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[1] Anisotropy of magnetic susceptibility (AMS) has been shown to provide specific useful information regarding the kinematics of deformation within subglacially deformed sediments. Here we present results from debris-rich basal glacier ice to examine deformation associated with glacier motion. Basal ice samples were collected from Tunabreen, a polythermal surge-type glacier in Svalbard. The magnetic fabrics recorded show strong correlation with structures within the ice, such as sheath folds and macroscopic stretching lineations. Thermomagnetic, low-temperature susceptibility, varying field susceptibility, and isothermal remanent magnetism acquisition experiments reveal that the debris-rich basal ice samples have a susceptibility and anisotropy dominated by paramagnetic phases within the detrital sediment. Sediment grains entrained within the basal ice are inferred to have rotated into a preferential alignment during deformation associated with flow of the glacier. An up-glacier directed plunge of magnetic lineations and subtle deviation from bulk glacier flow at the margins highlight the importance of noncoaxial strain during surge propagation. The results suggest that AMS can be used as an ice petrofabric indicator for investigations of glacier deformation and interactions with the bed.
Magnetic fabrics in the basal ice of a surge-type glacier
Edward J. Fleming,
Harold Lovell,
Carl T. E. Stevenson,
Michael S. Petronis,
Douglas I. Benn,
Michael J. Hambrey,
and Ian J. Fairchild
Received 25 March 2013; revised 6 September 2013; accepted 9 September 2013; published 31 October 2013.
[1]Anisotropy of magnetic susceptibility (AMS) has been shown to provide specic useful
information regarding the kinematics of deformation within subglacially deformed
sediments. Here we present results from debris-rich basal glacier ice to examine deformation
associated with glacier motion. Basal ice samples were collected from Tunabreen, a
polythermal surge-type glacier in Svalbard. The magnetic fabrics recorded show strong
correlation with structures within the ice, such as sheath folds and macroscopic stretching
lineations. Thermomagnetic, low-temperature susceptibility, varying eld susceptibility,
and isothermal remanent magnetism acquisition experiments reveal that the debris-rich
basal ice samples have a susceptibility and anisotropy dominated by paramagnetic phases
within the detrital sediment. Sediment grains entrained within the basal ice are inferred to
have rotated into a preferential alignment during deformation associated with ow of the
glacier. An up-glacier directed plunge of magnetic lineations and subtle deviation from bulk
glacier ow at the margins highlight the importance of noncoaxial strain during surge
propagation. The results suggest that AMS can be used as an ice petrofabric indicator for
investigations of glacier deformation and interactions with the bed.
Citation: Fleming, E. J., H. Lovell, C. T. E. Stevenson, M. S. Petronis, D. I. Benn, M. J. Hambrey, and I. J. Fairchild (2013),
Magnetic fabrics in the basal ice of a surge-type glacier, J. Geophys. Res. Earth Surf.,118, 2263–2278, doi:10.1002/jgrf.20144.
1. Introduction
[2] In this paper, we present a novel application of the
anisotropy of magnetic susceptibility (AMS) technique to
examine debris-rich basal ice. The ow of glacier ice can
produce similar structures to those produced in ductile defor-
mation within rocks [Maltman et al., 2000]. The analysis
of these structures and smaller-scale ice fabrics can provide
insight concerning the strain history and deformation of gla-
cier ice, since ice crystals are anisotropic [Castelnau et al.,
1998] and tend to develop a preferred orientation in response
to strain.
[3] The most commonly used method to examine fabrics
within glacier ice is the analysis of caxis crystallographic
orientations of ice crystals in thin section [e.g., Bader, 1951;
Rigsby, 1958] . This has been particularly useful in under-
standing how ice deforms under stress [Wilson, 2000; Wilson
and Sim, 2002]. More recently, the development of automated
techniqueshas introduced greater speed and objectivity [Wilen
et al., 2003]. However, Tison and Lorrain [1987] showed that
glacier ice can recrystallize over quite short timescales, so the
nal measured fabric may not represent the cumulative strain
but rather a more recent recrystallization event.
[4] Fabric analysis involving the measurement of the AMS
[Tarling and Hrouda, 1993] has provided considerable in-
sight into depositional [e.g., Ellwood and Ledbetter, 1977;
Hooyer et al., 2008; Lagroix and Banerjee, 2002] and defor-
mation histories [e.g., Borradaile and Jackson, 2004; Cifelli
et al., 2009; Parés et al., 1999] of rock and sediment. In
recent years, the technique has provided interesting new
information about various aspects of glaciology including
facilitating the interpretation of bed deformation [Hooyer
et al., 2008; Iverson et al., 2008], glacier ow direction
[Shumway and Iverson, 2009; Thomason and Iverson,
2009], and glaciotectonic history [Fleming et al., 2013] of
deformed glacial sediment. Despite the links between styles
of deformation seen in glaciers to those of rocks and sedi-
ment, there is (to our knowledge) no published research on
the AMS of glacier ice.
[5] Glacier ice formed by the rnication of snow, often
termed englacial ice [e.g., Hubbard et al., 2000], is dominated
by H
O and is therefore diamagnetic (negative susceptibility)
[Lancietal., 2001]. While the AMS of rocks dominated by
diamagnetic minerals has been used to investigate structural
deformation [e.g., Borradaile et al., 2012; de Wall et al.,
2000; Owens and Rutter, 1978], compared to that with ferro-
magnetic and paramagnetic dominated minerals, their relation-
ship to strain is not as well understood, and research into the
School of Geography, Earth and Environmental Sciences, University of
Birmingham, Birmingham, UK.
School of Geography, Queen Mary University of London, London, UK.
Department of Geology, University Centre in Svalbard (UNIS),
Longyearbyen, Norway.
Natural Resource Management, New Mexico Highlands University,
Las Vegas, New Mexico, USA.
Institute of Geography and Earth Sciences, Aberystwyth University,
Aberystwyth, UK.
Corresponding author: E. J. Fleming, School of Geography, Earth and
Environmental Sciences, University of Birmingham, Birmingham B15
2TT, UK. (
©2013. American Geophysical Union. All Rights Reserved.
