ncomms11897

ARTICLE
Received 11 Aug 2015 | Accepted 10 May 2016 | Published 10 Jun 2016
Massive subsurface ice formed by refreezing
of ice-shelf melt ponds
Bryn Hubbard1, Adrian Luckman2, David W. Ashmore1, Suzanne Bevan2, Bernd Kulessa2,
Peter Kuipers Munneke2, Morgane Philippe3, Daniela Jansen4, Adam Booth5, Heidi Sevestre6,
Jean-Louis Tison3, Martin O’Leary2& Ian Rutt2
Surface melt ponds form intermittently on several Antarctic ice shelves. Although implicated
in ice-shelf break up, the consequences of such ponding for ice formation and ice-shelf
structure have not been evaluated. Here we report the discovery of a massive subsurface ice
layer, at least 16 km across, several kilometres long and tens of metres deep, located in an
area of intense melting and intermittent ponding on Larsen C Ice Shelf, Antarctica. We
combine borehole optical televiewer logging and radar measurements with remote sensing
and firn modelling to investigate the layer, found to be B10 °C warmer and B170 kg mÀ 3
denser than anticipated in the absence of ponding and hitherto used in models of ice-shelf
fracture and flow. Surface ponding and ice layers such as the one we report are likely to form
on a wider range of Antarctic ice shelves in response to climatic warming in forthcoming
decades.
DOI: 10.1038/ncomms11897 OPEN
1 Centre for Glaciology, Department of Geography and Earth Sciences, Aberystwyth University, Aberystwyth SY23 3DB, UK. 2 Glaciology Group, Department
of Geography, Swansea University, Swansea SA2 8PP, UK. 3 Laboratoire de Glaciologie, De´partement Geósciences, Environnement et Socie´te´, Universite´ Libre
de Bruxelles, Bruxelles 1050, Belgium. 4 Alfred Wegener Institut, Helmholtz-Zentrum fu¨r Polar- und Meeresforschung D-27568, Bremerhaven, Germany.
5 School of Earth and Environment, University of Leeds, Leeds LS2 9JT, UK. 6 Department of Arctic Geology, University Centre in Svalbard, N-9171
Longyearbyen, Norway. Correspondence and requests for materials should be addressed to B.H. (email: byh@aber.ac.uk).
NATURE COMMUNICATIONS | 7:11897 | DOI: 10.1038/ncomms11897 | www.nature.com/naturecommunications 1

T
he B49,000 km2 Larsen C Ice Shelf (LCIS) is considered
susceptible to future collapse because of its exposure to
intense surface melting1 across the northern sector of the
Antarctic Peninsula (Fig. 1). Satellite and airborne radar data on
the LCIS indicate low firn-air content2, promoting the formation
of melt ponds and the potential for hydrofracturing3,4. Analysis of
remotely sensed images of Cabinet Inlet, located in the northwest
sector of the LCIS, reveals the presence during some summer
months of surface melt ponds that generally form in flow-parallel
troughs that are some tens of kilometres long and hundreds of
metres wide (Fig. 1). These ponds form in Cabinet Inlet as a result
of melting by fo¨hn winds that blow from the Graham Land
mountains and eastwards through the shelf’s northernmost-
fringing inlets5–7. Although appearing to generate substantial
surface melt, these winds do not persist for more than a few days
even through the summer, and ponds appear intermittently on
satellite images (Fig. 1). The additional stress and hydrofracturing
potential of such ponded surface water has been proposed as a
means of shelf destabilization8–10 and has been implicated in
the collapse of Larsen B Ice Shelf in 200211,12. However,
the implications of such ponding for the formation of new ice
and its influence on ice-shelf structure have not been reported.
In this study, we report the presence of a massive ice layer, at
least 16 km across, several kilometres long and tens of metres
deep, present beneath an area of intermittent pond formation on
LCIS, Antarctica. We combine field data with firn-modelling and
remotely sensed data to investigate the layer’s properties and
formation. The layer is found to be composed of two units: an
upper, solid ice unit formed largely from the continual refreezing
of ponded water; and a lower, infiltration ice unit formed
largely from the refreezing of meltwater that has percolated into
very dense firn. The layer is found to be B10 °C warmer and
B170 kg mÀ 3 denser than that which would have been present
in the absence of the influence of intense surface melting and
pond formation. The implications of the layer’s presence for
ice-shelf thickness estimates, flow and stability are explored.
