4.4). In agreement with the transverse GPR profiles, both rectangular and parabolic-shaped cross-sectional areas were used to calculate a range
in the estimated ice volumes.
d
Accumulation area (ACA).
e
Region of the ablation area. See text in section 4.4 for region definitions and details.
f
Defined as the percentage of the total (nadir) area of the actively flowing portion of the glacier.
g
Ice volume determined from GPR as above (see note (c)) for regions specified.
h
Ablation rates in ACZ determined from stake data. Ablation rates estimates for all other regions have been estimated following Kowalewski et al. [2011].
i
Surface parallel flow along the central flow line only, from Rignot et al. [2002]. The three data points listed for each region are in order: minimum, maximum, and average (given in parenthesis).
No data exist for the ACA on either glaciers or Region 1 of Mullins glacier. The values listed for Region 1 on Friedman are based on partial coverage of this region (the down valley 50% of the region)
only. In Region 4, velocities are below the noise threshold and assumed to be near zero.
j
Defined as the percent of the total ice surface covered by a continuous debris cover >1 cm thick.
k
The percentage of the total debris surface area exhibiting some level of patterned ground development.
l
Maximum distance across the polygon surface measured from a trough to the trough center opposite.
m
The percent of debris (by volume) within ice samples excavated from shallow (~50 cm) pits hand dug into the massive ice beneath the surface till. The three data points listed for each region
3
are in order: minimum, maximum, and average (given in parenthesis). Ice debris content estimated for a ~0.75 × 0.75 × 0.5 m (x × y × z) ice volume exposed and sampled at the base of each pits.
Debris content values include both dispersed fine-grain debris and large grain-size debris. It should be noted that the maximum debris content values in Region 2 for both glaciers were very localized and obtained by sampling directly within an inclined debris layer (IDL) (see section 4.4).
a
No
b
Region specific
General
Table 2. Physical Characteristics of Mullins and Friedman Debris-Covered Glaciers
Journal of Geophysical Research: Earth Surface
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Figure 6. Longitudinal transect Am–Am′ for Mullins Glacier. (first panel) Vertical aerial photograph (USGS TMA 3080 V0276)
showing transect location. (second panel) Unmigrated radar data. (third panel) Migrated data with a Hilbert magnitude
transform applied. (fourth panel) Sketch of major radar reflectors and physical interpretation. The vertical depth scale for all
1
panels is shown as two-way travel time (TWTT in ns) and as meters assuming an average radar travel time of 0.168 m ns .
Double-arrowed lines between the first panel and the second panel mark the location of identified arcuate surface discontinuities (ASD) and highlight the spatial correlation between ASD and inclined debris layers (IDL). Longitudinal distance along
the transect line is shown as both the distance from the start of the transect, and as the distance from the valley-headwall
datum (see also Figure 4). Data for the first ~900 m were collected using a 200 MHz antenna in distance mode; thereafter, we
used an 80 MHz antenna and collected data in point mode. The locations of Regions 1–4 are plotted on bottom axis.
4.4.1. Region 1: Mullins Glacier
4.4.1.1. GPR Data
Region 1 of Mullins Glacier extends from 0.6 to 1.5 km (Figure 4). The combined 200 MHz/80 MHz longitudinal
radar transect along Mullins Glacier, Am–Amʹ (Figure 6), is devoid of any major reflections for the first 600 m,
except for a prominent basal reflector. This basal reflector descends sharply from a depth (two-way travel time)
of ~675 ns at the beginning of the transection to a low point of ~1500 ns at ~1.1 km; thereafter, the basal
reflector begins a slight incline toward the surface (Figure 4). Small and isolated hyperbolic reflectors are rare
but locally present at all depths.
4.4.1.2. Ice-Core Data
Four shallow ice cores (maximum depth of 11 m for core MCI-11-02) obtained from Region 1 of Mullins Glacier
all show clean, bubbly ice that is largely devoid of debris (Figure 7). Cores that penetrate the region of
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Figure 7. Location and description of ice cores collected from Regions 1–4 of Mullins Glacier. Depth and general englacial
characteristics for each core are shown in the second panel. In general, the majority of ice cores from Regions 1 and 2 show
clean, bubbly glacial ice with isolated gravels and cobbles (existing cores in Region 2 do not penetrate IDL). Ice cores in
Region 3 uniformly show increasing amounts of gravel-and-cobble-sized clasts, as well as abundant sand veins filled with
sand-and-granule-sized grains of quartz and dolerite grus (e.g., local infill of thermal contraction cracks [Kowalewski et al.,
2011]). Ice from Region 4 contains up to 50% gravel-and-cobble-sized clasts, and, with increasing distance down valley, an
ever increasing concentration of sand veins. As such, ice-core penetration in Region 4 is minimal, and with existing drill
technology (KOCI drill [Green et al., 2007]) is limited to <1 m depth.
superposed ice in the distal end of Region 1 met refusal (additional penetration became impossible) at a
shallow, horizontal rocky layer at ~80–90 cm. This rocky layer, visible at the lateral margins of the superposed
ice, is well displayed in GPR profiles using a 200 MHz antenna [see Shean and Marchant, 2010].
4.4.1.3. Geomorphic Data
Boulders, cobbles, and gravels dotted across exposed ice in Region 1 cover from ~1 to 15% of the glacier
surface; concentrations typically increase down valley (Figures 2 and 6). Where present, small patches of gravel
attain a maximum thickness of ~1 to 2 cm. Isolated boulder-and-cobble-sized clasts are angular, and many
display visible impact scars that likely reflect collisions during rockfall from the valley headwall (percussion
marks [Kowalewski et al., 2011]). Overall, clasts show a subtle increase in slight surface staining (precursor to
oxidation [Salvatore et al., 2013]), which becomes more obvious with increasing distance down valley.
The centimeter-scale microtopography of the ice surface in Region 1 is closely coupled to the distribution,
size, and sorting of the overlying, scattered surface debris. The largest cobble-and-boulder-sized clasts of
Ferrar Dolerite are typically perched on ice pedestals that lie at the center of shallow (15 cm) moat-like
depressions (Figures 2c and 2d). These depressions are best developed on the up valley sides of rocks (e.g.,
upwind), with “tails” of ice and snow occurring in some places on the down valley sides. A sparse mixture of
gravel-sized clasts typically collects at the base of many depressions; in places, these smaller clasts are
lodged beneath the large, central rock, producing an unusual sorting pattern whereby large boulders and
cobbles rest directly on small gravel-sized clasts (Figure 2d). Elsewhere, dolerite cobbles of Region 1 are found
immediately below the ice surface, covered with a very thin (<1 cm) layer of refrozen meltwater (Figure 2e),
or embedded to greater depths (up to ~20 cm), with a “tube” of refrozen meltwater connecting them to
the ice surface (e.g., at the base of cryoconite holes [Fountain et al., 2004]). As noted above, the down valley
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Figure 8. Longitudinal transect Af–Afʹ for Friedman Glacier. (first panel) Vertical aerial photograph (USGS TMA 3080 V0276)
showing transect location. (second panel) Unmigrated radar data. (third panel) Migrated data with a Hilbert magnitude
transform applied. (fourth panel) Sketch of major radar reflectors and physical interpretation. The vertical depth scale for all
1
panels is shown as two-way travel time (TWTT in ns) and as meters assuming an average radar travel time of 0.168 m ns .
Double-arrowed lines between the first panel and the second panel mark the location of identified arcuate surface discontinuities (ASD) and highlight the spatial correlation between ASD and inclined debris layers (IDL). Longitudinal distance along
the transect line is shown as both the distance from the start of the transect and as the distance from the valley-headwall
datum (see also Figure 4). Data for the first ~750 m were collected using a 200 MHz antenna in distance mode; thereafter, we
used an 80 MHz antenna and collected data in point mode. The locations of Regions 1–4 are plotted on bottom axis.
limit of Region 1 on Mullins Glacier displays a small area of superposed ice (12,490 m2, 1.5–1.8 m deep) that in
plan view appears as an oblong, frozen “pond” (Figures 3 and 6).
4.4.2. Region 1: Friedman Glacier
4.4.2.1. GPR Data
Region 1 of Friedman Glacier extends from 0.9 to 1.7 km. Beginning ~900 m from the headwall of Friedman
Valley, the first 40 m of the combined 200/80 MHz longitudinal radar transect along Friedman Glacier, Af –Afʹ
(Figure 4), appears clean and lacks major reflections apart from a prominent basal reflector (Figure 8).
This basal reflector descends from a depth of 775 ns at the beginning of the transect and reaches a
maximum depth of 1240 ns at 1250 m before rising gradually toward the surface. Dispersed hyperbolic
point reflections and a strong horizontal reflection at 20 ns depth are apparent in the shallow 400 MHz
profile (Figure 8).
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4.4.2.2. Geomorphic Data
A mixture of boulders, cobbles, and gravels is sparsely distributed over 3–35% of the ice surface in Region 1 of
Friedman Glacier; where present, small patches of gravel-sized clasts can locally achieve a maximum
thickness of ~3 cm. Debris is most concentrated near the down valley end of the region. Clast characteristics
and ice surface microtopography are nearly identical to those found on Mullins Glacier. Just as for the case
of Mullins Glacier, the down valley limit of Region 1 displays a small area of superposed ice (14,880 m2,
1.5–1.8 m deep) that appears in plan view as an oblong, frozen pond (Figures 3 and 8).
4.4.3. Region 1: Interpretation
The salient subsurface features within Region 1 of both Mullins and Friedman Glaciers include a strong basal
reflector overlain by mostly homogeneous material that displays isolated hyperbolic reflectors. Along with
the data from shallow ice cores, this suggests that the predominately homogeneous and reflector-poor areas
represent clean glacier ice that is devoid of significant foliation or debris layers [e.g., Berthling et al., 2000]. The
strong coherent negative-positive-negative polarities of the basal reflector most likely represent the
intersection with valley beds. The smooth geometry of the basal reflections, as well as the sharp degradation
in signal penetration below them, suggests glacially sculpted bedrock rather than loose, unconsolidated, and
irregularly shaped subglacial debris/talus. The isolated hyperbolic reflections dotted throughout the
reflector-poor areas, and especially apparent in the higher-frequency transects (e.g., 400 MHz), most likely
represent individual/small groupings of clasts within the ice. This interpretation is supported by the high
signal amplitude and negative-positive-negative polarity signature of these reflections [Arcone et al., 2002].
The microtopography of the ice surface in Region 1 is directly related to the uneven distribution of surface
debris. The largest rocks alter proximal winds, producing up valley wind scoops and, in places, down valley
ridges. Further, enhanced ablation along clast margins (but not at their base) gives rise to ice pedestals
beneath the largest rocks [Doran et al., 2000; Hendy, 2000]. The inverse grading that places large cobbles and
boulders on top of small gravel-sized clasts most probably arises from the eventual toppling of these
“pedestal” rocks on gravel-sized clasts trapped in local depressions.
4.4.4. Region 2: Mullins Glacier
4.4.4.1. GPR Data
Region 2 of Mullins Glacier begins at the abrupt onset of the continuous supraglacial debris and extends from
1.5 to 4.1 km (Figures 3 and 6). The GPR data from this region are dominated by six distinct englacial reflectors
that, in longitudinal transects, predominately dip up valley, and in transverse profile appear parabolic in
shape (Figure 6). These reflectors are denoted as Mullins Inclined Reflectors 1, 2, 3, etc. (e.g., MIR 1 through
MIR 6) and are described in detail below. Interspersed with these inclined reflectors are isolated hyperbolic
reflectors, which are observed in all 80 MHz, 200 MHz, and 400 MHz transects. The basal reflector continues its
trajectory from Region 1 but is increasingly diffuse with distance down valley (Figure 6); it eventually shoals to
a depth of 190 ns at 3400 m along GPR transect Am–Amʹ before deepening slightly to 310 ns at 4000 m.
MIR1 emerges at ~1050 m from a position slightly above the basal reflector in Region 1 and intersects the ice
surface at 1600 m, at the apex of the first arcuate surface ridge and onset of continuous supraglacial cover
(Figure 6); the angle of intersection of MIR 1 with the ice surface is ~30°. Close inspection of transect Am–Amʹ
reveals a distinct change in MIR 1 at ~1500 m, where the reflector deepens slightly before continuing upward
toward the surface. Transverse profile Cm–Cmʹ (Figure 9) crosses this reflector and reveals its parabolic shape,
nearly 500 m wide at the transect location; the reflector intersects the ice surface just beneath the lateral
margins of the first surface ridge.
MIR 2 begins at ~1100 m and rises upward slightly before remaining nearly bed parallel for 300 m. Ultimately,
MIR 2 intersects the ice surface at 2000 m, showing an intersection angle of ~35° (Figure 6). Analysis of a
detailed portion of MIR 2 indicates that it is composed of a series of closely spaced, aligned hyperbolic
reflections (Figure S3 in the supporting information). Data from transverse profile Cm–Cmʹ show that MIR 2 is
parabolic in shape and that the deepest up valley portion of this reflector is nested beneath MIR 1 (Figure 9).
MIR 3 displays the most complex geometry of any of the identified inclined reflectors in Mullins glacier. The
basal portion of this reflector is difficult to differentiate from MIR 2, but it appears to emerge just above the bed
near ~1700 m; it intersects the ice surface at 2550 m, with an intersection angle of ~6° (Figures 6 and S3).
The general upward trend of MIR 3 is interrupted by three steps where, at each step, the reflector remains nearly
bed parallel for ~300–400 m (Figures 6 and S3). The last of these three bed-parallel steps occurs just beneath the
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ice surface where the primary reflector
splits into two reflectors: the primary
reflector intersects the surface at
2500 m, and the secondary reflector
remains below the surface for an
additional ~50 m before intersecting the
surface at 2550 m (Figure S3). MIR 3 is
~300 m wide where it intersects the ice
surface at the location of transverse
profile Em–Emʹ (Figure S4 in the
supporting information).
MIR 4 is the least distinct of the major
inclined reflectors. It emerges from near
the bed at ~1700 m and remains nearly
bed parallel for much of its length before
ascending and intersecting the ice
surface at ~2200–2300 m; the angle of
intersection is ~17° (Figure 6). MIR 4 is
410 m wide where it intersects the
surface along transverse profile Fm–Fmʹ
(Figure S5 in the supporting information).
MIR 5 has a unique geometry departing
significantly from the other major
reflectors. Beginning at 2800 m, it takes
the form of a large deformed U-shaped
reflector reaching a depth of ~25 m at
its lowest point (Figure 6). The down
valley extension of MIR 5 intersects the
surface at 3000 m with an intersection
angle of ~35°. MIR 5 is 510 m wide
where it intersects the surface along
transverse profile Fm–Fmʹ (Figure S5).