JOURNAL OF GEOPHYSICAL RESEARCH: EARTH SURFACE, VOL. 118, 22632278, doi:10.1002/jgrf.20144, 2013
magnetic anisotropy of H
O ice has not been carried out.
Unlike englacial ice, there is a zone of ice at the base of
glaciers and ice sheets which exhibits a distinct set of physical
and chemical properties formed by processes operating at
the bed, commonly referred to as basal ice [Hubbard et al.,
2009; Hubbard and Sharp, 1989; Knight, 1997]. This ice
is thought to have predominantly formed through processes
including adfreezing, regelation, and hydraulic supercooling
[Cook et al., 2006; Hubbard, 1991; Hubbard and Sharp,
1993] at the base of glaciers and ice sheets and, as such,
has the ability to incorporate signicant amounts of detrital
minerals or subglacial sediment en masse [Hambrey et al.,
2005]. Depending on the composition of the source material,
this detrital material is expected to contain paramagnetic and
ferromagnetic grains that will overwhelm the diamagnetic
signal and create fabrics which retain more of a signal re-
lated to ice deformation. The basal ice of glaciers and ice
sheets therefore represents a suitable candidate for potential
AMS investigations.
[6] Glacier ice ows in response to gravitational forces act-
ing on a sloping ice body; however, this ow is resisted by
friction at the bed and lateral margins. Being located in the
zone between the bed and the bulk of the glacier ice, basal
ice is shown to be strongly affected by glacial motion and
is commonly highly deformed [Larsen et al., 2010; Samyn
et al., 2010; Souchez et al., 2000]. As such, a variety of
structures are produced reecting compression, extension, or
simple shear, depending on the ow regime of the glacier.
Basal ice commonly exhibits a strong ice-crystal caxis fabric
[Samyn et al., 2008]. As a result, one may expect fabrics asso-
ciated with such deformation, as well as being recorded in the
diamagnetic ice, to be preserved through a preferred orienta-
tion of grains within the detrital sediment. Therefore, in theory,
an AMS fabric should develop within the detrital component
of basal ice that reects the cumulative strain history.
[7] In this study, we apply the AMS technique to basal ice
exposed at the margin of a surge-type tidewater glacier in
Svalbard. The aims of this study are to (i) characterize the
AMS fabric by determining the orientation and degree
of alignment and shape of the susceptibility ellipsoid.
Also, since different minerals can produce vastly different
fabric characteristics (e.g., inverse fabric in single domain
magnetite) [Ferré, 2002], the magnetic mineralogy of the
ice is investigated through rock magnetic experiments. (ii)
Determine the relationship of the fabric to other visible strain
indicators within the ice at both outcrop scale and through
the analysis of aerial photographs. (iii) Examine the rela-
tionship of the fabric to the recent surge activity of the
glacier. Through these investigations, the potential of the
AMS technique for the analysis of basal ice is evaluated
and future areas in which the technique could be applied
are suggested.
2. AMS Theory
[8] AMS is one of a group of techniques that can be used to
measure the physical arrangement of particles and minerals
(petrofabric) in rock or sediment. It works on the principle
that when subjected to an external magnetic eld, an induced
magnetism is generated in rock or sediment that is dependent
on the magnetic susceptibility, (K) represented by the equa-
tion M=KH, where Mis the induced magnetization, His
the applied eld, and Kis the susceptibility measured in SI
units[Tarling and Hrouda, 1993].
[9] Susceptibility is essentially a measure of the Fe content
in a sample but is also controlled by the alignment, distribu-
tion, or crystalline orientation of these mineral grains and so
is anisotropic. In this way, the magnetic fabric normally rep-
resents the petrofabric of the rock or sediment, thus providing
information on its formation/deformation. AMS can be used
to accurately determine fabric in three dimensions and is best
visualized as an ellipsoid with a long (K
), intermediate, (K
and minimum (K
) axis. While the AMS records the
petrofabric of a rock, it is an oversimplication to assume that
reects the mean orientation of the long axisof grains. This
is because mineral composition and grain size can greatly af-
fect how it behaves in response to an external magnetic eld,
and as such, the magnetic mineralogy needs to be explored
before reliable fabric interpretations can be made.
[10] Most minerals forming a rock or sediment can be
dened by three magnetic behaviors: ferromagnetic, para-
magnetic, and diamagnetic. Ferromagnetic minerals (which
include ferrimagnetic sensu strictu minerals, e.g., magne-
tite) have a strongly proportional relationship between M
(induced magnetism) and H(strength of applied eld), with
a maximum value of M. These grains retain their magnetism
when subjected to a high magnetic eld, and therefore carry
a remanent magnetism (as used in paleomagnetic analysis).
Ferromagnetic minerals have very high susceptibilities (e.g.,
1500 × 10
for magnetite) and will dominate the fabric even
if present in very small concentrations. They can easily be
identied based on their thermomagnetic properties as they
have a structure that limits thermal disruption up to the
Curie temperature, after which grains behave paramagneti-
cally [Dunlop and Özdemir, 1997]. In contrast, paramagnetic
minerals have a proportional, nonpermanent relationship
between Mand H. Paramagnetism is exhibited in silicate min-
erals that contain Fe in the crystal lattice (e.g., biotite and chlo-
rite). An important property in the detection of paramagnetism
is that susceptibility decreases with increase temperature
according to the Curie-Weiss law. Finally, diamagnetic min-
erals (e.g., quartz and calcite) have a slight negative response
to increasing H. Diamagnetism is present in all rocks but has
very weak, negative susceptibilities (1×10
and Hrouda, 1993] and is normally overshadowed by even
small amounts of paramagnetic or ferromagnetic grains.