Results
Borehole drilling and logging. In early austral summer
2014/2015, we drilled a B100-m long borehole into the flank of a
Cabinet Inlet trough indicated by satellite imagery to have
repeatedly hosted melt ponds over the past 15 years, but not since
2008/2009. Although the shelf was snow-covered at the time of
drilling, inspection of the wall of a 2-m deep pit revealed the
presence of both numerous ice layers within the snowpack and an
unusually thick ice layer at a depth of 2.0 m that prevented
continued excavation. The borehole was logged to a depth of
B97 m by optical televiewer (OPTV)13, providing a geometrically
accurate image of the complete borehole wall at a vertical and
lateral resolution of B1 mm. The resulting OPTV log (Fig. 2a)
contrasts starkly with those retrieved from other accumulating
b
a
Cabinet Inlet
Churchill
Peninsula
Cole
Peninsula
50 km 10 km
N
E
S
W
1 2 3
Figure 1 | Cabinet Inlet basemap and surface ponding. Location map of Cabinet Inlet study site on Larsen C Ice Shelf based on MODIS data from 3rd
December 2014. (a) Main figure location in Antarctica (red box). (b) Landsat expansion of Cabinet Inlet from 31st December 2001, showing surface
ponding (dark patches) and the locations of the borehole (yellow dot) and 200-MHz GPR transects (red lines) presented in Fig. 4.
ARTICLE NATURE COMMUNICATIONS | DOI: 10.1038/ncomms11897
2 NATURE COMMUNICATIONS | 7:11897 | DOI: 10.1038/ncomms11897 | www.nature.com/naturecommunications

ice-shelf and ice-sheet locations (for example, Fig. 2b), including
the interior of the Greenland Ice Sheet14 and an East Antarctic ice
shelf and ice rise15,16. In each of these other cases, OPTV log
luminosity decreases gradually with depth as surface snow
metamorphoses through firn to dense ice over depths of several
tens of metres. In contrast, our Cabinet Inlet OPTV log (Fig. 2a)
shows a sharp contact between high-luminosity surface snow or
firn and low-luminosity ice at a depth of only 2.9 m below the
ice-shelf surface. This ice extends to the 97-m deep base of
the log. Converting the luminosity of this OPTV log to
density reveals a mean density for this entire ice layer of
870 kg mÀ 3. Figure 2a also reveals a transition in ice type at a
depth of B45 m, with the overlying layer (named Unit 1)
hosting more frequent horizontal layers and being more dense
(mean ¼ 888 kg mÀ 3) than the underlying ice layer (Unit 2;
mean ¼ 854 kg mÀ 3), which is principally composed of bubbly
host ice containing coarse, irregularly dipping bubble-free ice
layers. The proximity of this generally massive ice layer to the
shelf surface, and the sharpness of its contact with the overlying
snowpack, preclude formation by the usual process of
compaction–metamorphism. Instead, we interpret this Unit 1
ice as having formed as a consequence of refreezing, following
periods of intense surface melting and intermittent pond
formation. The characteristics revealed by the OPTV image are
consistent with this unit (2.9 to B45 m) being largely bubble-free
pond ice formed from the refreezing of surface water ponded
during extended periods of intense, presumably summer, melting.