MIR 6, the final inclined reflector,
defines the boundary between Region 2
and Region 3 (Figures 3 and 6). It is
unique in that it does not display a clear
bed-parallel extension but instead
emerges at a steep angle from near the
bed at 3990 m and intersects the ice
surface at 4020 m (intersection angle of ~26°) (Figure 6). The surface intersection of MIR 6 coincides with the
down valley margin of a major ice lobe, showing ~15 m relief, which, in turn, roughly coincides with the initial
descent of Mullins Glacier into upper Beacon Valley (Figures 3, 4, and 6). MIR 6 is ~250 m wide where it
intersects the ice surface along transverse profile Hm–Hmʹ (Figure S6 in the supporting information).
4.4.4.2. Ice Excavations and Ice Core Data
Eleven shallow ice cores recovered from Region 2 of Mullins Glacier (all <24 m deep) show clean glacier ice
interspersed with isolated, sand-and-gravel-sized debris ≪ 1% by volume (Figure 7).
Figure 9. Transverse radar profile Cm–Cmʹ on Mullins Glacier, collected
with the 200 MHz antenna in distance mode (see Figure 4 for transect
location). (top) Unmigrated radar data. (middle) Migrated data with a
Hilbert magnitude transform applied. (bottom) Sketch of major radar
reflectors and physical interpretation; also shown are the locations of
corresponding arcuate surface discontinuities (ASD).
Fifty-five exposures of the buried ice surface beneath supraglacial debris in Region 2 document an overall
pattern of relatively clean ice (0.5–1% debris by volume) alternating with narrow bands of significant
englacial debris (each band containing up to 35% debris by volume). The regions of concentrated englacial
debris occur precisely where GPR data show inclined reflectors intersecting the glacier surface (e.g., MIR 1–6);
clasts in these bands dip up valley, at angles between 7 and 25°, and, on close inspection, appear to rest
directly on top of small pebbles; the latter produces an inverse grading pattern that is reminiscent of the
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Figure 10. Stable isotopic data and electrical conductivity measurements from ice recovered along Mullins and Friedman
18
18
Glaciers. (a) Plot of δD versus δ O. Also plotted is the local meteoric water line (LMWL) δD = 7.75 (δ O) 7.24 [Gooseff
18
et al., 2006] and the global mean water line (GMWL) is δD = 8.0 (δ O) + 10 [Craig, 1961]. (b) Deuterium excess versus δD for
samples plotted in Figure 10a. Deuterium excess values are both slightly negative and slightly positive with the majority
clustering around zero (min = 5.2‰, max = 3.3‰, mean = 0.6‰). Analytical precision is ±0.1‰. Isotope values are
18
presented as permil (‰) relative to Vienna SMOW. Deuterium excess values were calculated as d = δD 8 δ O after
Dansgaard [1964]. (c) Stable isotopic composition, electrical conductivity (EC), and total dissolved solids (TDS) plotted for
ice samples from Mullins and Friedman Glaciers. EC values have been adjusted to a reference temperature of 25°C using a
1
standard conversion factor 1.9% °C . EC precision is ±1.0%. Data are also presented as TDS where TDS = EC × 0.51.
stacking of clasts at the ice surface in Region 1. Inspection of clasts dug out from inclined bands reveals that
most display impact scars identical to those observed on the surface of clasts exposed in Region 1. No glacially
striated or faceted clasts were observed in the inclined debris bands (or elsewhere within Mullins Glacier).
Lastly, exposures of buried ice across Region 2 reveal the presence of near-vertical cracks, each typically less
than 2 cm wide and filled with stained quartz sand and dolerite grus (see also section 4.4.5.3 below).
Isotopic analyses of glacier ice (from 0.5 to 1 m depth) recovered from Region 2 fall along a local meteoric
water line (LMWL) when plotted on a graph of δ18O versus δD [Gooseff et al., 2006], with a best fit line
yielding a slope of δD = 8.0 × δ18O 2.14 (R2 = 0.97; Figure 10). Deuterium excess is near zero but displays
both slightly positive (maximum = 3.8) and negative values (minimum = 6.1). No trend is apparent with
proximity to inclined reflectors/observed englacial debris: ice within and beyond bands of debris/inclined
reflectors (e.g., tens of meters distant) show similar isotopic ratios. Similarly, ice samples from Region 2
display no significant trends in total dissolved solids (TDS) as a function of proximity to inclined reflectors
(Figure 10).
4.4.4.3. Geomorphic Data
The thin supraglacial debris of Region 2 increases steadily in average thickness from ~8 to 20 cm (Figure S7b
in the supporting information) and is arrayed in arcuate ridges and steps (Figures 3, 4, and 6). We refer to each
notable step/ridge by the nongenetic phrase “arcuate surface discontinuity,” or ASD; each ASD is numbered
sequentially down valley, with number 1 occurring at the onset of uniform debris cover (Figures 3 and 6).
Region 2 of Mullins Glacier contains six ASD (Mullins ASD 1 through 6). In plan view, Mullins ASD 1–5 all
display an elongated, arcuate pattern, with long tails extending several kilometers up valley; these tails all
converge toward similar points very near the valley headwall. Mullins ASD 6, which marks the down valley
boundary with Region 3, is unique in overall height (10 m relief) and form (lobate front), but it too can be
traced well up valley. Each ASD on Mullins Glacier occurs where GPR data show that inclined reflectors
intersect the ice surface.
Mullins ASD show the following subtle, but consistent stepwise transitions (Table 3 and Figures 11 and 12):
1. Slope change. The buried ice surface, and overlying debris, rise toward each ASD; this forms an asymmetrical step/ridge.
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1:7
NA : 51
NA : 45
NA : 38
0 : 90
NA : 15
NA : 2.7
0:8
Ridge
0.5–1 : 20
Increase
Increase
Increase
Ridge “couplet” formation
Increase
ASD 1
- : 0.25
5 : 2
No ridge
0.5–1 : 5 : 20
8 : 21
67 : 51
29 : 44
30 : 9
40 : 30
ASD 2
1.4 : 2.5
4 : 0
Ridge
0.5–1 : 25
4 : 25
7 : 20
68 : 52
25 : 44
12 : 23
5 : 95
ASD 3
- : 1.2
3 : 3
No ridge
0.5–1 : 3
2 : 20
11 : 15
-:
-:
-:
5 : 30
ASD 4
Mullins Glacier (Up Valley : Down Valley)
Increase
Decrease
Increase
Increase
Increase
Observed UV to DV Stepwise Δ
b
1.5 : 5.1
3 : 1
Ridge
1 : 30
15 : 69
17 : 30
85 : 10 : -:
40 : 100
ASD 5
NA : 2.1
0:5
Ridge
0.5–1 : 20
NA : 22
1 : 15
NA : 61
NA : 34
NA : 8
0 : 65
ASD 1
0.5 : 4.2
8 : 3
Ridge
0.5–1 : 5 : 25
7 : 15
81 : 71
16 : 21
10 : 25
15 : 100
ASD 2
1.1 : 4.5
7 : 0
No ridge
1 : 15
10 : 30
7 : 13
77 : 70
18 : 27
7 : 30
65 : 100
ASD 3
3 : 4.5
4 : 2
Ridge
-:
25 : 35
8 : 15
-:
-:
-:
80 : 100
ASD 4
2.5 : 5.2
7 : 1
No ridge
0.5 : 25 : 94
8 : 25
81 : 13 : 6:
80 : 100
ASD 5
Friedman Glacier (Up Valley : Down Valley)
Arcuate Surface Discontinuities (ASD)
No data indicated by “-.“ Average values are given for the zones immediately up valley of the ASD and down valley of the ASD, separated by a colon. The area that is considered to be immediately up valley and immediately down valley of the ASD varies with each parameter (see below) but ranges from 1 to 40 m. The stepwise changes in surface characteristics across ASD 6, which
mark the boundary of region 2 and region 3, are not represented here. Instead, they are explained in detail in the text.
b
The dominant stepwise change observed when comparing the area immediately up valley (UV) of and an ASD to the area down valley (DV) of and ASD.
c
Average calculated using the average till thicknesses for excavations within 100 m laterally from the central flow line and within 40 m (up or down valley) of the ASD. The number of samples
used for each thickness average varied from N = 3–10.
d
Average calculated using the grain-size fractions for all excavations within 100 m laterally from the central flow line and within 50 m (up or down valley) of the ASD. The number of samples used
for each average varied from N = 1–7.
e
Count of clasts >16 mm. Average of values for excavations within 30 m (up or down valley) of the ASD. The number of samples for each average varied from N = 2–7.
f
We define the area analyzed for ASD-related patterned ground analysis as the region bound longitudinally by the ASD and a parallel arc 50 m up or down valley of the ASD.
g
Elevation profiles were extracted from 1.5 m LiDAR [Schenk et al., 2004] DEM along the central flow line in the vicinity of each ASD and smoothed with a five-point moving average. The slope
values, derived from the elevation profiles were smoothed across a 10-point moving average to eliminate the impact of patterned ground. The reported up valley and down valley values are
averages for 35 m above and below the ASD, respectively.
h
3
Ice debris content estimated for a ~1 × 1 × 0.5 m (x × y × z) ice volume exposed at the base of selected pits. Debris content values include both dispersed fine-grain debris and large grain-size debris.
a
Thickness (cm)
d
Sand (wt %)
d
Gravel (wt %)
e
Lg gravel (N)
f
Mesoscale features Polygon prevalence
(% cover)
f
Polygon vertical
relief (cm)
f
Polygon diameter (m)
g
Local slope (deg)
Ridge?
Ice
Shallow ice debris
h
content (vol %)
Surface debris
c
General
a
Table 3. Variation in Surface Characteristics Across Arcuate Surface Discontinuities (ASD) in Region 2
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Figure 11. Sedimentological variation in supraglacial debris
on Mullins Glacier as a function of proximity to arcuate surface discontinuities (ASD 1–3) (locations highlighted as
shaded bars); grain sizes and sampling procedures for gravel
and sand as prescribed in Figure S7 in the supporting information. Smooth lines represent a two-point running average. Both grain size and debris thickness display consistent
and abrupt changes across ASD.
10.1002/2014JF003178
2. Debris thickness. The thickness of supraglacial
debris increases by 20–30% at, and immediately down valley from, each ASD.
3. Bulk grain-size distribution. The concentration of
gravel-sized clasts (<16 mm) and measured
cobble-and-boulder fractions increases by a
factor of ~1.5–2.0 across ASD.
4. Patterned ground. Thermal contraction crack polygons are more ubiquitous and possess larger
diameters and deeper troughs immediately
down valley of ASD.
5. Debris content of subjacent ice. The percentage of
observed englacial debris increases from its
typical value of ~ <1% by volume to as much as
35% by volume in the immediate vicinity of ASD.
Importantly, this ice lacks increased silt or amber
color that, for example, typically characterizes
basal ice in alpine glaciers in the Dry Valleys
region [Cuffey et al., 2000a; Mager et al., 2009];
see section 6 below.
Finally, we note that minor arcuate topographic
features are also present in Region 2 (“secondary
lineations” in Figure 3), but these are small, typically
discontinuous and do not display the suite of stepwise transitions in physical properties as observed in Mullins
ASD 1–5; furthermore, these secondary lineations are not associated with inclined, englacial reflectors.
Figure 12. Diagrammatic illustration showing the typical near-surface geomorphology for ASD. (a) Block diagram highlights changes across a typical ASD (see also Table 3). (b) Generalized cross section (longitudinal profile) showing topographic changes associated with ASD.
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4.4.5. Region 2: Friedman Glacier
4.4.5.1. GPR Data
Region 2 (Figure 3) extends from 1.7 to 3.4 km on Friedman Glacier and coincides with the abrupt onset of the
complete cover of supraglacial debris. The basal reflector, continuing from Region 1, is diffuse and gradually
shoals toward a minimum depth of 360 ns at 3.4 km (Figure 8). Six inclined reflectors (FIR 1 to FIR 6; e.g., Friedman
Inclined Reflectors 1, 2, and 3) dominate GPR transects across Region 2 (Figure 8); each inclined reflector
intersects the ice surface at a specific arcuate surface discontinuity, ASD, e.g., Friedman ASD 1 through 6
(Figures 3 and 8). Well-defined hyperbolic point reflectors are observed between adjacent FIR in both 80 MHz
(Figure 8) and 400 MHz (Figure S2 in the supporting information) transects. Details of the six major inclined
reflectors are presented below.
FIR 1 originates from just above the basal reflector and ascends upward to intersect the ice surface at Friedman
ASD 1 (~1700–1780 m) (Figures 8 and S8 in the supporting information). The smooth inclination of FIR 1 is
interrupted at ~1510 m where it shallows for 15 m, producing a short horizontal section before rising again
toward the ice surface (Figure 8). The shallowest portion of FIR 1 (0 to ~20 m depth) splits into two reflectors
that intersect the ice surface, at similar sub 25° angles, ~80 m apart (Figure S8). The 400 MHz transect
shows that at least the uppermost portion of FIR 1 is composed of individual, closely spaced, hyperbolic
reflections (Figure S2).
FIR 2 originates at a depth of ~1350 m near the onset of FIR 1 and rises smoothly to intersect the ice surface at
~1900 m, at a slight topographic ridge, e.g., Friedman ASD 2 (Figures 8 and S8); the angle of intersection with the
ice surface is ~27°. At the location of transverse profile Ff–Ffʹ, FIR 2 is roughly parabolic in shape and reaches a
maximum with of ~300 m where it intersects the buried ice surface (Figure 13). At depth, FIR 2 narrows and displays
moderate lateral asymmetry, with its central axis offset slightly to the east of the valley centerline (Figure 13).
FIR 3 originates above the basal reflection at ~1550 m (Figures 8 and S8). It extends down valley at a constant
bed-parallel orientation for ~450 m at 700 ns depth before rising sharply toward the ice surface at 2100 m; it
intersects the ice surface at 2200 m at ASD 3, with an intersection angle of ~36°. In transverse profile Ff–Ffʹ
(Figure 13), FIR 3 is ~450 m wide where it intersects the buried ice surface.
FIR 4 is the least distinct of the major inclined reflectors and originates above the basal reflection at ~2150 m
(Figure 8). Thereafter, it rises smoothly to intersect the ice surface at 2350 m, at Friedman ASD 4; the angle of
intersection with the ice surface is ~25°.