[11] Minerals can be classied as having shape, crystalline,
or distribution anisotropies. Shape anisotropy is common in
ferromagnetic grains (e.g., magnetite) and occurs when the in-
duced magnetization is preferentially orientated along the axis
of the grain. Crystalline anisotropy is common in paramag-
netic minerals (e.g., chlorite) and occurs where the induced
magnetization is dependent on the orientation of the crystal
lattices within the mineral (commonly with K
to the basal plane). In many examples, the fabrics produced
from shape and crystallographic anisotropy are directly com-
patible [e.g., Cifelli et al., 2009]. This is because paramagnetic
minerals (e.g., chlorite) tend to break preferentially along basal
planes and under extensional strain, these basal planes have
been shown to girdle about an axis parallel to extension, creat-
ing what is effectively an intersection lineation [Cifelli et al.,
2005]. Distribution anisotropy can play a role if ferromagnetic
grains are not randomly distributed through the matrix due to
magnetostatic interactions [Hargraves et al., 1991].
[12] In addition to this, grain size can play an important role
in the response of minerals in rock or sediment to an external
magnetic eld. Some minerals (e.g., magnetite), when present
in sizes below 0.03 μm, will exhibit single domain behavior
where susceptibility axes can switch creating inversefabrics
[Ferré, 2002]. As such, proper determination of the magnetic
mineralogy is vital (see section 4 for discussions of methods
used) to enable reliable conclusions to be drawn.
[13] AMS can characterize and quantify very weak or sub-
tle mineral fabrics and has been widely used in geology as a
means for investigating the processes involved in the forma-
tion of rocks and sediments [see references in Tarling and
Hrouda, 1993]. It is an important tool in understanding
how a material deforms in response to tectonic deformation
as stress acting on the sediment can cause grains to rotate
resulting in a preferential alignment. In glacial sedimentol-
ogy, the AMS of subglacial sediments has the ability to re-
veal subtle fabrics relating to ice deformation [Eyles et al.,
1987; Fleming et al., 2013; Gentoso et al., 2012; Shumway
and Iverson, 2009; Thomason and Iverson, 2009]. As well
as various eld-based applications, the technique has been ver-
ied through laboratory testing [Hooyer et al., 2008; Iverson
et al., 2008]. In these experiments, tills were sheared under
conditions thought to be operating at the bed. Microshears
were seen to develop that facilitate the rotation of grains into
the shear plane where they remain. This evidence was used
in support of the idea of March-type rotation [March, 1932],
where particles can role continuously in a viscous shearing
medium [Thomason and Iverson, 2006]. Basal ice generally
lies immediately above subglacial sediment and plays an im-
portant role in its formation through melt-out or lodgment
[Benn and Evans, 2010]. However, the way that sediment
particles within the ice respond to strain is not well understood.
The application of AMS to basal ice offers excellent opportu-
nity for some of these ideas to be investigated.
3. Glaciological and Geological Setting
[14] Tunabreen is a 33 km long tidewater glacier located
in central Svalbard (Figure 1). The glacier drains from the
Filchnerfonna and Lomonosovfonna ice caps and ows into
Tempelfjorden. The surrounding bedrock geology consists
of undeformed gently dipping Permian and Carboniferous
sediments composed of conglomerate, sandstone, and shale
of the Billefjorden Group. In turn, these are overlain by lo-
cally fossiliferous sandstones, carbonates, shales, and cherts
of the Dickson Land Subgroup [Cutbill and Challinor,
1965]. These strata were mostly deposited on a stable car-
bonate platform under shallow marine conditions [Harland
et al., 1997].
Figure 1. Geological map of Tunabreen and surrounding area, with glaciological structures and magnetic
fabric results. Geology redrawn after Dallmann et al. [2009, 2011]. Glaciological structures drawn from
aerial photographs dated July 2004. AMS results are plotted onto lower hemisphere, equal-area stereo-
graphic projections showing the mean susceptibility ellipsoids with the 95% condence ellipsoids and
the magnetic foliation (great circles) derived from the K
orientation. Inset map shows location of study
area within Svalbard.
[15] Radio echo-sounding records indicate that the glacier
is polythermal [Bamber, 1987]. Tunabreen is a surge-type
glacier and is the only one in Svalbard known to have surged
3 times, producing a consistent return period of approxi-
mately 40 years. Tunabreen last surged in 20032005, during
which the terminus advanced by up to 2km into
Tempelfjorden. Since surge termination, Tunabreen has
calved back to its present-day position, revealing spec-
tacular and easily accessible exposures of the basal zone
of the glacier at the lateral margins, including the glacier
bed interface. There are three dominant ice facies within
the exposures: a banded debris-rich facies composed of
alternating bands or laminae (110 mm thick) of ice-
containing diamicton and clean bubble-free ice, a solid
debris-rich facies composed of diamicton with some
stratication (hereafter referred to as banded facies
and solid facies,respectively) [after Hubbard et al.,
2009], and a clean, bubbly facies (hereafter termed
englacial facies)[afterHubbard et al., 2000].
[16] The ow regimes of Tunabreen are indicated through
structures exposed at the surface of the glacier (Figure 1). Ice
stratication and longitudinal foliation (utilizing glaciologi-
cal terminology of Hambrey and Lawson [2000]) are clearly
seen in aerial photographs. This stratication, which origi-
nates in an orientation dened by the margins of the ow
boundaries in the accumulation zone, becomes folded as
the ice ows. Fold tightness increases down-glacier, evolv-
ing to isoclinal toward the terminus. Fold limbs are rotated
a) SW
b) NW
c) SW
Figure 2. Field photographs of Tunabreen and the sections sampled. (a) NW section at the lateral margin
of Tunabreen. Blue ice in the right of the photograph represents englacial ice while the basal ice is shown by
the darker brown horizon in the center (snowmobile in foreground = 1 m). Height of section = 15 m. (b)
Photograph showing the locations of the NW and SE sections taken from the lateral moraine of Von
Postbreen. (c) SE section showing englacial ice (blue) overlying basal ice (brown and banded). Height of
section = 30 m.
parallel to the glacier margins and axial planes lie parallel to
glacier ow direction, creating ow-parallel structures
trending at 5°, which is commonly referred to as longitudinal
foliation [Hambrey and Lawson, 2000], a phenomenon well
known from Svalbard glaciers [e.g., Hambrey and Glasser,
2003; Hambrey et al., 2005].
[17] At the height of the most recent surge in 2004, almost the
entire length of Tunabreen exhibited intense surface crevassing.
Transverse crevasses dominated the pattern, forming perpendic-
ular to the longitudinal foliation and glacier ow direction.