Here the fine-scale horizontal layering apparent in Fig. 2a likely
reflects the episodic nature of the process, which would involve
four general stages: (a) snow accumulation, (b) surface melting,
(c) meltwater infiltration into underlying snow, eventually
resulting in its saturation and pond formation, and (d) the
freezing of that meltwater layer to form pond ice, probably during
early austral autumn. In contrast, Unit 2 (approximately 45–97-m
depth) also contains layers of bubble-poor ice, but in this zone
they are typically decimetres thick, contorted and isolated within
host ice that is optically brighter (likely due to the presence of
reflective bubbles; Fig. 2a). We interpret this lower unit as ice
dominated by infiltration refreezing formed by meltwater
percolating into underlying firn. This still involves the influence
of intense surface melting but, in contrast to the Unit 1,
of generally insufficient magnitude to form a continuous
quasi-massive layer of largely bubble-free pond ice. These
physical characteristics and depth ranges of the two units we
identify are consistent with the upper (Unit 1) ‘pond ice’ forming
within the Cabinet Inlet region of pond formation and the lower
(Unit 2) ‘infiltration ice’ forming, and being inherited from,
up-flow of the region of pond formation (Fig. 1).
Firn modelling and satellite image analysis. We evaluate this
hypothesis for the recent past through a one-dimensional firn
densification and hydrology model17, driven by surface mass
fluxes and temperature data from the RACMO2.3 regional
climate model for Cabinet Inlet18. Model results (Fig. 3a,b)
indicate the occurrence of intermittent, but substantial, surface
melt events at the borehole location, consistent with Unit 1
forming from the refreezing of surface meltwater. As well as
predicting the presence of such a layer, the model predicts
(Fig. 3b) that the layer’s upper surface was, in summer
2014–2015, expected to be 2.9-m below the snow surface, and
that the overlying snowpack contains a substantial ice layer
between B1.9 and 2.2 m, agreeing with our direct observations
from snow-pit digging, the OPTV log (Fig. 2a) and ground-
penetrating radar (GPR) data (Fig. 4) addressed below.
Further, this analysis is supported by a time series of moderate
resolution imaging spectroradiometer (MODIS) satellite images
(Fig. 3c), indicating that melt ponds formed annually between
2001 and 2009, while none appears in any of the images available
between early 2009 and the time of fieldwork in late 2014. This
reconstruction matches very closely the RACMO-based firn
densification reconstruction (Fig. 3b) with ‘pond years’,
coinciding with periods of intense near-surface firn
densification (for example, 2005–2007).
Ground-penetrating radar. An approximation of the lateral
extent of the massive subsurface ice layer we report is provided
by the area known from satellite images to host melt ponds:
B60-km across-flow and B20-km along-flow (Fig. 1). However,
the precise degree to which this zone is underlain by massive
N E S W N –20 –10 0
DIR OPTVLCIS OPTV Temp. (°C)
N E S W N
Depth
(m)
LCIS density
(kg m–3
)
DIR density
(kg m–3
)
0
500
1,000
250
750
0
500
1,000
250
750
cba
0
10
20
30
40
50
60
70
80
90
Figure 2 | Cabinet Inlet borehole data. (a) OPTV log of the LCIS borehole.
This raw log is of the complete borehole wall, unrolled to progress
north—east—south—west—north from left to right. The log is classified
into two units, described and interpreted in the text: Unit 1 extends from a
depth of B2.9–45 m, and Unit 2 from a depth of B45 m to the base.
(b) OPTV log of a typical ice-shelf borehole unaffected by significant
melting (Derwael Ice Rise, Roi Baudouin Ice Shelf, Antarctica) showing
(a) the normal reduction in luminosity with depth, interpreted as
progressive densification from snow, through firn to ice, and (b) the usual
presence of regular horizontal planes closing up with depth, interpreted as
annual layering. Density profiles derived from OPTV luminosity (described
in Methods) are superimposed as red lines on both OPTV logs. (c) Englacial
temperatures (a) measured by borehole thermistor string (solid dots),
(b) that would normally be used in an ice-shelf model for this site (open
circles), and (c) predicted by our firn model at a depth of 11 m (cross).