FIR 5 displays the most complex geometry of any of the major inclined reflectors in Friedman Glacier. The
basal portion of this reflector is difficult to differentiate from FIR 3, but it appears to originate above the basal
reflector near ~1500 m. It intersects the ice surface nearly 1200 m distant, at 2700 m at the location of
Friedman ASD 4; the intersection angle with the ice surface is ~13° (Figure 8). Two distinct bed-parallel
portions mark FIR 5: the first extends from ~1500 m to 2400 m and the second from 2450 to 2650 m; the latter
corresponds to a distinct gain in reflector thickness. The transverse profile across FIR 4, transect Gf–Gfʹ (Figure
S9 in the supporting information), is more complex than that of the other reflectors but shows a roughly
parabolic shape that is ~300 m wide at the buried ice surface.
FIR 6 defines the boundary between Region 2 and Region 3 on Friedman Glacier. FIR 6 begins at ~2800 m,
remains bed parallel for ~500 m, and ultimately rises to intersect the ice surface at a dipping angle of ~16° at
3400 m (Figure 8).
4.4.5.2. Ice Excavations
Exposures of the buried ice surface observed at the base of 26 soil pits in Region 2 of Friedman Glacier
document an overall pattern of relatively clean ice (0–1% debris by volume) alternating with narrow bands
of significant englacial debris (each band contains as much as 35% debris by volume). The regions with
elevated englacial debris occur at ASD and precisely where GPR data show inclined reflectors intersecting the
glacier surface (e.g., FIR 1–6); visible clasts in these debris bands dip up valley, at angles between 7 and
25°, and, on close inspection, appear to rest directly on top of small pebbles. Some of the clasts also show
impact scars, identical to those observed on clasts at the ice surface in Region 1, but none show striations or
faceting. Lastly, exposures of buried ice across Region 2 reveal the presence of near-vertical cracks, each
typically less than 2 cm wide and filled with stained quartz sand and dolerite grus. Isotopic and dissolved salt
content analysis of ice samples shows no trend related to proximity with ASD (Figure 10).
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4.4.5.3. Geomorphic Data
The thin supraglacial debris of Region 2
shows increasing trends in average
thickness (from ~6 to 30 cm) and sandsized fraction in the down-ice direction
(Figure S10 in the supporting
information). Friedman ASD displays
similar stepwise changes in underlying
ice surface slope, topography, and
grain-size characteristics (Figure 12) as
those mapped on adjacent Mullins
Glacier (see section 4.4.4.3 above and
Table 3). Each ASD on Friedman Glacier
occurs where GPR data show that
inclined reflectors appear to intersect
the ice surface and where englacial
debris increases abruptly to >30% by
volume. Additional minor arcuate
topographic features are also present in
Region 2 (secondary lineations in
Figure 3), but these are small, typically
discontinuous, and do not display the
suite of stepwise transitions in physical
properties observed in Friedman ASD 1
through 5; furthermore, they are not
associated with the inclined, englacial
reflectors.
4.4.6. Region 2: Interpretation
We interpret the set of six inclined
reflectors (MIR 1 to 6, FIR 1 to 6) as
narrow bands of closely spaced rocky
debris, hereafter termed inclined
debris layers (IDL, 1–6). This
interpretation is based on the high
signal amplitude of their radar returns,
Figure 13. Transverse radar profile Ff–Ffʹ collected with the 80 MHz
the negative-positive-negative
antenna in point mode (see Figure 4 for location). (top) Unmigrated radar
polarity signature of their radar
data. (middle) Migrated data with a Hilbert magnitude transform applied.
returns, and the significant volume of
(bottom) Sketch of major radar reflectors and physical interpretation.
englacial rocky debris present at their
points of surface intersection (e.g.,
exposed in surface ice at ASD). The basal reflector in Region 2 of both Mullins and Friedman Glaciers is
relatively diffuse and most consistent with scattered debris approximately 10–150 ns “thick” and with
the bedrock interface at some unknown depth below. Isotopic analyses of near-surface ice from
Region 2, which plot on a local meteoric water line, indicate minimal—if any—melting [Gooseff et al.,
2006; Sleewaegen et al., 2003; Souchez and Lorrain, 2006] regardless of proximity to ASD and IDL
(see Figure 10).
4.4.7. Region 3: Mullins Glacier
4.4.7.1. GPR Data
Region 3 on Mullins Glacier extends from 4.1 to 5.0 km. The 80 MHz longitudinal radar transect Am–Amʹ
reveals an increase in high-frequency noise and dispersed reflections. Region 3 displays areas of
chaotic, closely spaced reflections with near-vertical dipping angles (Figure 6). Though increasingly
diffuse with distance down valley, the basal reflector appears to deepen along the glacier axis, from
~260 ns at 4.1 km to a maximum of ~620 ns at 4.4 km before shoaling up to ~500 ns at
5.0 km (Figure 6).
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4.4.7.2. Ice Excavations and Ice Core Data
All cores in Region 3 of Mullins Glacier show clean bubbly ice that displays a subtle increase in englacial gravel
and cobble-sized clasts (~2–3% by volume) relative to cores from Regions 1 and 2 (Figure 7). Sand veins show
a marked increase in concentration, especially as observed at the ice surface, and reach a minimum depth of
~14 m (as observed in ice cores). These sand veins occur as subhorizontal to near-vertical layers, taper with
depth, and are filled with stained quartz sands and weathered dolerite grus.
4.4.7.3. Geomorphic Data
Region 3 of Mullins Glacier shows an abrupt increase in the thickness of supraglacial debris (Figure S7b) and
a concomitant increase in the maturation of thermal contraction crack polygons (Table 2); the latter
includes an increase in average polygon diameter and depth of bounding troughs [e.g., Levy et al., 2006;
Marchant et al., 2002]. Isotopic analysis of near-surface ice (0.5–1 m depth) fall on the local meteoric water
line (LMWL) (Figure 10).
4.4.8. Region 3: Friedman Glacier
4.4.8.1. GPR Data
Region 3 of Freidman Glacier extends from 3.4 to 4.8 km. The 80 MHz radar (Af–Afʹ) displays an overall
increase in noise and dispersed reflections relative to Regions 1 and 2 (Figure 8). No distinct IDLs are present.
The basal reflection is vague and difficult to define, though it appears to remain at a nearly constant depth
of ~400 ns.
4.4.8.2. Ice Excavations and Geomorphic Data
Region 3 shows an abrupt thickening in supraglacial debris and an increase in the maturation of thermal
contraction crack polygons (Table 2). Ice collected within the top 10–50 cm of buried glacier ice show an
increase in the concentration of both sand veins and sporadic gravel-and-cobble-sized debris (collectively up
to ~2% by volume).
4.4.9. Region 3: Interpretation
The GPR data suggest that Region 3 of both Mullins and Friedman Glaciers contains mostly clean ice, but with
an increasing concentration of dispersed debris relative to that found in Regions 1 and 2. The increase in
widespread englacial debris is inferred from (1) spurious reflections within the GPR data, (2) direct
observation of debris within shallow ice cores, including sand veins and isolated clasts, and (3) visible
englacial debris observed at the base of soil pits and within shallow ice pits.
4.4.10. Region 4: Mullins Glacier
4.4.10.1. GPR Data
Ice in Region 4 of Mullins Glacier (4300 m to 4550 m along Am–Amʹ (Figure 6)) lacks distinct reflectors and
instead displays chaotic signal returns; the 80 MHz signal penetration is less than that of Region 3. A distinct
basal reflector is not observed.
4.4.10.2. Ice Excavations and Geomorphic Data
Ice exposed at the base of soil pits in Region 4 shows a large increase in the concentration of gravel-andcobble-sized clasts. Visible englacial debris reaches ~40–50% by volume, and ice coring is nearly impossible.
At present, only one shallow core, 1 m deep, has been extracted; the core contains 35% dispersed sand and
gravel, with isolated cobbles containing fresh faces and preserved impact scars (percussion marks
[Kowalewski et al., 2011]). Supraglacial debris in Region 4 ranges from ~40 cm to 70 cm thick, typically
increasing down valley; the debris is cut by well-developed contraction-crack polygons, each with diameters
approaching 20 m (Table 2). As noted above, data from satellite interferometry suggest that horizontal ice
velocity in Region 4 approaches zero, essentially <2 mm a1 and indistinguishable from measurement noise
[Rignot et al., 2002].
4.4.11. Region 4: Friedman Glacier
4.4.11.1. GPR Data
GPR data from Region 4 of Friedman Glacier (from 2850 m to 3050 m along Af Afʹ) show chaotic signal
returns; the 80 MHz signal penetration depth is significantly reduced relative to that in Region 3 (Figure 8).
A distinct basal reflector is not observed.
4.4.11.2. Geomorphic Data
Region 4 commences immediately below the lobate front of Friedman Glacier. The surface exhibits large
diameter polygons (Table 2) (20+ m diameter) that cut into unconsolidated debris at least 70 cm in
thickness. At present, it is unclear if glacier ice exists beneath this thick debris cover; if so, it is unreachable
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in hand-dug excavations and not
resolvable in our GPR data. According
to data from satellite interferometry,
surface velocities in this region
approach zero and become
indistinguishable from measurement
noise [Rignot et al., 2002].
4.4.12. Region 4: Interpretation
The closely spaced, high-amplitude
reflections in the GPR data for Mullins
Glacier are interpreted as debris-laden
ice of unknown depth. Individual
hyperbolic reflectors cannot be
identified within GPR transects,
suggesting that individual clasts in
Region 4 are not spaced sufficiently far
to be resolved separately in the 80 MHz
data, e.g., most likely < ~40 cm apart.
Our assertion that calls for debris-laden
Figure 14. (a) On Mullins Glacier, 400 MHz transect Km–Kmʹ collected ice (e.g., ice volume > pore space)
rather than ice-cemented debris (e.g.,
across the site of ice core MCI-08-01 (see Figures 4 and S1 for location).
(left) Unmigrated data; (right) migrated data with a Hilbert magnitude
ice concentrations ≤ pore space) is
transform applied. The location of the core is indicated by the grey vertical supported by visual inspection of the
bars. At the location of the borehole, the GPR indicates a major layer of
buried ice surface at dozens of localities
debris at a two-way travel time of ~270 ns. This debris represents MIR-4/
IDL-4. (b) Sketch map of MCI-08-01 showing very little debris except for a and shallow ice cores that penetrate
this region. Friedman Glacier shows
few cobbles and gravels at 8 and 16 m and ultimately refusal at 22.7 m.
The cobble-rich layer intersected at 22.7 m is almost certainly the same
similar GPR data (as that observed in
layer imaged at 271 ns. This depth equates to an average radar travel
Mullins Glacier), but because buried ice
1
velocity of 0.168 m ns .
was not physically encountered at the
base of soil pits in this region, the relative abundance of ice (ice cement versus excess ice, or even a lack of
subsurface ice) remains unknown in Region 4 of Friedman Glacier.
4.5. Determining the Depths of GPR Reflectors
In the absence of a common midpoint profile, we determined a representative average radar travel velocity
using measured depths to reflectors from ice-core data (refusal) and confirmed this estimate indirectly
through iterative tests of hyperbola collapse during GPR data migration. Borehole MCI-08-01 in Region 2 of
Mullins Glacier met refusal at an impenetrable layer at 22.7 m (MIR 4) (Figure 14). Correlating this depth with a
strong reflector observed in the colocated 400 MHz GPR profile (Km–Kmʹ) yields a travel velocity of
0.168 m ns1 (dielectric of 3.18) (Figure 14). This value is slightly higher than the velocity value determined by
hyperbola collapse for the headwall area of Mullins Glacier (0.167 m ns1 [Shean and Marchant, 2010]) and
lower than values determined for a debris-covered glacier on James Ross Island, Antarctica (0.17 m ns1) [e.g.,
Fukui et al., 2008]. However, it is within the range of typical radar velocities expected for cold glacial ice
[Plewes and Hubbard, 2001, and references therein] and provided excellent hyperbola collapse during
iterative tests of radar data migration.
Although this calculated radar velocity is robust at the locations of measured boreholes, it is not
necessarily representative of heavily debris-laden ice. In these areas, the travel velocity will likely be slower
than our reported velocities, and more consistent with other GPR-travel velocities through debris-rich ice,
e.g., 0.12 m ns1 [Degenhardt and Giardino, 2003], 0.15 m ns1 [Lehmann et al., 1998], 0.14 m ns1 [Isaksen
et al., 2000], and 0.12 m ns1 [Monnier et al., 2008]. The net effect of using a single-average radar travel
velocity of 0.168 m ns1 will be to slightly overestimate the thickness of englacial debris layers and, as a
result, slightly overestimate the depths for rays that travel through these layers. For this reason, we
consider our depth-conversion estimates as maximum values and assign a conservative error of 5% for
reported values (equivalent to using an average velocity of 0.160 m ns1 or a dielectric of 3.51).
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In summary, using an average radar travel velocity of 0.168 m ns1, we arrive at the following: Mullins Glacier
reaches a maximum thickness of ~125 m near its headwall at 1.1 km in Region 1; it thins to a depth of ~22 m
where its ice surface slope increases near the terminus of Region 2 (3.4 km) and thickens again to ~45 m at
4.6 km (ice thicknesses beyond 4.6 km were not measured in this study, but see Shean and Marchant [2010]
for additional data). Friedman Glacier reaches its maximum depth of ~105 m at ~1.25 km in Region 1;
thereafter, it thins gradually to ~45 m depth at ~3.6 km in Region 3; if ice continues beyond this point [e.g.,
Kowalewski et al., 2011], it is unresolvable in GPR data using 80 MHz antennas. Both glaciers display parabolic
cross-sectional geometries and rest on a base of inferred rocky debris 3–10 m thick (Regions 2–4) or directly
on inferred bedrock (Region 1).
4.6. Ground Penetrating Radar, Ice Cores, and Surface Geomorphology: Interpretation Synthesis
Mullins and Friedman Glaciers share similar surface and internal characteristics that permit subdivision into
four distinct regions. Region 1 consists of relatively clean glacier ice, >100 m thick, with scattered englacial
debris (<1% by volume). This ice rests on smooth reflectors, most probably bedrock, and is overlain by a
discontinuous cover of sporadic cobble-and-gravel-sized clasts. Region 2 also contains relatively clean
glacier ice, ~100–25 m thick but additionally includes six, parabolic-shaped, inclined debris layers (IDL)
composed of concentrated, rocky debris from 1 to 3 m thick. IDL originate from above the glacier bed and
intersect the ice surface at marked, arcuate surface discontinuities (ASD) (Figure S11 in the supporting
information). ASD typically occur at longitudinal intervals of ~200–400 m and exhibit long tails that extend
back toward ice accumulation areas; isotopic analyses and total dissolved salt content of ice show no
systematic variation across individual IDL/ASD. Each ASD shows near-identical changes in ice surface slope,
supraglacial debris thickness, patterned ground, and changes in bulk grain-size distribution. Regions 3 and
4 lack well-defined IDL and show stagnant to near-stagnant ice [Rignot et al., 2002] with increasing
concentrations of dispersed, englacial debris and a concomitant increase in the thickness of overlying,
supraglacial debris.