Tunabreen has a tidewater margin and is dominated by a
strong extensional ow regime during surges, a characteris-
tic often seen in other Svalbard tidewater surge-type gla-
ciers [cf., Hodgkins and Dowdeswell, 1994; Murray et al.,
2003] without the compressional deformation commonly
exhibited at the terminus of land-terminating Svalbard glaciers
[Hambrey et al., 2005]. However, toward the terminus, the
eastern margin of Tunabreen reaches a conuence with the
neighboring Bogebreen, Phillippbreen, and Von Postbreen.
Here a component of oblique compressional deformation is
seen through the presence of structures that crop out at
the surface which truncate foliation and crevasse patterns,
interpreted as thrusts (Figure 1). This, combined with a
changing coastal morphology, results in the deviation of
ow at this location from a predominantly southward direc-
tion into a SSW direction.
4. Methods
[18] Two sections were analyzed at the lateral margins of
Tunabreen, hereafter referred to as the northwest (NW) and
southeast (SE) sections (Figure 2). Six sites were chosen
from the banded basal ice facies, covering both lateral and
vertical changes. In order to increase the chances of the
acquisition of reliable fabrics, sites were chosen where the
sediment concentration was greater than 10% by volume.
Cores were collected during April 2011, using a portable
rock drill with a 2.5 cm diameter, nonmagnetic, diamond-
tipped drill bit, and orientated using a Brunton compass by
scratching a ducial mark on to the side of the core. Cores
were subsequently transported to a cold room (at 20°C) at
the University Centre in Svalbard and cut using a nonmagnetic,
diamond-tipped circular rock saw into 21 mm sections, making
one to two samples from each core. Sedimentological and
structural data were collected in the eld using standard
procedures [cf. Evans and Benn, 2004]. Structural data from
the measurement of mineral lineations were collected in
March 2012.
[19] The AMS was measured using an AGICO KLY-3
Kappabridge operating at 875 Hz with a 300 A/m applied
eld at the University of Birmingham and an AGICO MFK-
1A Kappabridge operating at 976 Hz with a 200 A/m applied
eld at New Mexico Highlands University. In total, 71 sam-
ples were analyzed with an average of 10 subsamples per site.
The following parameters were used to evaluate the suscepti-
bility ellipsoid [cf. Tarling and Hrouda, 1993]. The mean
susceptibility (K
Kmean ¼K1þK2þK3
where (K
) are the principal susceptibilities
(SI units). The shape of the ellipsoid can be characterized
using lineation (L) and foliation (F)parameters[Khan, 1962]
and are calculated as
[20] Also used are the corrected anisotropy degree (P
), to
determine the strength of the fabric, and the shape parameter
(T), to dene the shape of the susceptibility ellipsoid [Jelínek,
1981], which respectively are
Pj¼exp ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
þln K1
ln K1
[21] Because of the low susceptibility of the samples,
careful cleaning and calibration of the sample holder were
undertaken between each site, as even small amounts of fer-
romagnetic or paramagnetic dust may swamp the suscepti-
bility signal of the samples [Borradaile et al., 2012]. In
spite of the Kappabridge being sensitive to 0.5× 10
SI with
an accuracy of 0.1%, the anisotropy values near zero can be
anomalously high [Biedermann et al., 2013; Hrouda and
Kapička, 1986; Rochette, 1987]. Although this is not thought
to affect fabric orientations [Callot et al., 2010; Hrouda,
2004], its effect can cause problems when calculating the
Table 1. Mean Site AMS Data (See Section 3 for Calculation)
Site NK
95% Error K
95% Error K
95% Error LFP
TB2 6 9.04E-06 4/12.522/6160/76 49/19 273/649/4 1.018 1.052 1.077 -0.046
TB3 13 2.28E-05 338/26 36/21 243/11 36/21 132/62 35/16 1.017 1.009 1.027 -0.303
TB4 7 1.28E-05 345/15 34/13 254/335/12 152/75 18/12 1.022 1.035 1.058 0.227
TB5 12 2.37E-05 355/25 19/14 261/971/13 152/63 71/14 1.076 1.024 1.107 -0.512
TB7 11 1.87E-05 2/11 46/993/746/26 214/77 33/13 1.012 1.025 1.038 0.335
TB8 12 1.89E-05 21/26 35/12 280/21 38/22 157/56 31/7 1.037 1.028 1.066 -0.142
N= Number of Samples; K
= Mean Susceptibility; K
= Orientations (Declination and Inclination) of the Principal Susceptibility Axes with 95%
Condence Ellipses; L= Lineation (L=K
); F= Foliation ( F=k
); P
= Anisotropy Degree; T= Shape Parameter.
200 400
400 600
c) i) ii)
Temperature oC Temperature oC
Absolute susceptibility (10-6)
Absolute susceptibility (10-6)
Absolute susceptibility (10 t)
Applied field (A/M) Applied field (A/M)
tb3 tb5
Normalised intensity (J/Jmax)
Normalised intensity (J/Jmax)
0 100 100
0 0.5
0.5 1 1 1.5
1.5 2 2 2.5
Applied field (T)
Applied field (T)
200 200
300 300
400 400
500 500
600 600
700 700
800 800
a) i) ii)
b) i) ii)
d) i) ii)
Normalised suceptibility (K/K0)
Normalised suceptibility (K/K0)
Normalised suceptibility (K/K0)
Normalised suceptibility (K/K0)
Temperature (Kelvin)
Temperature (Kelvin)
0.4 0.4
0.2 0.2
0.6 0.6
0.8 0.8
1.0 1.0
50 50
100 100
150 150
200 200
250 250
300 300 20 40 60 80 100
Figure 3. Rock magnetic experiments. (a) Low-eld susceptibility (K) versus temperature curves for (i) NW
and (ii) SE sections. In each case, the heating curve is black and the cooling curve is gray. (b) Low-eld sus-
ceptibility versus applied eld (A/M) curves normalized to the lowest susceptibility value (K/K
and (ii) SE sections. (c) The normalized reciprocal (K/K
) susceptibility versus temperature for (i) NW and
(ii) SE sections and (iii) the ratio of the susceptibility at the lowest temperature to the susceptibility at room
temperature (K
), plotted against magnetic susceptibility. (d) IRM experiments showing variation in
normalized magnetic intensity (J/J
) with applied eld (T) for (i) NW and (ii) SE sections.
anisotropy parameters, and as such, any subsamples with sus-
ceptibilities in the range of 5μSI to 5 μSI were discounted,
as recommended by Hrouda [2004].