NATURE COMMUNICATIONS | DOI: 10.1038/ncomms11897 ARTICLE

pond and/or infiltration ice, and the depth of such layers where
they are present beyond our borehole, cannot be determined from
temporally discrete satellite images. To evaluate this we carried
out GPR profiling at 200 MHz along three transects focused on
the borehole location (Fig. 1). The resulting radargrams (Fig. 4)
reveal the presence of numerous near-surface reflectors and one
substantial reflector at a depth that varies between B1 and B3 m
along the transects. Very little radar energy above the background
noise level was received from below this reflection. This absence
of signal return is consistent with the presence of a refrozen ice
layer that, although characterized by minor density stratification,
is physically and chemically uniform compared with firn layering
unmodified by meltwater. This main reflector is B2.9-m deep at
the borehole location, coincident with field-based digging, drilling
and OPTV logging, as well as with firn modelling (above).
We therefore infer that this strong GPR reflection represents the
upper surface of a spatially extensive ice layer that is both thick
and relatively homogeneous. Since meltwater ponds are confined
to surface troughs that persist in MODIS images for decades
within this region of Cabinet Inlet, it is unlikely that this ice layer
is ubiquitously composed of Unit 1 pond ice. Away from the
troughs, this near-surface reflector therefore probably indicates
the uppermost surface of an ice layer that is more similar to
Unit 2 infiltration ice, that is, still influenced by intense surface
melting and subsurface refreezing, but in these areas not actually
forming standing ponds. In the absence of further direct inves-
tigation, such as by ice coring or borehole analysis, it is not yet
possible to determine at high resolution the lateral variability of
the two units we identify. Nonetheless, our OPTV and GPR data
together indicate that a widespread ice layer extends at least
across all of our B16-km flow-orthogonal and B6-km flow-
parallel GPR study area. Our borehole OPTV log also indicates
that Units 1 and 2 combined extend to a depth of at least 97 m at
the location of our borehole. While it is highly likely that this
thickness—defined by the depth of the base of Unit 2—varies
spatially across Cabinet Inlet, it is not currently possible to specify
this distribution without further borehole and/or GPR data.
Discussion
The presence of the refrozen ice layer we report above affects the
physical properties of the LCIS in at least two ways. First,
the layer’s density is substantially higher than that of the snow and
firn that would otherwise have formed over this depth range by
standard compaction–metamorphism. For example, a recent LCIS
model19 used a density of B700 kgmÀ 3 for the shelf’s uppermost
B100 m, based on inverting seismic data recorded in the shelf’s
southern sector. In contrast, our OPTV-derived densities indicate a
measured density of B870 kgmÀ 3 over this depth range in the
Cabinet Inlet area, 24% higher. Such a density enhancement
influences calculations of the shelf’s thickness based on the surface
elevation data20. In this case, using an ice density of 917 kgmÀ 3, a
sea water density of 1,026 kgmÀ 3, a surface elevation of 63.5m
(Fig. 4) and the assumed firn density of 700 kgmÀ 3 for the
uppermost 97 m of the shelf yields a total shelf thickness of 382 m.
Repeating the calculation with our OPTV-reconstructed density of
870 kgm3 for the uppermost 97 m yields a total shelf thickness of
551 m. Although substantially thicker than that calculated from the
‘standard firn’ model, a thickness of 551 m is close to that of 564 m
reconstructed for the area by the Bedmap2 consortium21. This
close correspondence reflects the fact that the (altimetry-based)
Bedmap reconstructions include a correction to account for
enhanced densification, resulting from active surface melting on ice
shelves22. Our OPTV-based density reconstruction also yields a
firn-air content (the column-length equivalent accounted for by
material with a density lower than that of bubble-free glacial ice) of
5.0 m. While this value is substantially lower than the 23.0m that
would result from firn of density 700 kgmÀ 3, it is far closer to the
spatially distributed range of 0–4 m predicted for the area based on
recent combined analyses of remotely sensed surface elevation and
shelf thickness fields9,10. However, direct comparisons such as
these are somewhat confounded by our firn-air content of 5.0 m
being based on a single-point measurement (recorded, as noted
above, on the limb of an elongate trough hosting an ephemeral
surface pond), whereas the reconstructions based on remotely
sensed data integrate data over a coarser spatial field.