5. Results: Modeling
5.1. Modeled Glacier Thickness
We utilized a simple analytical steady state flow model to generate expected centerline ice thicknesses for
comparison with those derived from GPR. This exercise was carried out in order to (1) provide an independent
test of our calculated average GPR travel velocity, (2) confirm our estimates of glacier bed depth—especially in
the distal portion of each glacier where the basal reflector is diffuse—and (3) provide initial estimates of the ice
hardness parameter for use in future numerical modeling studies. A similar approach was adopted by Shean
and Marchant [2010] to estimate ice thickness in upper Mullins Glacier and by Rignot et al. [2002] to estimate
first-order ice thickness throughout a large portion of the Quartermain Mountains. Similar to Konrad et al. [1999]
and Konrad and Humphrey [2000], we use Glen’s flow law and assume flow without basal slip (due to cold-ice
temperatures, see below); we also include changes in overburden pressures arising from the presence of
supraglacial debris and incorporate a parabolic shape factor [Hooke, 2005] to approximate the effect of lateral
drag by valley sidewalls:
h¼
1
v ðn þ 1Þ
ρd nþ1 ðnþ1Þ
ρd
d
d
þ
ρi
ρi
2AðSf ρi g sin αÞn
(1)
where h is the depth of the deforming ice body, v is the surface velocity, n is the flow law exponent
(3), ρi is the average bulk density of ice (900 kg m3), ρd is the average density of the supraglacial
debris (1800 kg m3), g is the acceleration of gravity, d is the supraglacial debris thickness, Sf is a
parabolic shape factor, α is the surface slope, and A is the ice-flow hardness parameter. The
supraglacial debris thickness (ranging from 0 to 0.5 m) is derived from a linear fit to measured debris
depths (Figure S7a) along the centerline of each glacier and the parabolic shape factor is estimated via
interpolating the values from Paterson [1994] using valley geometry derived from GIS measurements
and longitudinal and transverse GPR-derived ice thicknesses estimates (see sections 4.4.4.1 and 4.5).
The value for the ice hardness parameter, A, was initially estimated based on the average measured
and modeled ice temperatures (section 5.2) and then iteratively tuned to investigate the goodness of
fit to GPR ice thickness measurements.
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This simple steady state analytical model is successful in broadly duplicating GPR-derived ice thicknesses
along the central flow line of each glacier; results confirm maximum ice thicknesses of ~125 m and ~105 m
for the studied portions of Mullins and Friedman Glaciers, respectively (Figure S12 in the supporting
information). The best fit for both glaciers occurs with a constant ice hardness flow parameter, A, equal
to ~7.7 × 1025 s1 Pa1/3. This value of A is slightly higher than what would be expected of cold polar ice
[e.g., Paterson, 1994] but is within the range of values calculated for Mullins Glacier (1.27 × 1025 to
2.34 × 1024 s1 Pa1/3) via seismic depth modeling by Shean et al. [2007] and broadly consistent with the
value used for 23°C ice in central Beacon Valley (1.36 × 1024 s1 Pa1/3) by Rignot et al. [2002].
5.2. Modeled Thermal Regime
Although the thicknesses and environmental settings of both Mullins and Friedman Glaciers are consistent with
cold-based glacial thermal regimes, we explicitly calculate their current maximum basal temperature in order to
determine the magnitude of temperature change necessary to reach polythermal basal conditions. Our
estimates for the basal thermal regime of both Mullins and Friedman Glaciers are derived from ice thickness
estimates (GPR data/modeling) and shallow ice temperatures derived from borehole measurements and
modeling. First, we consider a vertical column of ice in one dimension, ignoring strain heating and advection,
whereby the heat transfer equation reduces to a simple linear form with a slope of G/k, where G is the
geothermal flux and k is the average thermal conductivity [Paterson, 1994]. When this linear approximation is fit
to borehole temperatures at 10 m and 30 m at MCI-09-03, we arrive at an estimated geothermal flux of
105 mW/m2 and a maximum estimated basal ice temperature of 17.4°C at 125 m (the measured maximum
depth of Mullins Glacier) (Figures 5 and S13 in the supporting information). Employing the full range of reported
values of G for the MDV region (68–105 mW/m2), we derive an envelope for basal ice temperatures at 125 m
depth of between 17 and 19°C. We also modeled the thermal regime using the thermal solution within the
thermomechanically coupled Ice Sheet Systems Model [Larour et al., 2012] in order to investigate any potential
heat contributions from strain heating/advection. As expected for these thin and slow-moving glaciers, we
found contributions from these sources to be negligible (Figure S13). The only model scenarios that produce
basal ice temperatures near the pressure melting point are those that require mean annual ice surface
temperatures well above modern values (e.g., an increase in mean annual atmospheric temperature by ~17°C)
and unrealistically high values of G (approaching ~180 mW/m2).
5.3. Modeled Rockfall Runout Distance
In an effort to understand better the potential sources for englacial and supraglacial debris on Mullins and
Friedman Glaciers, we utilized the numerical simulations within CRSP-3d (Colorado Rockfall Simulation
Program) [Andrew et al., 2012; Haramy et al., 2013] to model run out distances for falling clasts sourced along
the headwall of Mullins Glacier. At the core of CRSP-3d is a discrete element method (DEM) simulation of rock
velocity (translational and rotation in three dimensions) and contact forces between rock and slope [Andrew
et al., 2012; Bartingale et al., 2009; Leine et al., 2013]. Falling rocks are defined by shape, size, and density, and
slope material is parameterized by physical surface roughness and a hardness coefficient—a value that
subsumes both the coefficient of restitution and a dampening coefficient; see additional model
parameterizations and run scenarios in Figure 15.
The modeled rockfall run out distances using the CRSP-3d model are provided in Figure 15. As expected,
rockfall from isolated rock outcrops yields narrow, linear tracks of concentrated debris, rather than
widespread layers across the glacier surface. In addition, the modern headwall geometry and ice slope
characteristics favor long run out distances, with deposition primarily within the lower accumulation area
(~70%) and the upper ablation area (30%). Debris trapped in the accumulation area generally includes the
smaller clast sizes representative of most Mullins supraglacial debris (<2.0 m). This debris would become
buried beneath subsequent snow and ice to undergo rotation and transportation along internal flow lines
[Berthling et al., 2000]; those rocks that initially come to rest out on the ablation area include the largest clast
sizes (>2.0 m); these clasts would travel solely as supraglacial debris (Figures 16).
6. Discussion
We evaluate potential origins for englacial and supraglacial debris within Mullins and Friedman Glaciers. We
begin with supraglacial debris, for which a candidate hypothesis for its origin has already been proposed
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Figure 15. Simulated rockfall events from Mullins Valley headwall using the Colorado Rockfall Simulation Program (CRSP3d; see section 5.3). Model topography is derived from the digital elevation model (DEM) from Schenk et al. [2004] and a
5 m DEM derived from a stereo pair of GeoEye01scenes (upper headwall only). Material domains were assigned within the
model to match visual inspections of imagery and ground-truthed in field. Surface roughness values for ice and hard snow
were estimated to be 0.1–0.5 cm. The hardness coefficient for packed snow was based on iterative tuning of the model to
match observed rockfall. Model rocks include a range of idealized shapes (sphere, plate, and cylinder) and sizes. (a) Model
geometry and representative rockfall run out paths for 1 m spherical rocks and n = 300. Individual paths are delineated by
colored streaks down the headwall with color representing the kinetic energy contained in the rock at each location on the
fall path. (b) Histograms showing the horizontal stopping distance away from the point of origin for individual rockfall
events: (top) the combined run out distance statistics for all clasts (0.5 m, 1 m, 2 m, and 3 m diameters; n = 50 for each size
class); (bottom) run out distance for falling rocks of 0.5 m and 3 m diameters (n = 50). The results show that a partitioning in
run out distance is based on clast size (larger rocks travel greater distances) but that the majority of total rockfall debris will
be deposited in the lower accumulation area. (c) Spatial domain of material types within the model.
[Kowalewski et al., 2011, and references therein]; we then evaluate potential origins for the enigmatic layers of
inclined rocky debris (IDL). We end with an assessment of whether cold-based, debris-covered glaciers
containing distinct englacial debris bands can yield reliable records for long-term environmental change.
6.1. The Origin of Supraglacial Debris on Mullins and Freidman Glaciers
The prevailing hypothesis for the origin of supraglacial debris on Mullins and Friedman Glaciers is rockfall
sourced from rock cliffs at valley heads [Kowalewski et al., 2011; Rignot et al., 2002; Shean and Marchant, 2010]
(Figure 16). This model is entirely consistent with data presented above. However, the relative percentage of
rockfall that lands directly onto ice ablation areas (and travels solely as supraglacial debris) versus the
percentage that lands in ice accumulation areas—and thereby includes an additional component of englacial
transport [Kowalewski et al., 2011; Marchant et al., 2013]—is currently unknown. The distinction is important, not
only for understanding transport pathways but also for assessing potential origins for the inclined layers of
concentrated englacial debris (IDL) found throughout Mullins and Friedman Glaciers (see section 6.2).
Based on results from GPR data, the measured thickness of supraglacial debris, and the concentration of
debris observed in shallow ice cores, it appears that most of the supraglacial debris on Mullins and Friedman
Glaciers could have formed directly from sublimation of ice containing rocky debris. For example, if we (1)
accept 1% as the maximum concentration of dispersed englacial debris (by volume) in Regions 1 and 2 of
Mullins and Friedman Glaciers (excluding IDL, see below), (2) assume that ice sublimation decreases from
~101 mm a1 to ~102 mm a1 across Region 2 [Kowalewski et al., 2011]; that horizontal ice flow for Region
2 is ≤40 mm a1 [Rignot et al., 2002]; and (3) consider surface erosion is essentially negligible over the
time scales considered here [Morgan et al., 2010a, 2010b], then nearly 70% of the observed thickness of
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Figure 16. Diagrammatic illustration showing how two end-member scenarios of rockfall type/magnitude might impact
the distribution of englacial and surficial debris on Mullins and Friedman Glaciers (cross section and plan view). (a)
Rockfall is low magnitude and stochastically distributed in space and time. This rockfall debris (grey dots) is initially distributed across the surface of the accumulation area. Following burial beneath fresh snow and ice, clasts travel as scattered
englacial debris (top: black dots) and reemerge down valley at the base of a thickening layer of surficial debris (top: grey
stipple; bottom: black dots). (b) Rockfall is dominated by low-frequency high-magnitude events. At time t = 1, a highmagnitude rockfall deposit is emplaced (top: grey line; bottom: grey dots). At time t = 2, the rockfall debris has traveled
englacially to intermediate depth; however, debris that initially landed on the ablation area has remained at the ice surface.
At time t = 3, much of the debris traveling englacially has reemerged at the buried ice surface (bottom: black dots).
supraglacial debris in Region 2 can be attributed to sublimation of underlying ice with 1% scattered englacial
debris (by volume). If correct, then the remaining 30% of the supraglacial debris would likely come from
rockfall deposited directly on top of ice ablation areas and from the addition of concentrated englacial debris
located within inclined debris layers (IDL, see below).
To test this hypothesis, we turn to the results of the modeled expected rockfall runout distances (section 5.3),
which show, to a first order, that given the present-day ice surface slopes and headwall geometries, the
majority of rockfall derived from exposed rock outcrops at valley heads would come to rest on ice
accumulation areas and subsequently follow englacial transport pathways; only exceptionally large clasts,
≥2–3 m in diameter (rare on Mullins and Friedman Glaciers), would typically be deposited in ice ablation areas
(Figure 15). The subtle increase in the thickness of supraglacial debris across lobate ridges and steps on
Mullins and Friedman Glaciers (a 10–20% increase across arcuate surface discontinuities, ASD), as well as the
consistent change in grain size across these boundaries, suggests that sublimation of ice in IDL adds an
important but relatively minor contribution to supraglacial debris. Taken together, the results suggest that
the vast majority of supraglacial debris (~70% of its thickness) is derived from sublimation of ice with
scattered englacial debris, with the remaining ~30% derived from a combination of rockfall events that result
in some component of long-distance runout and, locally, by the addition of rocky debris within IDL.
6.2. The Origin of Inclined Debris Layers (IDL) Within Mullins and Friedman Glaciers
Inclined layers of englacial debris are relatively uncommon features in alpine glaciers but have been
observed consistently in debris-covered glaciers and traditional rock glaciers [Berthling et al., 2000;
Degenhardt, 2009; Degenhardt and Giardino, 2003; Degenhardt et al., 2003; Isaksen et al., 2000; Monnier et al.,
2008]. Although purported origins vary, most posited origins require some type of basal entrainment
mechanism followed by active transport toward the ice surface. Two related models applicable to Mullins and
Friedman Glaciers are shearing and thrusting of subglacial, ice-rich permafrost [e.g., Chinn and Dillon, 1987;
Fukui et al., 2008; Strelin and Sone, 1998] and basal entrainment beneath cold-based ice [e.g., Atkins, 2013;
Atkins et al., 2002; Cuffey et al., 2000a; Mager et al., 2009]. A third model, which involves the concentration of
rockfall debris in ice accumulation areas, followed by subsequent burial beneath snow and ice, and englacial
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Figure 17. Diagrammatic illustration showing the origin of inclined debris layers (IDL) following the climate-hollow-lag (CHL)
model. (a) Sketches of hypothetical accumulation, differential ablation, and mass balance during times of net accumulation
and net ablation. (b) Diagrams depicting formation of IDL via alternating climate regimes. The initial condition (time t = 1) calls
for a debris-covered glacier with ice exposed in the upper ablation area and surface debris gradually thickening down valley.
At time t = 2, a climate shift favoring ablation (right) results in differential ice loss, with significant loss near the valley headwall
(unprotected ice) and minimal loss beneath thick debris cover. As ice loss progresses, a spoon-shaped hollow forms near the
headwall that ultimately is capped by a lag of rocky debris. At time t = 3, the climate shifts to a net accumulation phase again
and the spoon-shaped hollow is in-filled with snow and glacier ice. The surface lag of time step 2 has now become an inclined
debris layer (IDL). At time t = 4, the climate has again shifted to the net ablation phase and a new ablation hollow forms near
the head of the glacier. The IDL formed in the previous climate cycle continues to deform and flow down valley.
transport, is also possible [e.g., Berthling et al., 2000; Degenhardt, 2009; Degenhardt and Giardino, 2003;
Degenhardt et al., 2003; Isaksen et al., 2000; Monnier et al., 2008]. The merits of all three models as potential
origins for IDL in Mullins and Friedman valleys are presented below.