[22] In all AMS investigations, determination of the mag-
netic mineralogy is of importance because of the different
fabric characteristics which can be produced by different
minerals. The detrital component, typically 1040% volume
of sample volume, was extracted from the diamagnetic H
by sublimation, and investigations of the magnetic mineralogy
were conducted at New Mexico Highlands University. The
variation of low-eld magnetic susceptibility with temperature
and eld strength was conducted on an AGICO MFK-1A
Kappabridge with a CS4 high-temperature susceptibility
attachment. Thermomagnetic experiments were conducted for
six samples from two sites measuring variations of magnetic
susceptibility on heating at a 6°C interval from room tempera-
ture (20°C) to 700°C. Low-temperature susceptibility experi-
ments were conducted on all sites using an in-house cryostat
system coupled with the Kappabridge. The samples were
cooled to 77K in liquid nitrogen, and the bulk susceptibility
measured every 18 s during warming to room temperature.
The low-eld variation of AC susceptibly was measured in
the following elds: 5, 10, 20, 30, 40, 50, 60, 70, 80, 100,
150, 200, 250, 300, 350, 400, 500, 600, and 700 A/M follow-
ing the procedure of Hrouda et al. [2006]. In addition, the fer-
romagnetic fraction of six samples was analyzed through the
acquisition of Isothermal Remanent Magnetization (IRM) ex-
periments (i.e., partial hysteresis loops), rst by demagnetizing
the sample in an alternating eld (AF) to remove the natural
remanent magnetism (NRM), followed by applying an exter-
nal eld at progressive stronger elds up to a peak of 2.5 tesla
(T) eld. This experiment was measured on an AGICO JR6-A
dual-speed spinner magnetometer in a magnetically shielded
room that attenuates Earthseldtolessthan0.1%.
5. Results
5.1. Magnetic Mineralogy
[23] The mean susceptibility (K
) of the samples ranges
from 9 to 23 μSI (average 20 μSI, Table 1), well within the
paramagnetic realm [Tarling and Hrouda, 1993]. The mag-
netic susceptibility of the samples in which the detrital sedi-
ment was separated from the ice are 96μSI for TB3 and
38 μSI for TB5, reecting the absence of diamagnetic H
[24] Low-temperature susceptibility measurements (Figure 3c)
can be used to distinguish between the contribution of para-
magnetic from ferromagnetic phases, since antiferromag-
netic, diamagnetic, and most ferromagnetic minerals have a
temperature-independent susceptibility in the 77 to 295 K
temperature range [Richter and van der Pluijm, 1994]. The
curves show good Curie-Weiss temperature dependence,
where susceptibility decreases with increasing temperature
[Nagata, 1961]. The ratio of low-temperature/room tempera-
ture versus the mean value of room temperature susceptibil-
ity are plotted (Figure 3c) and the ratio of all samples is
above 3.2, indicating a substantial paramagnetic component
to the low-eld AMS.
[25] The variation of low-eld susceptibility with temper-
atures from 20 to 700°C (Figure 3b) shows a decreasing sus-
ceptibility with increase of temperature on some curves
within the range of 20 to 250°C following Curie-Weiss behavior
(Figure 3bi), whereas others show an independent or slight
increase in susceptibility within this range (Figure 3bii).
Above 250°C, all samples show an increase in susceptibility
1.04 1.08 1.12
K -6
x 10
a) b)
Figure 4. Variation of AMS parameters for all sites. (a) Anisotropy parameter (P
) versus susceptibility
(K). (b) L% versus F% (magnetic lineation versus magnetic foliation). (c) Shape parameter (T) versus
anisotropy parameter (P
with increasing temperature and exhibit strong peaks at 560°C,
presumably indicating either the growth of new ferromagnetic
phases on heating, or the Hopkinson peakowing to a minor
amount of Fe-Ti oxide present within the sample.
[26] The variation of eld strength with susceptibility can
also be used to provide constraints on the magnetic mineral-
ogy (Figure 3c). This experiment works on the principle that
diamagnetic and paramagnetic minerals exhibit a linear rela-
tionship between magnetization and the magnetization
eld, whereas the susceptibility of some ferromagnetic
minerals exhibit a strong eld-dependent susceptibility
[Hrouda et al., 2006]. Nonsystematic behavior is seen in
all samples in the 0 to 200 A/M range reecting the high
error margin in the measurement of susceptibility at these
frequencies in low-susceptibility samples. However, above
200 A/M, all samples show a eld-dependent susceptibility,
which increases up to 500 A/M before decreasing. This pre-
sumably represents a minor contribution to the susceptibility
by a ferromagnetic component.
[27] This ferromagnetic component is investigated further
through the acquisition of IRM (Figure 3d). This works on
the principle that the coercivity of a mineral varies with
composition and grain size [Dunlop and Özdemir, 1997].
N= 69
K1 = 359/20
K2 = 089/02
K3 = 184/70
95% confidence
Figure 5. Stereographic projection of AMS results from all
samples showing K
(black squares), K
(gray triangles), and
(white circles) with 95% condence ellipses. Refer to
section 3 for derivation of anisotropy parameters.
Figure 6. Two-dimensional section logs of (a) NW section and (b) SE section (no vertical exaggeration),
with magnetic fabrics for all sites showing the three mean principal susceptibility axes plotted on to lower
hemisphere stereographic projections. See Figures 1 and 2 for locations.
For example, the saturation magnetization of hematite is near
3 T while magnetite is fully saturated by 300 mT. The IRM
acquisition curves all fail to show complete saturation at
2.5 T indicating the presence of a high-coercivity phase,
presumably hematite.