Second, as well as altering its density, the ice column will be
warmed by latent heat released by the freezing of ponded and
percolating meltwater. This effect is quantified by a thermistor
4
3
2
1
0
M:Aratio
Depth(m)
a
b
c
Year
0
2
4
6
8
10
Density (kg m–3
)
900
800
700
600
500
400
No ponds
Ponds
2014
2012
2010
2000
2008
2006
2004
2002
1990
1998
1996
1994
1992
Figure 3 | Firn model and remotely sensed outputs for Cabinet Inlet for
the period 1st January 1990 to the time of fieldwork on 1 st December
2014. (a) Ratio of melt to accumulation (M:A) predicted by the firn model
for each 12-month period (1st March–28th February). The horizontal
dashed line marks an M:A ratio of 1. (b) Predicted firn density plotted
against depth. The predicted massive shallow ice layer (dark blue) is never
more than B8 m below the shelf surface and extends to within B1 m of the
surface in years 1993–2001 and 2005–2009, both containing individual
years of high M:A ratio. (c) Classification of time series of summer MODIS
satellite images (available since 2001) of Cabinet Inlet according to the
presence (red bars) or absence (grey bars) of surface ponds.
50
–2,330,000
Easting
(m
)
–2,320,000
1,160,000
1,170,000
Northing (m)
60
70
Ele.(ma.s.l.)
1
2
3
Figure 4 | Cabinet Inlet radargrams. 200 MHz GPR profiles along the
three transects identified in Fig. 1, overlaid on a Landsat image of the region
from 31st December 2001 when surface melt ponds were present. The
borehole location is marked by the yellow dot.

string installed into the borehole following OPTV logging
(Fig. 2c). The mean annual temperature measured at a depth of
11 m was À 5.9 °C, while the mean annual surface temperature at
the site (which normally defines that at a depth of 10–15 m in the
absence of significant refreezing) is À 16.9 °C. Further, the
entire ice profile recorded by our thermistor string (Fig. 2c) is
warmer than would be expected without factoring in refreezing
of ponded meltwater. This effect is confirmed by our firn
model, which under-predicts englacial temperatures despite
accounting for heat released by the refreezing of percolating
meltwater. In this case, modelled pore-water refreezing
contributes B1.5 °C to the firn, yielding a predicted temperature
at 11 m of À 15.4 °C (Fig. 2c). This is still almost 10 °C colder
than our measured firn temperature at that depth, demonstrating
that the refreezing of ponded meltwater (not included in the
model) provides significant, hitherto unconsidered, englacial
heating in this region. As well as being denser, this ice is
therefore also warmer than that which would otherwise be used in
numerical models of the flow of the LCIS19.
Replacing on the order of 10 Â 10 Â 0.1 km of standard firn
with relatively warm and dense ice will exert some influence on
ice shelf flow and stability. However, evaluating this influence at
the ice-shelf scale is not straightforward. In the first instance, the
warmer ice layer we identify will be less viscous than colder ice,
and this may go some way to accounting for the anomalously
high-rate factor that has in the past been necessary for numerical
models of the flow of the LCIS, and in particular its northern
sector, to match empirical data23,24. A consequence of such a
temperature-induced acceleration would be a general reduction in
back stress within confined embayments, such as Cabinet Inlet,
potentially increasing shear stresses along the flow unit’s
lateral margins. This process is consistent with observations
that the northern sector of Larsen B Ice Shelf experienced an
increase in lateral rifting before its break up in 2002 (ref. 25). In
contrast, the influence of the layer’s presence on brittle
deformation may be in the opposite direction, with enhanced
ductile flow accommodating strain and the solid ice we report
being more resistant to tensile fracture than lower-density and
finer-grained firn. In such cases, the flow-parallel alignment of
elongate solid ice bodies such as that we report herein might serve
to resist the flow-orthogonal crevassing that commonly develops
in response to longitudinal tensile stresses, approaching the
marine limit of ice shelves. Finally, the spatial extent of this layer
at the ice-shelf scale, and the way in which its deformation
interacts with that of other material units, such as suture ice, and
basal channels and crevasses, will also influence the way in which
the layer’s presence affects overall shelf stability. In particular,
longitudinal troughs located on the surface of ice shelves have
been related to spatially coincident basal channels26, which have
themselves been associated with shelf instability27 through local
thinning and crevassing28,29. Although it is not yet known
whether the surface troughs on LCIS investigated herein are
associated with such basal channels, the recent finding that
material density is locally enhanced below similar surface troughs
on the Roi Baudouin Ice Shelf, Antarctica30, is consistent with
the enhanced melting and massive ice formation reported herein.