6.2.1. Origin of IDL by Basal Shearing and Thrusting of Preexisting, Ice-Rich Debris
Although shearing and thrusting of subglacial debris has been postulated to explain the origin of debris
layers in polythermal debris-covered glaciers [e.g., Chinn and Dillon, 1987; Fukui et al., 2008; Strelin and Sone,
1998], it has yet to be applied to cold-based alpine glaciers like Mullins and Friedman Glaciers (e.g., with basal
temperatures <16°C). The basic assumption of this model is that heterogeneous ice movement, especially
across areas of enhanced/reduced basal flow typically associated with thermal transitions, may induce shear
stresses and thrusts [Nye, 1952] that could potentially provide both an entrainment mechanism and transport
vector for uplifting subglacial debris [Alley et al., 1997; Chinn and Dillon, 1987; Fukui et al., 2008; Glasser et al.,
2003; Hambrey et al., 1999; Hubbard et al., 2004]. In addition to differential basal flow, the model requires the
persistent availability of ice-rich subglacial debris. For Mullins and Friedman Glaciers, the model must also
account for punctuated entrainment that is uniform in both space and time across two separate glaciers. Our
GPR data suggest that although ice-rich debris may underlie glacier ice in Regions 2–4, such debris is most
probably absent from Region 1. Rather, GPR data from Region 1—where IDL originate—show profiles with
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smooth, basal returns, most plausibly interpreted as coherent bedrock. Where these smooth reflectors are
most visible in GPR profiles (Figures 6 and 9), they appear to lie well below the onset of mapped IDL,
suggesting individual IDL are most probably sourced above the bed. Additionally, given the inherent
complexities in bedrock topography, ice thicknesses, and ice surface slopes, it is unlikely that punctuated
basal shearing would be consistent across two separate glaciers. One possibility is that past episodes of
climate warming may have temporarily yielded polythermal ice conditions, leading to episodic basal
entrainment and thrust formation [e.g., Hubbard et al., 2004]. However, there is little evidence for recent
climate warming of the magnitude to produce polythermal conditions (~17°C increase for the current ice
thickness; Figure S13); rather, local evidence calls for enduring hyperarid, cold-desert conditions for the
last several million years and extreme landscape stability over this duration [Marchant et al., 2002, 2013;
Schaefer et al., 2000; Sugden et al., 1995b]. In support of this assertion, we highlight the observation that clasts
within IDL do not exhibit striations or faceting that would be consistent with basal entrainment. For these
reasons, we consider basal shearing and thrusting of preexisting, ice-rich debris to be an unlikely origin for
IDL within Mullins and Friedman Glaciers.
6.2.2. Origin of IDL Through Basal Entrainment Beneath Cold-Based Ice and Subsequent Transport
Along Shear Planes
Recent studies have shown that minor erosion is possible beneath cold-based ice [e.g., Atkins, 2013; Atkins et al.,
2002; Cuffey et al., 2000a]. Atkins et al. [2002] and Atkins [2013] have demonstrated convincingly that, in places,
cold-based ice can rotate and transport loose debris across bedrock, producing gouges and scratches and
depositing thin gravelly till. Others have postulated far more significant erosion and basal entrainment; however,
on close inspection, erosion and entrainment in these latter cases has been localized at the margins of thawed
patches in polythermal ice, rather than beneath the broad areas of cold-based ice [e.g., Chinn and Dillon, 1987].
Finally, Cuffey et al. [2000a] modeled erosion rates beneath a fully cold-based alpine glacier in the MDV (Meserve
Glacier, with basal temperatures everywhere ≤ 17°C) and calculated rates on the order of 107 m Ma1.
In further consideration of a basal-entrainment mechanism for IDL within Mullins and Friedman Glaciers, we
first discuss whether the requisite erosion rates and subsequent entrainment of rocky debris in IDL would
greatly exceed postulated and calculated rates of erosion beneath typical cold-based ice [Cuffey et al., 2000a].
Assuming a conservative ice flow rate of 35 mm a1 [Rignot et al., 2002], an average IDL thickness of 2 m, and
an average debris concentration in IDL of ~25% by volume, the volume flux of IDL exiting Region 1 is on
the order of 0.0175 m3 m2 s1. Averaged across the entire bed of Region 1 (required by the widespread
distribution of IDL’s observed in transverse GPR profiles), this flux would require aglacier-wide erosion rate on
the order of 4 × 105 m/yr, roughly 2 orders of magnitude greater than that reported for the cold-based
Meserve Glacier [e.g., Cuffey et al., 2000a]. Furthermore, because entrainment would likely be localized and
not uniform across all of Region 1, this calculation represents a conservative minimum erosion rate and the
localized entrainment rate would be much higher.
In addition, we note that the well-documented, tell-tale signs for basal entrainment beneath cold-based ice are
consistently lacking from IDL in Mullins and Friedman Glaciers. Missing indicators for debris entrainment
beneath cold-based ice include the following: the presence of ubiquitous amber-colored ice [Cuffey et al.,
2000a; Mager et al., 2009; Samyn et al., 2005; Sleewaegen et al., 2003; Tison et al., 1993]; a several fold increase in
the concentration of total dissolved solids [Cuffey et al., 2000b; Holdsworth, 1974; Holdsworth and Bull, 1970;
Mager et al., 2009]; an increase in silt and dispersed fine debris (as much as 5%) [Fitzsimons, 2003; Mager et al.,
2009]; striated or glacially molded clasts; or the appearance of “muddy clots” within otherwise fine-grained
debris [Tison et al., 1993]. None of these characteristics are present within IDL of Mullins and Friedman Glaciers.
In fact, apart from the noted increase in the abundance of gravel-and-cobble-sized clasts in mapped IDL’s, the
associated ice is identical in isotopic composition, bulk-dissolved solute content, and visible appearance to that
observed elsewhere throughout Regions 1–4 of Mullins and Friedman Glaciers.
Finally, and perhaps most significantly, we note that typical cold-based alpine (nondebris-covered) glaciers in
the MDV lack the englacial debris bands that so characterize the surface and internal structure of Mullins and
Friedman Glaciers. If cold-based ice were capable of significant basal entrainment followed by subsequent
upward movement of debris along thrust planes [Alley et al., 1997; Fukui et al., 2008; Glasser et al., 2003;
Hambrey et al., 1999], then we would expect to find similar IDL within other cold-based alpine glaciers
throughout the Dry Valleys, which is not the case [Chinn, 1980, 1985; Fountain et al., 2006]. Some englacial
debris bands do occur in glaciers of the MDV, but these bands are invariably found near glacier snouts [Chinn,
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1989], as they are within cold-based glaciers outside the MDV region [Fitzsimons, 2003; Fitzsimons and
Colhoun, 1995; Hooke, 1973; Hooke and Hudleston, 1978], and are most likely related to overriding of terminal
debris aprons [Hooke, 1973; Shaw, 1977]. Further, the spatial distribution of IDL along the length of Mullins
and Friedman Glaciers requires that basal erosion and entrainment must be punctuated in time, with similar
pulses of roughly equal-magnitude erosion and transport for both Mullins and Friedman Glaciers, a scenario
we consider unlikely. Thus, on the basis of the above discussion, we do not consider basal entrainment
beneath cold-based ice a likely origin for mapped IDL.
Other potential basal entrainment mechanisms for IDL are equally problematic. These might include the
trapping of debris within basal crevasses [Bennett et al., 1996; Ensminger et al., 2001], the folding of debrisladen basal ice [e.g., Glasser et al., 1998], and, as noted above, the overriding of terminal debris aprons [Hooke,
1973; Shaw, 1977]. However, these processes are typically only active within thin terminal regions of
polythermal glaciers and therefore are not likely applicable to Mullins and Friedman Glaciers.
Notwithstanding the above, we wish to emphasize that the data presented do not preclude the possibility of
localized shear along established IDL; however, they are first formed. In fact, the discrete rheological contrasts
presented by debris-laden IDL [Arenson et al., 2002; Colaprete and Jakosky, 1998] in otherwise clean glacier ice
would provide preferred planes for differential shear under longitudinal compressive stress [Hooke et al.,
1972]. Thus, shear and enhanced vertical transport may occur along some of the IDL’s in Mullins and Friedman
Glaciers [e.g., Shean and Marchant, 2010], but this does not imply that they formed by this mechanism.
Minor differential shear would provide one explanation for very subtle changes in horizontal ice surface
velocity observed in InSAR data [Rignot et al., 2002]. Future analysis of ice fabrics may help illuminate the
nature and location of any potential shear [e.g., Alley, 1988] associated with IDL.
6.2.3. Rockfall Origin for IDL Within Mullins and Friedman Glaciers
As noted above, the vast majority of alpine glaciers in the MDV lacks supraglacial debris; the few
debris-covered glaciers that do exist form only in the lee of cliffs, principally below cliffs of Ferrar
Dolerite and Beacon Supergoup sandstones [Marchant et al., 2013]. Moreover, all clean-ice alpine
glaciers (e.g., those without supraglacial debris) in the MDV appear to lack IDL [Chinn, 1980, 1985;
Fountain et al., 2006; Mager et al., 2009, and references therein]. Taken together, these simple
observations suggest that the origin of supraglacial debris and IDL’s may be related and both may
require rockfall from valley headwalls.
In consideration of this model, we note that changes in rockfall—and subsequent burial of debris in ice
accumulation areas—has often been attributed as a causal mechanism for internal debris layers and
multiunit structures within debris-covered glaciers and rock glaciers outside of Antarctica [Berthling et al.,
2000; Degenhardt, 2009; Degenhardt and Giardino, 2003; Degenhardt et al., 2003; Isaksen et al., 2000; Monnier
et al., 2008]. We postulate two end-member scenarios for a rockfall-centric origin for IDL in Mullins and
Friedman Glaciers. The first involves high-magnitude rockfall events that would place significant debris on ice
accumulation areas; such an event, or series of events closely spaced in time (e.g., <~10 years apart) would
yield a geologically instantaneous rocky layer that could ultimately be buried beneath subsequent snow and
ice, producing inclined layers of englacial debris [e.g., Berthling et al., 2000] (Figure 16b). We refer to this
model as the large-magnitude rockfall model (LMR). In contrast, the second scenario relies on maintaining
low rates of rockfall (similar to today’s) but instead requires marked changes in ice accumulation and ablation.
A decrease in net ice accumulation would favor lowering of exposed glacier ice near valley headwalls (e.g., ice
not protected beneath debris) but negligible lowering for glacier ice buried beneath supraglacial debris. This
differential ablation would produce spoon-shaped hollows near valley headwalls [Ackert, 1998; Humlum,
1996; Johnson and Lacasse, 1988; Potter, 1972; Whalley, 1979] that—owing to the continued sublimation of ice
with scattered englacial debris and persistent low-amplitude rockfall events—would ultimately become
covered with a uniform and widespread rocky lag. The overall grain size, sorting, and spatial arrangement of
clasts within IDL, where observed directly in hand-dug ice trenches, tends to mimic the distribution and
stacking of large cobbles on top of small pebbles observed in the upper ablation areas of Mullins and
Freidman glaciers. This similarity hints at the possibility that IDL may have developed in ice accumulation
areas through slow sublimation processes similar to those currently operating in the ablation areas today.
Renewed net-ice accumulation would bury these lags, initiate englacial transport, and ultimately produce
englacial debris bands (IDL). We refer to this model as the climate-hollow-lag model (CHL).
MACKAY ET AL.
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In assessing the LMR model, we note that the strong correlation between the number and spacing of IDL
(and associated ASD) across Mullins and Friedman Glaciers implies regional forcing, rather than the linkage
of stochastic rockfall at the heads of Mullins and Friedman valleys. On the other hand, simultaneous
detachment of large volumes of rockfall could occur if forced by regional seismic activity [e.g., Reznichenko
et al., 2011]. If seismicity was the driving force for IDL, then one would expect to see widespread evidence
for recent and catastrophic rockfall beneath all steep cliffs throughout the western Dry Valleys region, which
is not the case [Denton et al., 1993; Marchant et al., 1996; Putkonen et al., 2012; Sugden et al., 1995a; Swanger
and Marchant, 2007]. Instead, current models for landscape evolution suggest incredible landscape stability,
with extremely low rates of sporadic rockfall; the region possesses some of the lowest rates of downslope
movement on Earth [Marchant et al., 2013; Putkonen et al., 2012] and minimal bedrock erosion, the latter
averaging <10 cm/Ma [Summerfield et al., 1999a, 1999b]. Even if synchronous, large-magnitude rockfall
events were possible; modeling studies from Okura et al. [2000] suggest that for bedrock geometries
comparable to those of Mullins and Friedman valleys, large-magnitude rockfall events would result in longdistance transport; in this case, clasts would likely travel well out onto ice ablation areas and essentially
bypass ice accumulation areas, thereby preventing development of IDL (see also Figure 16b). Finally, the
uniform thickness and widespread areal extent of IDL, especially as observed in transverse GPR profiles, is
difficult to reconcile via large-magnitude rockfall. Deposits from large rockfall events are typically elongated
in plan view, with high length-to-width ratios, and are inconsistent with the observed distribution (across the
entire glacier width) and uniform thickness of mapped IDL.
One possibility is that regional climate change may induce marked changes in ice accumulation and
ablation as well as an increase in the overall frequency rockfall events [Gruber and Haeberli, 2007; Gruber
et al., 2004]. Together, these changes could produce widespread lags in upper ablation areas that, with
renewed ice accumulation could form IDL. Linkages between climate change and rockfall frequency have
been noted for cliffs in subpolar and temperate regions, where changes in the rate of meltwater
development have been implicated with climate warming [e.g., Gruber and Haeberli, 2007; Gruber et al.,
2004; Huggel et al., 2012; Ravanel and Deline, 2011]. However, such linkages have not been documented for
cliffs in the Transantarctic Mountains, where mean annual temperatures are well below zero Celsius
(< 23°C in this case) and where the necessity of frequent rockfall events appear at odds with
geomorphic data that call for long-term landscape stability throughout the region [Marchant et al.,
2013; Putkonen et al., 2012]. On the basis of the above data, we cannot rule out a combination of
increased rockfall rate and decreased ice accumulation as co-drivers in the development of IDL, but
the most parsimonius explanation is that rockfall rates remain low and relatively constant—as
suggested by allied geomorphic data—and that lags form predominantly by reduced ice
accumulation and the development of spoon-shaped hollows.
6.3. The Potential for Debris-Covered Glaciers in the MDV to Register Long-Term Climate Change
Figure 17 presents a diagrammatic illustration of our preferred model for the development of IDL, the
Climate-Hollow-Lab model (CHL). As shown in Figure 17b, during episodes of reduced ice accumulation
(t = 2), ablation is most pronounced at the surface of exposed (unprotected) ice in the upper portions of the
glacier [Ackert, 1998; Konrad et al., 1999; Kowalewski et al., 2011; Potter et al., 1998; Scherler et al., 2011]. Such
differential ablation forms the localized topographic depression (termed “ablation hollows”) near glacier
heads; most of the protected portion of the glacier is little affected and continues to slowly flow down valley.