5.2. Anisotropy of Magnetic Susceptibility
[28] Samples yield susceptibility ellipsoids that are pre-
dominantly triaxial (Figure 4), where Fis roughly equal to
L, although variation exists between subsamples, ranging
from strongly oblate to strongly prolate (possibly in part
a) b)
c) d)
e) f)
g) h)
5 cm 2 m
4 m 1 m
50 cm 20 cm
30 cm 20 cm
Figure 7. Field photographs of typical structures present within the basal ice. (a) Banded debris-rich ice fa-
cies from which most of the samples were collected showing subhorizontal alternating bands (>110 mm) of
debris-rich and debris-poor ice. (b) Thick (8 m) section of basal ice at NW section with lenses of clean bubbly
(englacial) ice. The basal ice is partially obscured by icicles. (c) Banded and solid debris-rich basal ice facies
thrust over blocks of clean englacial ice. (d) Foldedbanded ice showing a double vergence pattern in the folds.
(e) Isoclinal recumbent folds verging to the right. (f) Close up of isoclinal Zfold showing vergence to the
right. The fold axis can be traced through the ice giving a three-dimensional view of the fold. (g) Mineral
stretching lineations in debris-poor basal ice above debris-rich horizon. (h) Mineral stretching lineations
and elongated bubbles on the surface of debris-rich horizons within clean ice.
arising from the high error margins when calculating parame-
ters at low susceptibilities) [Biedermann et al., 2013; Hrouda,
2004]. The mean corrected anisotropy degree (P
) is relatively
high (1.05) (Figure 4a) compared with the typical values within
sediments dominated by paramagnetic minerals. However,
hematite can have very high (>100) anisotropies [Guerrero-
Suarez and Martín-Hernández, 2012; Tarling and Hrouda,
1993] and this high value may reect the presence of a minor
amount of hematite contributing to the anisotropy.
[29] AMS results are shown on lower hemisphere, equal-
area stereographic projections (Figure 5) and the corresponding
AMS results from individual sites and their sampled locations
are shown in Figure 6. The mean maximum susceptibility ori-
entation (K
) plunges gently (20°) to the north, with a general
north-south trend (mean = 359°) (Figure 5), subparallel to dom-
inant glacier ow direction, calculated from the trend of the gla-
cier and orientation of the macroscopic longitudinal foliation.
The minimum susceptibility axes (K
) are subvertical, dening
the pole to the magnetic foliation (K
plane). At the NW
cluster at 20° to 001° and K
axes cluster at 030° to 188°.
The SE section, despite being close to the opposite margin of
the glacier, gives a broadly similar fabric orientation to the
SE section with K
axes clustering at 20° to 355° and K
clustering at 27° to 097°.
5.3. Analysis of Visible Structures
[30] In subglacial sediments investigated at other sites, the
orientation of magnetic fabrics has been shown to reect
glacier-induced simple shear relating to the ow direction
of glacier ice [e.g., Fleming et al., 2013; Hooyer et al.,
2008; Iverson et al., 2008; Shumway and Iverson, 2009;
Thomason and Iverson, 2009]. Basal ice lies at this crucial
boundary between the bulk glacier ice and deforming bed
and, as such, has been interpreted to deform similarly in a
way strongly related to the ow of the glacier [Knight,
1997]. Evidence for deformation is seen at both sections as
a variety of structures including folds, faults, and lineations
(Figure 7). One of the unique features of the study of defor-
mation within glacier ice is that, as opposed to most other
geological materials, ice is often translucent or transparent.
This allows structures to be seen in three dimensions through
the ice face (e.g., Figures 7f7h), aiding analysis and inter-
pretation. These structures can be analyzed to provide insight
into the kinematics of deformation, thus providing indepen-
dent verication of the state of strain within the basal ice.
As such, comparisons can be made with the magnetic fabric
to determine its relationship to strain within the ice.
[31] Folding and boudinage are common within the basal
ice at both sections, especially at the SE section. Here the
banded ice facies (Figure 7a), which presumably formed at
an orientation parallel to the glacier bed or overriding obsta-
cles, is highly folded in places (Figures 7d7f). Folds are
typically steeply inclined to recumbent and strongly asym-
metric, with interlimb angles ranging from tight to isoclinal.
One interesting and, at rst somewhat confusing, aspect of
these folds is that vergence direction can appear on the
two-dimensional ice face (Figure 7d) to be in both direc-
tions. Folds also occasionally form concentric augen-like
rings (Figure 9a). The axes of these folds lie in a north-south
orientation, generally parallel to the glacier ow direction
and parallel to the maximum susceptibility orientations
). This indicates that rather than being purely cylindrical,
which is often assumed, folds are highly noncylindrical
in a style often referred to as sheath folding [Alsop and
Carreras, 2007; Alsop and Holdsworth, 2004; Alsop
et al., 2007].
[32] The fabric of the debris and bubbles within the banded
ice facies is not planar. In contrast, a strong linear component
is present (Figures 7g and 7h). Debris is observed to be ar-
ranged in linear aggregates and has, in places, been strongly
smeared along an axis. Lineations, measured at the SW
section, cluster at 10° to 005° (Figure 8). In places, strongly
elongated bubbles are orientated in the same direction as
the debris lineations (Figure 7h). Debris lineations are
also seen to form generally parallel to fold axes and almost
completely parallel to the magnetic lineation which, in most
previous studies of AMS of deformed sediments, represents
the direction of stretching [e.g., Cifelli et al., 2005; Liss
et al., 2002; Parés and van der Pluijm, 2002].
[33] Faulting is common, illustrating that as well as ductile
folding, brittle deformation has also occurred within the basal
ice at both sections (Figures 7b and 7c). The NW section con-
tains a number of faults that are typically orientated N-S to
NE-SW, shallow to moderately dipping to the east, parallel
or subparallel to the glacier margins. At the SE section, faults
strike in an N-S orientation; however, both dip angle and dip
direction are variable. The majority of the faults have a re-
verse offset, but many contain subhorizontal debris lineations
on their surface, indicating oblique or even transverse slip in
some cases and suggesting a transpressional glaciotectonic
regime. Thrusting has clearly resulted in the tectonic thicken-
ing of basal ice, for example, Figure 7c where banded debris-
rich ice has been thrust up over blocks of clean englacial ice.