In the light of these complexities, identifying and exploring the
net influence of pond ice formation on ice-shelf stability can
only be achieved with confidence by including the full spatial
extent and physical properties of the refrozen ice layer into
a shelf-wide flow and fracture model, which may be regarded as a
future research priority.
The surface ponding responsible for the subsurface ice layer we
report herein is currently restricted to warmer regions of
Antarctica’s fringing ice shelves and in particular, but not
exclusively, to the northern and western sectors of the Antarctic
Peninsula. However, regional warming is predicted to spread
southwards and intensify substantially over forthcoming
decades5,31, so ponding is expected also to become more
widespread. Similarly, massive layers of warm, dense ice are
therefore highly likely to form within ice shelves present across a
substantial area of Antarctica, perhaps the entire continent if
warming continues into the 22nd century5, with important
consequences for ice-shelf flow, hydrofracture and stability.
Methods
Borehole drilling and logging. The borehole was drilled by pressurized hot water
and logged by OPTV13. The resulting OPTV log, analysed by WellCAD and BIFAT
software32, provides a geometrically accurate image, with a pixel size of B1 mm, of
the material composition of the complete borehole wall, as well as the structural
geometry of layers and inclusions intersecting it33. Since the OPTV log records an
image of reflected light, it also provides a proxy for the density of compacted snow,
firn and ice due to the progressive decrease in reflectivity, as voids close and bubbles
are occluded and collapsed13,15. Hubbard et al.15 exploited this relationship and
identified an exponential relationship between OPTV luminosity (L) and material
density (D) on the basis of samples recovered from a core retrieved from a borehole
logged by OPTV on the Roi Baudouin Ice Shelf, East Antarctica. However, due to
equipment loss, the light-emitting diode brightness of the OPTV probe used in the
current study was different from that of Hubbard et al.15, and a new calibration was
undertaken. This was based on the correlation of 40 core samples recovered from a
logged borehole, again located on the Roi Baudouin Ice Shelf (Fig. 5). The new
calibration yields a best-fit regression equation of D ¼ 950–40.1 e(0.0101L) (R2 ¼ 0.82),
with root mean square values of the residuals of 40.4, 35.2 and 21.7 kg mÀ 3 for the
density ranges 600–700, 700–800 and 800–900 kg mÀ 3, respectively. While these
values provide an error range for absolute densities derived from OPTV luminosity,
the relative changes in density reported herein, that is, along a single borehole log,
reduce to the precision of the method. This is approximated by the density range
(910 kgmÀ 3) divided by the luminosity range (256), or B3.5kgmÀ 3.
Borehole temperatures (Fig. 2c) were recorded by negative temperature
coefficient thermistors, recorded across a Wheatstone half-bridge by Campbell
Scientific micro-loggers. Resistances were converted to temperature using a
polynomial34 fitted to the manufacturer’s calibration curve refined via a second-stage
calibration in a distilled water/ice bath. Once recalibrated, sensors were replaced into
a new water/ice bath to determine temperature error, yielding a root mean square
error of ±0.03 °C. Borehole temperatures were logged for at least 100 days, and the
undisturbed ice temperature was calculated from an exponential function fitted to
the cooling curve35. By this time, all temperatures recorded below the near-surface
zone of seasonal thermal disturbance had stabilized to values that varied (over
timescales of days to weeks) by o0.05 °C. Thus, the englacial temperatures we report
herein were not influenced by the hot-water drilling process.