Due to reversed ice surface slopes in the vicinity of the down valley portion of the hollow (Figure 17b),
localized ice flow in this region may flow very slowly up valley, partially infilling and elevating the nescient
hollow. However, given the low ice temperatures, thin ice conditions, and relatively minor changes in ice
surface elevations, any reverse flow into the hollows would be minimal compared to ice loss from
sublimation. As the ablation hollow forms, its thickening mantle of surface debris (the evolving lag from
sublimation of ice with rocky debris and continued, minor rockfall) forms an insulating layer that halts
runaway ice loss. Following a climate transition favoring net ice accumulation (t = 3, Figure 17b), the ablation
hollow and overlying surface debris are buried. This now-buried debris (IDL) is subsequently deformed and
advected down valley [e.g., Berthling et al., 2000; Degenhardt, 2009; Degenhardt and Giardino, 2003; Goodsell
et al., 2005; Hooke and Hudleston, 1978; Isaksen et al., 2000; Vaughan et al., 1999]. The processes repeats,
forming IDL at similar times in both Mullins and Friedman Glaciers (t = 4).
MACKAY ET AL.
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Journal of Geophysical Research: Earth Surface
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Although the duration and magnitude of climate forcing required to produce discrete IDL is unknown, the
extremely slow horizontal ice-flow velocities recorded in InSAR [Rignot et al., 2002], as well as the reported
low sublimation and high potential for long-term ice preservation [Kowalewski et al., 2011], suggest causation
by relatively low-frequency climate perturbations, perhaps on the order of ~104 years.
Our preferred model for the development of IDL builds upon aspects of previous models that invoked some
combination of climate change and partially overriding flow lobes in talus-derived rock glaciers [Degenhardt,
2009; Monnier et al., 2008], variations in debris input or ice accumulation/ablation [Barsch, 1996; Olyphant,
1987], and alternating rockfall regimes driven by fluctuating headwall geometries [Ackert, 1998]. In our
departure from these previous models, we conclude that the salient features of Mullins and Friedman
Glaciers that are necessary (if not required) to form IDL (and associated ASD) via the CHL model mechanism
include the following: (1) persistent cold-based ice conditions, (2) rockfall source areas in the headwall region,
(3) increasingly thick supraglacial debris in the down valley direction, and (4) low overall ablation rates.
Therefore, other debris-covered glaciers that meet these criteria may also contain records of long-term
climate change. However, we caution that not all surface lineations/ridges on Mullins and Friedman glaciers
(or similar debris-covered glaciers elsewhere) can be considered ASD, as defined here, or must be associated
with climate perturbations. These locally discontinuous surface ridges that lack associated IDL’s and do not
show the stepwise changes in surface geomorphology that define ASD (e.g., the secondary lineations shown
in Figure 3) are most likely related to compressive flow [e.g., Burger et al., 1999; Kääb and Weber, 2004],
perhaps exacerbated by changes in flow velocity related to the alternate development of ablation hollows
near valley heads, followed by renewed ice accumulation and enhanced longitudinal flow.
7. Conclusions
We used ground-penetrating radar, ice core analyses, geomorphic mapping, and numerical modeling to
investigate the surface and internal structure of two cold-based, debris-covered glaciers situated within
neighboring valleys in the cold-and-dry Quartermain Mountains of southern Victoria Land, Antarctica.
Friedman Glacier, ~100 m thick at its deepest point near the valley headwall, thins steadily along its length
before terminating at 3.8 km. Mullins Glacier thins from ~125 m thick near its headwall toward a local
minimum of ~22 m at 3.4 km distance; the thickness of Mullins Glacier ice at the termination of GPR coverage
(5.2 km down valley) could not be determined due to incomplete penetration of the radar signal.
A simple application of a modified Glen’s flow law reasonably predicted measured ice thicknesses and flow
velocities calculated from InSAR [Rignot et al., 2002]. Both Mullins and Friedman Glaciers are cold based and
frozen to their beds.
Both glaciers are overlain by a layer of dry supraglacial debris (8–75 cm thick), each of which is marked with
near-identical patterns of arcuate surface ridges and steps (ASD); each ASD, in turn, displays uniform changes in
bulk-sediment grain size, meter-scale surface topography, and the geometry and maturity of surface polygons.
Results from 24 km of ground-penetrating radar (80, 200, and 400 MHz antennas) show that each ASD is
underlain by a distinct layer of rocky englacial debris, inclined up glacier, and containing as much as 35%
debris by volume. Between these inclined debris layers, ice in Mullins and Friedman Glaciers is nearly devoid
of englacial debris <1% by volume; the most likely source for the scattered englacial debris (<1% by volume)
is sporadic rockfall at valley headwalls.
The similarity in number and pattern of englacial debris bands and corresponding ASD across both glaciers,
along with model results and observations that call for negligible basal entrainment as a source for englacial
debris, is best explained by episodic environmental change at valley headwalls. We envision three distinct
components to this model: (1) continued low-amplitude rockfall (similar to today’s rate) to produce scattered
englacial debris; (2) a reduction in net ice accumulation that favors a marked reduction in the ice surface
elevation of (unprotected) ice in former ice accumulation areas but negligible ice surface lowering beneath
debris-covered (protected) portions of each glacier; ultimately, this differential ablation produces a spoonshaped hollow near valley headwalls that, eventually, is capped with rocky lags produced from continued ice
sublimation and low-amplitude rockfall events; (3) renewed net ice accumulation buries these lags, initiates
englacial transport, and ultimately produces the observed englacial debris bands and accompanying arcuate
surface discontinuities.
MACKAY ET AL.
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Journal of Geophysical Research: Earth Surface
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If the model is correct, the implication is that cold-based, debris-covered glaciers with distinct englacial
debris bands and corresponding changes in surface morphology may record environmental change. Given
the propensity for long-lived debris-covered glaciers in Antarctica, this opens up the possibility that a subset
of debris-covered glaciers could serve as reliable proxies for long-term climate change, potentially covering
the last several hundred thousand years duration.
Acknowledgments
Funding for this research was provided
by NSF Graduate Research Fellowship
awarded to S.L. Mackay and NSF Polar
Programs grants NSF ANT-0944702 and
NSF ANT-0739700 to D.R. Marchant.
Logistical support for this project in
Antarctica was provided by the U.S.
National Science Foundation through
the U.S. Antarctic Program. We are very
grateful to Geophysical Survey Systems,
Inc. (GSSI) for technical assistance, support, and temporary loan of the MLF
antenna. We thank Alistar Hayden,
Michael Dyonisius, Greg Wissink, and
Andrew Knott for tireless assistance
with ground-penetrating radar in the
field; we also thank Michael Bender and
Audrey Yao for their excellent discussions on the Mullins Glacier system.
Finally, we thank Ole Humlum, an
anonymous reviewer, and the Editor,
Bryn Hubbard, for their detailed and
very helpful comments on an earlier
version of this paper. Data are available
on the NASA global change master
directory: (http://gcmd.nasa.gov/getdif.
htm?[GCMD]BU-054-001).
MACKAY ET AL.
References
Ackert, R. (1998), A rock glacier/debris-covered glacier system at Galena Creek, Absaroka Mountains, Wyoming, Geografiska Ann. Ser. A,
80(3–4), 267–276, doi:10.1111/j.0435-3676.1998.00042.x.
Alley, R., K. Cuffey, E. Evenson, J. Strasser, D. Lawson, and G. Larson (1997), How glaciers entrain and transport basal sediment: Physical
constraints, Quat. Sci. Rev., 16(9), 1017–1038, doi:10.1016/S0277-3791(97)00034-6.
Alley, R. B. (1988), Fabrics in polar ice sheets: Development and prediction, Science, 240(4851), 493–495, doi:10.1126/science.240.4851.493.
Andrew, R., H. Hume, R. Bartingale, A. Rock, and R. Zhang (2012), CRSP-3D user’s manual Colorado rockfall simulation program, F. H.
Administration, Rep. FHWA-CFL/TD-12-007. [Available at http://www.cflhd.gov/programs/techDevelopment/geotech/CRSP-3D/documents/01_CRSP-3D_Users_Manual_entire_report.pdf.]
Arcone, S. A. (1996), High resolution of glacial ice stratigraphy: A ground-penetrating radar study of Pegasus Runway, McMurdo Station,
Antarctica, Geophysics, 61(6), 1653–1663, doi:10.1190/1.1444084.
Arcone, S. A., M. L. Prentice, and A. J. Delaney (2002), Stratigraphic profiling with ground-penetrating radar in permafrost: A review of
possible analogs for Mars, J. Geophys. Res., 107(E11), 18-11–18-14, doi:10.1029/2002JE001906.
Arenson, L., M. Hoelzle, and S. Springman (2002), Borehole deformation measurements and internal structure of some rock glaciers in
Switzerland, Permafrost Periglacial Process., 13(2), 117–135, doi:10.1002/ppp.414.
Atkins, C. B. (2013), Geomorphological evidence of cold-based glacier activity in South Victoria Land, Antarctica, Geol. Soc. London Spec. Publ.,
381, 299–318, doi:10.1144/SP381.18.
Atkins, C. B., P. Barrett, and S. Hicock (2002), Cold glaciers erode and deposit: Evidence from Allan Hills, Antarctica, Geology, 30(7), 659,
doi:10.1130/0091-7613(2002)030<0659:CGEADE>2.0.CO;2.
Augustinus, P., and M. Selby (1990), Rock slope development in McMurdo Oasis, Antarctica, and implications for interpretations of glacial
history, Geografiska Ann. Ser. A, 72A, 55–62, doi:10.2307/521237.
Barsch, D. (1996), Rock Glaciers: Indicators for the Present and Gormer Geoecology in High Mountain Environments, Springer, Berlin,
doi:10.1007/978-3-642-80093-1.
Barsch, D., and M. Jakob (1998), Mass transport by active rock glaciers in the Khumbu Himalaya, Geomorphology, 26(1–3), 215–222,
doi:10.1016/S0169-555X(98)00060-9.
Bartingale, R., J. D. Higgins, R. Andrew, A. Rock, and R. Zhang (2009), Colorado rockfall simulation program version 5.0, in 60th Highway
Geology Symposium, pp. 189–200, Buffalo, New York.
Bennett, M. R., M. J. Hambrey, D. Huddart, and J. Ghieene (1996), The formation of a geometrical ridge network by the surge type glacier
Kongsvegen, Svalbard, J. Quat. Sci., 11(6), 437–449.
Berthling, I., B. Etzelmüller, K. Isaksen, and J. Sollid (2000), Rock glaciers on Prins Karls Forland. II: GPR soundings and the development of
internal structures, Permafrost Periglacial Process., 11(4), 357–369, doi:10.1002/1099-1530(200012)11:4<357::AID-PPP366>3.0.CO;2-6.
Burger, K., J. J. Degenhardt, and J. Giardino (1999), Engineering geomorphology of rock glaciers, Geomorphology, 31(1–4), 93–132,
doi:10.1016/S0169-555X(99)00074-4.
Campbell, I. B., and G. G. Claridge (2006), Permafrost properties, patterns and processes in the Transantarctic Mountains region, Permafrost
Periglacial Process., 17(3), 215–232, doi:10.1002/ppp.557.
Chinn, T. (1980), Glacier balances in the dry valleys area, Victoria Land, Antarctica, in Proceedings of the Riederalp Workshop, vol. 125,
pp. 237–247, IAHS-AISH, Wallingford, U. K.
Chinn, T. (1985), Structure and equilibrium of the Dry Valleys glaciers, N. Z. Antarct. Rec., 6, 73–88.
Chinn, T. (1989), Single folds at the margins of dry-based glaciers as indicators of a glacial advance, Ann. Glaciol., 12, 23–30.
Chinn, T., and A. Dillon (1987), Observation on a debris-covered polar glacier ’Whisky Glacier’, James Ross Island, Antarctic Peninsula,
Antarctica, J. Glaciol., 33(115), 300–310.
Clark, D., E. Steig, N. Potter, J. Fitzpatrick, A. Updike, and G. Clark (1996), Old ice in rock glaciers may provide long-term climate records, Eos
Trans. AGU, 77(23), 217–221, doi:10.1029/96EO00149.
Clark, D., E. Steig, N. Potter Jr., and A. Gillespie (1998), Genetic variability of rock glaciers, Geografiska Ann. Ser. A, 80, 175–182, doi:10.1111/
j.0435-3676.1998.00035.x.
Colaprete, A., and B. Jakosky (1998), Ice flow and rock glaciers on Mars, J. Geophys. Res., 103(E3), 5897–5909, doi:10.1029/97JE03371.
Craig, H. (1961), Isotopic variations in meteoric waters, Science, 133(3465), 1702–1703, doi:10.1126/science.133.3465.1702.
Cuffey, K., H. Conway, A. Gades, B. Hallet, R. Lorrain, J. Severinghaus, E. Steig, B. Vaughn, and J. White (2000a), Entrainment at cold glacier
beds, Geology, 28(4), 351–354, doi:10.1130/0091-7613(2000)28<351:EACGB>2.0.CO;2.
Cuffey, K., H. Conway, A. Gades, B. Hallet, C. Raymond, and S. Whitlow (2000b), Deformation properties of subfreezing glacier ice: Role of
crystal size, chemical impurities, and rock particles inferred from in situ measurements, J. Geophys. Res., 105(B12), 27,895–27,915,
doi:10.1029/2000JB900271.
Dansgaard, W. (1964), Stable isotopes in precipitation, Tellus, 16(4), 436–468, doi:10.1111/j.2153-3490.1964.tb00181.x.
Degenhardt, J. J. (2009), Development of tongue-shaped and multilobate rock glaciers in alpine environments—Interpretations from
ground penetrating radar surveys, Geomorphology, 109(3–4), 94–107, doi:10.1016/j.geomorph.2009.02.020.
Degenhardt, J. J., and J. R. Giardino (2003), Subsurface investigation of a rock glacier using ground-penetrating radar: Implications for
locating stored water on Mars, J. Geophys. Res., 108(E4), 8036, doi:10.1029/2002JE001888.
Degenhardt, J. J., J. R. Giardino, and M. B. Junck (2003), GPR survey of a lobate rock glacier in Yankee Boy Basin, Colorado, USA, Geol. Soc.
London Spec. Publ., 211(1), 167, doi:10.1144/GSL.SP.2001.211.01.14.
Denton, G. H., D. E. Sugden, D. R. Marchant, B. Hall, and T. Wilch (1993), East Antarctic ice sheet sensitivity to Pliocene climatic change from a
Dry Valleys perspective, Geografiska Ann. Ser. A, 75A, 155–204, doi:10.2307/521200.