6. Discussion
6.1. Control on AMS Fabric
[34] The low susceptibility of the samples indicates a vol-
umetrically signicant proportion of diamagnetic minerals,
presumably quartz, calcite, and ice. Yet, the presence of
paramagnetic and ferromagnetic phases provides a positive
Mean K1 orientation with
95% confidence ellipse
Figure 8. Mineral stretching lineations and mean K
tation with condence ellipse for SE exposure with 95% con-
dence ellipses.
susceptibility which probably controls the magnetic fabric
[Tarling and Hrouda, 1993]. The dependence of susceptibility
on temperature follows Curie-Weiss behavior at low temper-
atures, suggesting a dominance of paramagnetic minerals
[Richter and van der Pluijm, 1994]. At high temperatures,
the increase in susceptibility can be attributed to the growth
of new ferromagnetic minerals with the peak at 550°C, possi-
bly representing a suppressed Hopkinson peakof a minor
ferromagnetic contribution. The dependence of the suscepti-
bility on eld strength could be attributed to a ferromagnetic
contribution, since pure paramagnetic minerals yield eld-
independent behavior [Hrouda et al., 2006], but given the
low susceptibility and the strong dependence of susceptibility
with temperature, its inuence on the AMS is considered
minor. The high coercivity picked out bythe IRM experiments
indicates that hematite most likely controls this ferromagnetic
contribution. Therefore, we interpret the origin of the AMS
signal as having a mixed magnetic mineralogy. This is domi-
nated by paramagnetic phases which, given the composition
of the material, are likely to be phyllosilicate clays with possi-
bly a minor contribution of a high-coercivity ferromagnetic
phase, presumably hematite.
[35] The presence of ow-parallel magnetic lineations as-
sociated with sediment dominated by phyllosilicate clay min-
erals and hematite may at rst seem counterintuitive as both
minerals typically display crystalline anisotropy, where the
maximum susceptibility axis lies in the basal plane of the
mineral [Tarling and Hrouda, 1993]. As such, K
tions are not parallel to the long axis of grains but rather
depend on the crystallographic structure with the minimum
susceptibility perpendicular to the basal plane. In spite of
this, magnetic lineations are common in rocks dominated
Figure 9. Schematic diagram illustrating the relationship of structures to AMS fabrics. (a) Three-dimensional
cartoon of Tunabreen (vertical scale exaggerated) showing the structure of the foliation/stratication,
faults, and basal ice in relation to the orientation of the AMS lineation. (b) Sketch of banded basal ice show-
ing the preferred alignment of grains. (c) Visualization of subsequent AMS fabric through the AMS ellip-
soid with K
(maximum), K
(intermediate), and K
(minimum) susceptibility axes. (d) Presentation of
ellipse through stereonet displaying the mean northerly orientated K
parallel to glacier ow direction.
by phyllosilicate minerals and are shown to form parallel to
the direction of stretching [Cifelli et al., 2005, 2009; Parés
and van der Pluijm, 2002]. Phyllosilicate minerals tend to
break along their basal plane, which when under extensional
stresses, become disposed about an axis parallel to stretching,
thus creating a magnetic lineation that is directly compatible
with fabrics created through shape anisotropy.
6.2. Relationship of Structures to AMS
[36] The magnetic fabrics show strong apparent correspon-
dence with the orientations of macroscopic structures present
within the ice. The stratication (mapped in Figure 1 and sche-
matically drawn in Figure 9), which would have originally
formed in an orientation parallel to ow boundaries in the accu-
mulation zone, has been tightly folded forming a longitudinal
foliation (Figure 9a) under a strong extensional regime. This fo-
liation generally lies parallel to the AMS lineation. The close
relationship of the strike of the longitudinal foliation and the
AMS lineation within the basal ice suggests that, as one would
expect, the basal ice has been deformed by glacier motion.
[37] At the outcrop scale, the banded ice facies within the
basal ice have been folded under noncoaxial stretching and
simple shear (Figure 10d). In these conditions, folding initiates
during the initial stage of shear where the eld of compression
occurs at a high angle to bedding (Figure 9dii). As deforma-
tion continues, the strain ellipse rotates to a low angle to bed-
ding and extensional processes become dominant, resulting in
boudinage (Figure 10diii). The folds created within the basal
ice at Tunabreen have fold axes which are strongly curvilin-
ear (Figure 9a). This represents a noncylindrical style of fold-
ing, commonly referred to as sheath folding [Alsop et al.,
2007]. Sheath folds normally form when perturbations
during the initial stages of folding are greatly exaggerated
in high-strain conditions [Cobbold and Quinquis, 1980]. As
folding progresses, fold noses become stretched and elon-
gated, and fold axes rotate toward the direction of shear
within the ice and the fold axes becomes parallel with the
main stretching direction (Figure 10c). Sections perpendic-
ular to the shearing direction are characterized by concen-
tric, eye-shaped folds and doubly verging fold directions
(Figure 10b). Sheath fold noses lie parallel to the orientation
of AMS lineations as fold axes are essentially indistinct
from stretching lineations.
[38] Deformation of the ice at Tunabreen has also resulted
in the formation of distinct linear features within the basal
ice (Figures 7g and 7h). Clusters of debris are smeared out
and aligned about an axis. The smearing of grains in basal
ice has been referred to in the past [Hubbard and Sharp,
1995; Hubbard et al., 2000], but its relationship to cumu-
lative strain has not. Similar lineations are often seen in
structurally deformed metamorphic rocks [Neves et al.,
2005; Twiss and Moores, 1992], commonly referred to as
stretching lineations. Stretching lineations in deformed rocks
form in an orientation parallel to the direction of stretching
during ductile deformation [Ramsay and Huber, 1983].
Thus, structural analysis of their orientation can provide
useful information about the kinematics of deformation and
deformational history.