Firn modelling. The firn model used is IMAU-FDM v1.0, which takes into account
firn compaction, meltwater percolation and refreezing17,18. At the surface, the firn
model is forced with mass fluxes (snowfall, snowmelt, rain, sublimation, and
snowdrift sublimation and erosion) and surface temperature from the regional
500
600
700
800
900
20 40 60 80 100 120 140 160 180 200 220
OPTV luminosity (0 – 255)
D = 950 – 40.1 e
(0.0101L)
R2
= 0.82
Density(kgm–3
)
Figure 5 | Optical televiewer density–luminosity calibration. Material
density, measured gravimetrically on core samples, plotted against
equivalent OPTV luminosity, both measured along a borehole located on
Roi Baudouin Ice Shelf, East Antarctica.
NATURE COMMUNICATIONS | DOI: 10.1038/ncomms11897 ARTICLE

climate model RACMO2.3, run at a 5.5-km horizontal resolution in a domain over
the Antarctic Peninsula.
Satellite image analysis. MODIS Terra and Aqua level 1 (250 m) data were
ordered via the level 1 and atmosphere archive and distribution system, and used to
check for the presence of melt ponds in Cabinet Inlet. A total of 577 cloud-free
band 2 (848 nm) images between 1st December and 30th April from 2000 (Aqua)
or 2002 (Terra) to 2015 were classified to generate the time series shown in Fig. 3c.
Ground-penetrating radar. GPR data (Fig. 4) were collected using a Sensors and
Software Pulse Ekko Pro system operated in common offset mode with a 0.8-ns
sample interval and a 4,000-ns sample window. The console was mounted on the
skidoo and the 200 MHz antennae towed 15m behind on a plastic sledge at
B10kmhÀ 1 with an eight stack trace recorded every 3 m. A Leica VIVA GS10
GNSS rover unit was connected directly to the GPR console; the base station was
located at the drill site. Surface elevation was referenced to the Earth Gravitational
Model 1996 geoid, taken to represent sea level, and no correction was made for
dynamic topography. GPR data were processed in Reflex-W. Processing steps
included de-wow of 25-ns filter length, spherical divergence compensation, spectral
whitening between 70 and 200 MHz, and a two-dimensional mean-averaging filter.
The static correction was based on Leica Geo Office post-processed GNSS data,
resulting in a vertical accuracy of ±0.5 m. Radar wave propagation velocities used to
convert travel times to the depths shown in Fig. 4 were 0.2m nsÀ 1 for the upper-
most 3 m (based on a local 500 MHz common midpoint gather from the uppermost
reflector of Unit 1) and a typical value for glacial ice36 of 0.17 m nsÀ 1 below that.
Data availability. The data that support the findings of this study are available
from http://www.projectmidas.org/data/hubbard2016/ and from the corresponding
author upon request.
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Acknowledgements
This research was funded by the Natural Environment Research Council, grants NE/
L005409/1 and NE/L006707/1, and Aberystwyth University’s Capital Equipment fund.
The Leica VIVA GS10 GNSS system and Sensors and Software Pulse Ekko Pro radar
transmitter were loaned by the Natural Environment Research Council Geophysical
Equipment Facility, loan number 1028. We thank the British Antarctic Survey for
logistical support, and in particular the project’s field assistants Ashly Fusiarski and
Nicholas Gillett.
Author contributions
A.L. and B.H. led the overall project. Fieldwork was carried out by D.A., S.B., B.H., B.K.,
H.S., A.B. and A.L. D.J., M.O.L. and I.R. contributed to ice-shelf modelling, guiding
the field programme and interpretations. S.B. and A.L. led the remote-sensing
contributions. P.K.M. carried out the firn modelling. M.P., J.-L.T. and B.H. contributed
to the OPTV calibration. All authors contributed to writing the manuscript, which
was led by B.H.
Additional information
Competing financial interests: The authors declare no competing financial interests.
Reprints and permission information is available online at http://npg.nature.com/
reprintsandpermissions/
How to cite this article: Hubbard, B. et al. Massive subsurface ice formed by
refreezing of ice-shelf melt ponds. Nat. Commun. 7:11897 doi: 10.1038/ncomms11897
(2016).
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users will need to obtain permission from the license holder to reproduce the material.
To view a copy of this license, visit http://creativecommons.org/licenses/by/4.0/

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