De Pascale, G., W. Pollard, and K. Williams (2008), Geophysical mapping of ground ice using a combination of capacitive coupled resistivity
and ground-penetrating radar, Northwest Territories, Canada, J. Geophys. Res., 113, F02S90, doi:10.1029/2006JF000585.
©2014. American Geophysical Union. All Rights Reserved.
2537
Journal of Geophysical Research: Earth Surface
10.1002/2014JF003178
Doran, P. T., R. Wharton, W. Lyons, D. D. Marais, and D. Andersen (2000), Sedimentology and geochemistry of a perennially ice-covered
epishelf lake in Bunger Hills Oasis, East Antarctica, Antarct. Sci., 12(02), 131–140, doi:10.1017/S0954102000000171.
Ensminger, S. L., R. B. Alley, E. B. Evenson, D. E. Lawson, and G. J. Larson (2001), Basal-crevasse-fill origin of laminated debris bands at
Matanuska Glacier, Alaska, USA, J. Glaciol., 47(158), 412–422, doi:10.3189/172756501781832007.
Fitzpatrick, J., E. Steig, D. Clark, N. Potter Jr., A. Updike, G. Clark, and S. Jacobsen (1995), Rock glaciers as archives of old glacier ice: Evidence
from an ice core at Galena Creek, Wyoming, Eos Trans. AGU, 76, F214.
Fitzsimons, S. J. (2003), Ice-marginal terrestrial landsystems: Polar continental glacier margins, in Glacial Landsystems, pp. 89–110, Arnold, London.
Fitzsimons, S. J., and E. A. Colhoun (1995), Form, structure and stability of the margin of the Antarctic ice sheet, Vestfold Hills and Bunger Hills,
East Antarctica, Antarct. Sci., 7, 171–171, doi:10.1017/S095410209500023X.
Fountain, A. G., M. Tranter, T. H. Nylen, K. J. Lewis, and D. R. Mueller (2004), Evolution of cryoconite holes and their contribution to meltwater
runoff from glaciers in the McMurdo Dry Valleys, Antarctica, J. Glaciol., 50(168), 35–45, doi:10.3189/172756504781830312.
Fountain, A. G., T. H. Nylen, K. L. MacCLUNE, and G. L. Dana (2006), Glacier mass balances (1993–2001), Taylor Valley, McMurdo Dry Valleys,
Antarctica, J. Glaciol., 52(178), 451–462, doi:10.3189/172756506781828511.
Fountain, A. G., T. Nylen, A. Monaghan, H. Basagic, and D. Bromwich (2009), Snow in the McMurdo Dry Valleys, Antarctica, Int. J. Climatol., 30,
633–642, doi:10.1002/joc.1933.
Frauenfelder, R., and A. Kääb (2000), Towards a palaeoclimatic model of rock-glacier formation in the Swiss Alps, Ann. Glaciol., 31(1), 281–286,
doi:10.3189/172756400781820264.
Fukui, K., T. Sone, J. Strelin, C. Torielli, J. Mori, and Y. Fujii (2008), Dynamics and GPR stratigraphy of a polar rock glacier on James Ross Island,
Antarctic Peninsula, J. Glaciol., 54(186), 445–451, doi:10.3189/002214308785836940.
Giardino, J., J. Shroder, and J. Vitek (1987), Rock Glaciers, Australia, Allen & Unwin.
Glasser, N. F., M. J. Hambrey, K. R. Crawford, M. R. Bennett, and D. Huddart (1998), The structural glaciology of Kongsvegen, Svalbard, and its
role in landform genesis, J. Glaciol., 44(146), 136–148.
Glasser, N. F., M. J. Hambrey, J. L. Etienne, P. Jansson, and R. Pettersson (2003), The origin and significance of debris-charged ridges at the
surface of Storglaciaren, northern Sweden, Geografiska Ann. Ser. A, 85(2), 127–147, doi:10.1111/1468-0459.00194.
Goodsell, B., M. Hambrey, and N. F. Glasser (2005), Debris transport in a temperate valley glacier: Haut Glacier d’Arolla, Valais, Switzerland,
J. Glaciol., 51(172), 139–146, doi:10.3189/172756505781829647.
Gooseff, M. N., W. B. Lyons, D. M. McKnight, B. H. Vaughn, A. G. Fountain, and C. Dowling (2006), A stable isotopic investigation of a polar
desert hydrologic system, McMurdo Dry Valleys, Antarctica, Arct. Antarct. Alp. Res., 38(1), 60–71, doi:10.1657/1523-0430(2006)038[0060:
ASIIOA]2.0.CO;2.
Green, J., B. Koci, and J. Kyne (2007), Koci drill for drilling in ice, sand and rock: Drill requirements, design, performance, difficulties, Ann.
Glaciol., 47(1), 105–108, doi:10.3189/172756407786857677.
Gruber, S., and W. Haeberli (2007), Permafrost in steep bedrock slopes and its temperature-related destabilization following climate change,
J. Geophys. Res., 112, F02S18, doi:10.1029/2006JF000547.
Gruber, S., M. Hoelzle, and W. Haeberli (2004), Permafrost thaw and destabilization of Alpine rock walls in the hot summer of 2003, Geophys.
Res. Lett., 31, L13504, doi:10.1029/2004GL020051.
Haeberli, W. (1985), Creep of mountain permafrost: Internal structure and flow of alpine rock glaciers, Mitt. der Versuchsanst. fur Wasserbau,
Hydrol. und Glaziologie an der ETH Zurich, 77, 5–142.
Haeberli, W., and D. Vonder Mühll (1996), On the characteristics and possible origins of ice in rock glacier permafrost, Z. Fur Geomorphol., 104, 43–57.
Haeberli, W., M. Hoelzle, A. Kääb, F. Keller, D. Vonder Mühll, and S. Wagner (1998), Ten years after drilling through the permafrost of the active
rock glacier Murtel, Eastern Swiss Alps: Answered questions and new perspectives, paper presented at 7th International Conference on
Permafrost, Yellowknife, Canada.
Haeberli, W., B. Hallet, L. Arenson, R. Elconin, O. Humlum, A. Kääb, V. Kaufmann, B. Ladanyi, N. Matsuoka, and S. Springman (2006), Permafrost
creep and rock glacier dynamics, Permafrost Periglacial Process., 17(3), 189–214.
Hambrey, M., M. Bennett, J. Dowdeswell, N. Glasser, and D. Huddart (1999), Debris entrainment and transfer in polythermal valley glaciers,
J. Glaciol., 45(149), 69–86.
Haramy, K. Y., R. Zhang, and A. Rock (2013), CRSP-3D application for remediating a rockfall at Yosemite National Park, in Geo-Congress 2013:
Stability and Performance of Slopes and Embankments III, edited by C. Meehan, pp. 592–603, Am. Soc. of Civ. Eng., San Diego, Calif.,
doi:10.1061/9780784412787.062.
Hauck, C., and C. Kneisel (2008), Applied Geophysics in Periglacial Environments, 240 pp., Cambridge Univ. Press, Cambridge, doi:10.1017/
CBO9780511535628.
Hausmann, H., K. Krainer, E. Bruckl, and W. Mostler (2007), Internal structure and ice content of Reichenkar rock glacier (Stubai Alps, Austria)
assessed by geophysical investigations, Permafrost Periglacial Process., 18(4), 351–367, doi:10.1002/ppp.601.
Head, J. W., D. R. Marchant, J. L. Dickson, A. M. Kress, and D. M. Baker (2010), Northern mid-latitude glaciation in the Late Amazonian period of
Mars: Criteria for the recognition of debris-covered glacier and valley glacier landsystem deposits, Earth Planet. Sci. Lett., 294(3), 306–320,
doi:10.1016/j.epsl.2009.06.041.
Hendy, C. H. (2000), The role of polar lake ice as a filter for glacial lacustrine sediments, Geografiska Ann. Ser. A, 82(2–3), 271–274, doi:10.1111/
j.0435-3676.2000.00125.x.
Holdsworth, G. (1974), Meserve Glacier Wright Valley, Antarctica: Part I. Basal processes, Rep. 37, pp. 1–104 Institute of Polar Studies.
Holdsworth, G., and C. Bull (1970), The Flow Law of Cold Ice: Investigations on Meserve Glacier, Antarctica, IASH Publ., vol. 86, pp. 204–216,
Hanover, N. H, Antarctic Claciological Exploration (ISAGE).
Hooke, R. L. (1973), Flow near the margin of the Barnes Ice Cap, and the development of ice-cored moraines, Geol. Soc. Am. Bull., 84(12),
3929–3948, doi:10.1130/0016-7606(1973)84<3929:FNTMOT>2.0.CO;2.
Hooke, R. L. (2005), Principles of Glacier Mechanics, 2nd ed., Cambridge Univ. Press, Cambridge, U. K.
Hooke, R. L., and P. J. Hudleston (1978), Origin of foliation in glaciers, J. Glaciol., 20(83), 285–299.
Hooke, R. L., B. Dahlin, and M. Kauper (1972), Creep of ice containing dispersed fine sand, J. Glaciol., 11, 327–336.
Hubbard, B., and N. Glasser (2005), Field Techniques in Glaciology and Glacial Geomorphology, 400 pp., John Wiley, Chichester, U. K.
Hubbard, B., N. Glasser, M. Hambrey, and J. Etienne (2004), A sedimentological and isotopic study of the origin of supraglacial debris bands:
Kongsfjorden, Svalbard, J. Glaciol., 50(169), 157–170, doi:10.3189/172756504781830114.
Huggel, C., J. J. Clague, and O. Korup (2012), Is climate change responsible for changing landslide activity in high mountains?, Earth Surf.
Processes Landforms, 37(1), 77–91, doi:10.1002/esp.2223.
Humlum, O. (1996), Origin of rock glaciers: Observations from Mellemfjord, Disko Island, central West Greenland, Permafrost Periglacial
Process., 7(4), 361–380, doi:10.1002/(SICI)1099-1530(199610)7:4<361::AID-PPP227>3.0.CO;2–4.
MACKAY ET AL.
©2014. American Geophysical Union. All Rights Reserved.
2538
Journal of Geophysical Research: Earth Surface
10.1002/2014JF003178
Humlum, O. (1998), The climatic significance of rock glaciers, Permafrost Periglacial Process., 9(4), 375–395, doi:10.1002/(SICI)1099-1530
(199810/12)9:4<375::AID-PPP301>3.0.CO;2–0.
Humlum, O. (2000), The geomorphic significance of rock glaciers: Estimates of rock glacier debris volumes and headwall recession rates in
West Greenland, Geomorphology, 35(1–2), 41–67, doi:10.1016/S0169-555X(00)00022-2.
Isaksen, K., R. Odegard, T. Eiken, and J. Sollid (2000), Composition, flow and development of two tongue-shaped rock glaciers in the permafrost of Svalbard, Permafrost Periglacial Process., 11(3), 241–257, doi:10.1002/1099-1530(200007/09)11:3<241::AID-PPP358>3.0.CO;2-A.
Johnson, P. (1984), Rock glacier formation by high-magnitude low-frequency slope processes in the southwest Yukon, Ann. Assoc. Am.
Geogr., 74(3), 408–419, doi:10.1111/j.1467-8306.1984.tb01463.x.
Johnson, P., and D. Lacasse (1988), Rock glaciers of the Dalton Range, Kluane Ranges, south-west Yukon Territory, Canada, J. Glaciol., 34(118),
327–332.
Kääb, A., and M. Weber (2004), Development of transverse ridges on rock glaciers: Field measurements and laboratory experiments,
Permafrost Periglacial Process., 15(4), 379–391, doi:10.1002/ppp.506.
Kääb, A., V. Kaufmann, R. Ladstädter, and T. Eiken (2003), Rock glacier dynamics: Implications from high-resolution measurements of surface
velocity fields, paper presented at 8th International Conference on Permafrost, Zurich, Switzerland.
Konrad, S., and N. Humphrey (2000), Steady-state flow model of debris-covered glaciers (rock glaciers), in Debris Covered Glaciers, edited, pp.
255–266, IAHS Publication, Wallingford, U. K.
Konrad, S., N. Humphrey, E. Steig, D. Clark, N. Potter, and W. Pfeffer (1999), Rock glacier dynamics and paleoclimatic implications, Geology,
27(12), 1131–1134, doi:10.1130/0091-7613(1999)027<1131:RGDAPI>2.3.CO;2.
Kowalewski, D., D. R. Marchant, J. S. Levy, and J. W. W. Head (2006), Quantifying low rates of summertime sublimation for buried glacier ice in
Beacon Valley, Antarctica, Antarct. Sci., 18(3), 421–428, doi:10.1017/S0954102006000460.
Kowalewski, D., D. R. Marchant, K. M. Swanger, and J. W. Head (2011), Modeling vapor diffusion within cold and dry supraglacial tills of
Antarctica: Implications for the preservation of ancient ice, Geomorphology, 126(1), 159–173, doi:10.1016/j.geomorph.2010.11.001.
Kowalewski, D., D. R. Marchant, J. W. Head, and D. W. Jackson (2012), A 2D model for characterising first-order variability in sublimation of
Buried Glacier Ice, Antarctica: Assessing the influence of polygon troughs, desert pavements and shallow subsurface salts, Permafrost
Periglacial Process., 23(1), 1–14, doi:10.1002/ppp.731.
Lancaster, N. (2002), Flux of eolian sediment in the McMurdo Dry Valleys, Antarctica: A preliminary assessment, Arct. Antarct. Alp. Res., 34,
318–323, doi:10.2307/1552490.
Larour, E., H. Seroussi, M. Morlighem, and E. Rignot (2012), Continental scale, high order, high spatial resolution, ice sheet modeling using the
Ice Sheet System Model (ISSM), J. Geophys. Res., 117, F01022, doi:10.1029/2011JF002140.
Lehmann, F., and A. G. Green (2000), Topographic migration of georadar data: Implications for acquisition and processing, Geophysics, 65(3),
836, doi:10.1190/1.1444781.
Lehmann, F., D. V. Miihll, M. v. d. Veen, P. Wild, and A. G. Green (1998), True topographic 2-D migration of georadar data, in Symposium on
the Application of Geophysics to Engineering and Environmental Problems, edited by R. S. Bell, M. H. Powers, and T. Larson, pp. 107–114,
Environ. and Eng. Geophys. Soc., Chicago, Ill.
Leine, R., A. Schweizer, M. Christen, J. Glover, P. Bartelt, and W. Gerber (2013), Simulation of rockfall trajectories with consideration of rock
shape, Multibody Syst. Dyn., 32(2), 241–271, doi:10.1007/s11044-013-9393-4.
Levy, J. S., D. R. Marchant, and J. W. W. Head (2006), Distribution and origin of patterned ground on Mullins Valley debris-covered glacier,
Antarctica: The roles of ice flow and sublimation, Antarct. Sci., 18(3), 385–397, doi:10.1017/S0954102006000435.
Lundstrom, S., A. McCafferty, and J. Coe (1993), Photogrammetric analysis of 1984–89 surface altitude change of the partially debris-covered
Eliot Glacier, Mount Hood, Oregon, USA, Ann. Glaciol., 17, 167–170.
Mager, S., S. Fitzsimons, R. Frew, D. Samyn, and R. Lorrain (2009), Composition and origin of amber ice and its influence on the behaviour of
cold glaciers in the McMurdo Dry Valleys, Antarctica, J. Glaciol., 55(190), 363–372, doi:10.3189/002214309788608688.
Marchant, D. R., and J. W. Head (2007), Antarctic dry valleys: Microclimate zonation, variable geomorphic processes, and implications for
assessing climate change on Mars, Icarus, 192(1), 187–222, doi:10.1016/j.icarus.2007.06.018.
Marchant, D. R., G. H. Denton, C. Swisher III, and N. Potter Jr. (1996), Late Cenozoic Antarctic paleoclimate reconstructed from volcanic ashes in
the Dry Valleys region of southern Victoria Land, Bull. Geol. Soc. Am., 108(2), 181, doi:10.1130/0016-7606(1996)108<0181:LCAPRF>2.3.CO;2.
Marchant, D. R., A. R. Lewis, W. M. Phillips, E. J. Moore, R. A. Souchez, G. H. Denton, D. E. Sugden, N. Potter, and G. P. Landis (2002), Formation of
patterned ground and sublimation till over Miocene glacier ice in Beacon Valley, southern Victoria Land, Antarctica, Geol. Soc. Am. Bull.,
114(6), 718–730, doi:10.1130/0016-7606(2002)114<0718:FOPGAS>2.0.CO;2.
Marchant, D. R., S. Mackay, J. W. Head, and D. Kowalewski (2010), Documenting microclimate variation and the distribution of englacial
debris in Mullins Glacier, Antarctica: Implications for the origin, flow, and modification of LDA and LVF on Mars, paper presented at 41st
Lunar and Planeteray Sciences, Houston, Tex.
Marchant, D. R., S. L. Mackay, J. L. Lamp, A. T. Hayden, and J. W. Head (2013), A review of geomorphic processes and landforms in the Dry
Valleys of southern Victoria Land: Implications for evaluating climate change and ice-sheet stability, Geol. Soc. London Spec. Publ., 381,
319–352, doi:10.1144/SP381.10.
Martin, E., and W. Whalley (1987), Rock glaciers: Part 1: Rock glacier morphology: Classification and distribution, Prog. Phys. Geogr., 11(2), 260.
Mihalcea, C., C. Mayer, G. Diolaiuti, A. Lambrecht, C. Smiraglia, and G. Tartari (2006), Ice ablation and meteorological conditions on the debriscovered area of Baltoro glacier, Karakoram, Pakistan, Ann. Glaciol., 43, 292, doi:10.3189/172756406781812104.
Monnier, S., C. Camerlynck, and F. Rejiba (2008), Ground penetrating radar survey and stratigraphic interpretation of the Plan du Lac rock
glaciers, Vanoise Massif, northern French Alps, Permafrost Periglacial Process., 19(1), 19–30, doi:10.1002/ppp.610.
Morgan, D., J. Putkonen, G. Balco, and J. Stone (2010a), Degradation of glacial deposits quantified with cosmogenic nuclides, Quartermain
Mountains, Antarctica, Earth Surf. Processes Landforms, 36(2), 217–228, doi:10.1002/esp.2039.
Morgan, D., J. Putkonen, G. Balco, and J. Stone (2010b), Quantifying regolith erosion rates with cosmogenic nuclides 10Be and 26Al in the
McMurdo Dry Valleys, Antarctica, J. Geophys. Res., 115, F03037, doi:10.1029/2009JF001443.
Nye, J. (1952), The mechanics of glacier flow, J. Glaciol., 2(12), 82–93.
Okura, Y., H. Kitahara, T. Sammori, and A. Kawanami (2000), The effects of rockfall volume on runout distance, Eng. Geol., 58(2), 109–124,
doi:10.1016/S0013-7952(00)00049-1.
Olyphant, G. (1987), Rock glacier response to abrupt changes in talus production, in Rock Glaciers, pp. 55–64, Allen and Unwin, London.
Ostrem, G. (1959), Ice melting under a thin layer of moraine, and the existence of ice cores in moraine ridges, Geogr. Ann., 41(4), 228–230.
Paterson, W. S. B. (1994), The Physics of Glaciers, 3rd ed., 480 pp., Pergamon, Oxford.
Plewes, L. A., and B. Hubbard (2001), A review of the use of radio-echo sounding in glaciology, Prog. Phys. Geogr., 25(2), 203, doi:10.1177/
030913330102500203.
MACKAY ET AL.
©2014. American Geophysical Union. All Rights Reserved.
2539
Journal of Geophysical Research: Earth Surface
10.1002/2014JF003178
Potter, N. (1972), Ice-cored rock glacier, Galena Creek, northern Absaroka Mountains, Wyoming, Geol. Soc. Am. Bull., 83(10), 3025–3058.
Potter, N., E. Steig, D. Clark, M. Speece, G. Clark, and A. Updike (1998), Galena Creek rock glacier revisited: New observations on an old
controversy, Geografiska Ann. Ser. A, 80A, 251–265, doi:10.1111/j.0435-3676.1998.00041.x.
Putkonen, J., D. J. Morgan, and G. Balco (2012), Regolith transport quantified by braking block, McMurdo Dry Valleys, Antarctica,
Geomorphology, 155, 80–87, doi:10.1016/j.geomorph.2011.12.010.
Ravanel, L., and P. Deline (2011), Climate influence on rockfalls in high-Alpine steep rockwalls: The north side of the Aiguilles de Chamonix
(Mont Blanc massif) since the end of the ‘Little Ice Age’, Holocene, 21(2), 357–365, doi:10.1177/0959683610374887.
Reynolds, J. M. (1997), An Introduction to Applied and Environmental Geophysics, 796 pp., John Wiley, New York.
Reznichenko, N. V., T. R. Davies, and D. J. Alexander (2011), Effects of rock avalanches on glacier behaviour and moraine formation,
Geomorphology, 132(3–4), doi:10.1016/j.geomorph.2011.05.019.
Rignot, E., B. Hallet, and A. Fountain (2002), Rock glacier surface motion in Beacon Valley, Antarctica, from synthetic-aperture radar
interferometry, Geophys. Res. Lett., 29(12), 1607, doi:10.1029/2001GL013494.
Salvatore, M., J. Mustard, J. W. Head, D. R. Marchant, and M. Wyatt (2013), Characterization of spectral and geochemical variability within the
Ferrar Dolerite of the McMurdo Dry Valleys, Antarctica: Weathering, alteration, and magmatic processes, Antarct. Sci., 26(1), 49–68,
doi:10.1017/S0954102013000254.
Samyn, D., A. Svensson, S. J. Fitzsimons, and R. D. Lorrain (2005), Ice crystal properties of amber ice and strain enhancement at the base of
cold Antarctic glaciers, Ann. Glaciol., 40(1), 185–190, doi:10.3189/172756405781813618.
Schaefer, J. M., H. Baur, G. H. Denton, S. Ivy-Ochs, D. R. Marchant, C. Schluchter, and R. Wieler (2000), The oldest ice on Earth in Beacon Valley,
Antarctica: New evidence from surface exposure dating, Earth Planet. Sci. Lett., 179(1), 91–99.
Schenk, T., B. Csatho, Y. Ahn, T. Yoon, S. W. Shin, and K. I. Huh (2004), DEM generation from the Antarctic LiDAR data: Site report. [Available at
http://usarc.usgs.gov/lidar/lidar_pdfs/Site_reports_v5.pdf.]
Scherler, D., B. Bookhagen, and M. R. Strecker (2011), Spatially variable response of Himalayan glaciers to climate change affected by debris
cover, Nat. Geosci., 4(3), 156–159, doi:10.1038/ngeo1068.
Shaw, J. (1977), Till body morphology and structure related to glacier flow, Boreas, 6(2), 189–201, doi:10.1111/j.1502-3885.1977.tb00348.x.
Shean, D. E., and D. R. Marchant (2010), Seismic and GPR surveys of Mullins Glacier, McMurdo Dry Valleys, Antarctica: Ice thickness, internal
structure and implications for surface ridge formation, J. Glaciol., 56(195), 48, doi:10.3189/002214310791190901.
Shean, D. E., J. W. W. Head, and D. R. Marchant (2007), Shallow seismic surveys and ice thickness estimates of the Mullins Valley debriscovered glacier, McMurdo Dry Valleys, Antarctica, Antarct. Sci., 19(4), 485–496, doi:10.1017/S0954102007000624.
Shroder, J., M. Bishop, L. Copland, and V. Sloan (2000), Debris-covered glaciers and rock glaciers in the Nanga Parbat Himalaya, Pakistan,
Geografiska Ann. Ser. A, 82(1), 17–31, doi:10.1111/j.0435-3676.2000.00108.x.
Sleewaegen, S., D. Samyn, S. J. Fitzsimons, and R. D. Lorrain (2003), Equifinality of basal ice facies from an Antarctic cold-based glacier, Ann.
Glaciol., 37(1), 257–262, doi:10.3189/172756403781815708.
Souchez, R., and R. Lorrain (2006), The environmental significance of deuterium excess in meteoric and non-meteoric antarctic ice, in Glacier
Science and Environmental Change, edited by P. G. Knight, pp. 179–184, doi:10.1002/9780470750636.ch35.
Steig, E., J. Fitzpatrick, N. Potter Jr., and D. Clark (1998), The geochemical record in rock glaciers, Geografiska Ann. Ser. A, 80(3), 277–286,
doi:10.1111/j.0435-3676.1998.00043.x.
Strelin, J., and T. Sone (1998), Rock glaciers on James Ross Island, Antarctica, paper presented at 7th International Permafrost Conference,
Collection Nordicana, Univ. Laval, Yellowknife, Canada.
Sugden, D. E., G. H. Denton, and D. R. Marchant (1995a), Landscape evolution of the Dry Valleys, Transantarctic Mountains: Tectonic implications, J. Geophys. Res., 100, 9949–9967, doi:10.1029/94JB02875.
Sugden, D. E., D. R. Marchant, N. Potter, R. A. Souchez, G. H. Denton, C. C. Swisher, and J. L. Tison (1995b), Preservation of miocene glacier ice in
East Antarctica, Nature, 376(6539), 412–414, doi:10.1038/376412a0.
Summerfield, M., F. M. Stuart, H. A. P. Cockburn, D. E. Sugden, G. H. Denton, T. Dunai, and D. R. Marchant (1999a), Long-term rates of
denudation in the Dry Valleys, Transantarctic Mountains, southern Victoria Land, Antarctica based on in-situ-produced cosmogenic
Ne-21, Geomorphology, 27(1–2), 113–129, doi:10.1016/S0169-555X(98)00093-2.
Summerfield, M., D. E. Sugden, G. H. Denton, D. R. Marchant, H. Cockburn, and F. Stuart (1999b), Cosmogenic isotope data support previous
evidence of extremely low rates of denudation in the Dry Valleys region, southern Victoria Land, Antarctica, Geol. Soc. London Spec. Publ.,
162(1), 255, doi:10.1144/GSL.SP.1999.162.01.20.
Swanger, K. M., and D. R. Marchant (2007), Sensitivity of ice-cemented Antarctic soils to greenhouse-induced thawing: Are terrestrial archives
at risk?, Earth Planet. Sci. Lett., 259(3–4), 347–359, doi:10.1016/j.epsl.2007.04.046.
Swanger, K. M., D. R. Marchant, D. E. Kowalewski, and J. W. Head III (2010), Viscous flow lobes in central Taylor Valley, Antarctica: Origin as
remnant buried glacial ice, Geomorphology, 120(3), 174–185, doi:10.1016/j.geomorph.2010.03.024.
Swanger, K. M., D. R. Marchant, J. M. Schaefer, G. Winckler, and J. W. Head III (2011), Elevated East Antarctic outlet glaciers during warmerthan-present climates in southern Victoria Land, Global Planet. Change, 79(1), 61–72, doi:10.1016/j.gloplacha.2011.07.012.
Tison, J. L., J. R. Petit, J. M. Barnola, and W. Mahaney (1993), Debris entrainment at the ice-bedrock interface in sub-freezing temperature
conditions (Terre Adelie, Antarctica), J. Glaciol., 39(132), 303–315.
Trabant, P. K. (1984), Applied High Resolution Geophysical Methods, 265 pp., International Human Resources Development Corporation,
Boston, doi:10.1007/978-94-009-6493-8.
Vaughan, D. G., H. F. J. Corr, C. S. M. Doake, and D. Waddington (1999), Distortion of isochronous layers in ice revealed by ground-penetrating
radar, Nature, 398(6725), 323–326, doi:10.1038/18653.
Wade, F. A. (1945), The physical aspects of the Ross Shelf Ice, Proc. Am. Philos. Soc., 89(1), 160–173.
Welch, B. C., W. T. Pfeffer, J. T. Harper, and N. F. Humphrey (1998), Mapping subglacial surfaces of temperate valley glaciers by two-pass
migration of a radio-echo sounding survey, J. Glaciol., 44(146), 164–170.
Whalley, W. (1979), The relationship of glacier ice and rock glacier at Grubengletscher, Kanton Wallis, Switzerland, Geografiska Ann. Ser. A,
61(1), 49–61, doi:10.2307/520514.
Whalley, W., and F. Azizi (1994), Rheological models of active rock glaciers: Evaluation, critique and a possible test, Permafrost Periglacial
Process., 5(1), 37–51, doi:10.1002/ppp.3430050105.
Whalley, W., and F. Azizi (2003), Rock glaciers and protalus landforms: Analogous forms and ice sources on Earth and Mars, J. Geophys. Res.,
108(E4), 8032, doi:10.1029/2002JE001864.
Whalley, W., and H. Martin (1992), Rock glaciers: II. Models and mechanisms, Prog. Phys. Geogr., 16(2), 127, doi:10.1177/030913339201600201.
Yilmaz, O. (2001), Seismic Data Analysis: Processing, Inversion, and Interpretation of Seismic Data, 2027 pp., Society of Exploration Geophysics,
Tulsa, Okla., doi:10.1190/1.9781560801580.
MACKAY ET AL.
©2014. American Geophysical Union. All Rights Reserved.
2540