[39] At Tunabreen, these lineations lie at an orientation
parallel to the fold axes of sheath folds and the strike of macro-
scopic surface lineations, and subparallel to the ow direction
of the glacier. Also, these lineations lie almost completely
parallel to the magnetic lineations (Figure 8), thus providing
independent verication that these form in an orientation paral-
lel to stretching, and as such, we interpret them as stretching
lineations. Under high-strain conditions, detrital grains within
a) b)
i) ii)
30 cm
magnetic lineation
Figure 10. (a) Photograph of sheath fold within banded ice facies from SE section showing augen pattern
with concentric rings and double verging folds. (b) Three-dimensional model of sheath folding observed
at SE section. (c) Interpretation of sheath fold development and associated strain ellipse, showing (i)
predeformation and original conguration, (ii) initial folding, and (iii) evolved sheath folding and rotation
of fold axes parallel to ow.
the ice will rotate into the most stable orientation about an axis
parallel to stretching, forming the lineations. As these lineations
are parallel to the interpreted direction of stretching, they can be
used in a similar way to which they are in structural geology
and the analysis of tectonically deformed rocks in order to give
the kinematics of deformation within the ice.
6.3. Kinematics of Deformation Within the Basal Ice
[40] The up-glacier dip of K
is a feature commonly seen
within subglacial sediments under simple shear [Shumway
and Iverson, 2009; Thomason and Iverson, 2009]. The mean
plunge of the K
lineation at 20° up glacier may (shown in
Figure 5 and drawn schematically in Figures 9b9d) indicate
that within the basal ice, as well as pure shear, there is a com-
ponent of noncoaxial strain and simple shear causing the
updip rotation of K
orientations, matching the strain ellipse.
Ring shear experiments of subglacial tills subject to simple
shear reveal that steady state AMS fabrics develop at strains
of 730, in which K
lies parallel to shear direction dipping
28° away from shear direction [Hooyer et al., 2008; Iverson
et al., 2008]. These experiments produced almost identical
fabric characteristics and clustering patterns as those
displayed in Figure 5. One could argue a similar model for
the rotation of grains within basal ice, where slip between
the grains and the ice keeps particles from rotating through
the shear plane (as suggested by March [1932]) therefore
rejecting Jeffery rotation [Jeffery, 1922] within ice. However,
as the magnetic mineralogy of the tills used are different,
caution is applied when making direct comparisons, and
conclusions should not be made until further laboratory
testing on materials with a similar mineralogy is obtained.
6.4. Relationship to Surge Dynamics
[41] At the NW section (Figure 11a), magnetic lineations
lie in an orientation that deviates slightly away from the dom-
inant glacier ow direction in this area. If the fabrics formed
purely by stretching and shear due to friction at the bed, one
may expect the magnetic lineations to trend parallel to ice
ow. However, the ow of the glacier ice is not uniform
across the ice surface. At the margins, lateral drag can result
in the development of marginal shear zones such as those
recorded after the 19821983 surge of Variegated Glacier
[Lawson et al., 1994; Sharp et al., 1988]. At Tunabreen,
the deviation of the magnetic lineations from parallel to
glacier ow is probably caused by the rotation of the strain
ellipse away from glacier ow direction at the margins un-
der noncoaxial strain (Figure 11bi).
[42] At the SE section (Figure 11a), the orientation of the
magnetic lineations cannot be explained in the same way,
since lateral shear would cause the inclination of lineations
in the opposite direction to that observed. However, magnetic
lineations lie parallel to longitudinal foliation identied on
aerial photographs at that location, indicating that ice ow
has been rotated in the opposite direction to that expected.
This can be explained by the presence of an irregular presurge
carving margin (Figure 10bi). During the surge, local splaying
into an embayment facilitated rotation of the ice in an anti-
clockwise direction (Figure 10bii), causing the slight deviation
b) i)
a) 2002 pre-surge
2004 surge-maximum
2012 - present configuration
Irregular margin of
Von Postbreen and
formation of
lateral shear at
margins causing
deviation of strain
ellipse from flow
Lateral spread
and clockwise
rotation of ice
into bay
Calving of margin to
present day
Lateral shear plane
Inferred ice flow
Longitudinal foliation
Figure 11. (a) Aerial photograph mosaic of the Tunabreen terminus at surge maximum in 2004 with the
location of the sections studied. (b) Interpretation of the formation of the magnetic lineations showing (i)
2002 presurge conguration and irregular margin of Von Postbreen. (ii) 2004 surge maximum showing
the orientations of shear in the NW section and the lateral spreading and clockwise rotation of surface
foliation and magnetic lineation at the SE section. (iii) Present conguration of Tunabreen at time of
study (2012).
from overall ow direction of both the AMS fabrics and sur-
face foliation (Figure 10biii).
6.5. The Use of AMS for the Analysis of Deformation
Within Basal Ice
[43] This study has shown that the detrital component of
basal ice contains sediment from which an AMS fabric
can be measured and provide insight into subglacial pro-
cesses. The magnetic fabric appears to be a direct reection
of the petrofabric of the detrital grains within the ice. A
magnetic lineation is recorded, parallel to the inferred direc-
tion of stretching and simple shear within the ice. This result
provides support for the validity of the AMS of subglacial
sediment, where magnetic lineations are also seen to form
parallel to stretching/shear direction within the sediment
[Fleming et al., 2013; Gentoso et al., 2012; Shumway and
Iverson, 2009; Thomason and Iverson, 2009]. The potential
preservation of AMS fabrics from basal ice to sediment dur-
ing melt-out requires further study. However, as an AMS
fabric is seen within basal ice, caution should be taken when
interpreting AMS fabrics in subglacial sediments as being
formed solely by bed deformation, especially when an ori-
gin through melt-out is suspected.
[44] Utilizing the methodology described here, the AMS
technique can be directly reproduced and applied to other gla-
ciers. AMS has several advantages over other petrofabric tech-
niques [Iverson et al., 2008]. The fabric can be determined
relatively quickly, accurately, and objectively, and the suscep-
tibility ellipsoid can be calculated in three dimensions. AMS
represents the volume average of many grains in each subsam-
ple and many subsamples make up a site. Being sensitive
to minor changes in the state of strain, investigations of the
AMS of basal ice has the potential to provide knowledge on
the processes occurring at the ice-bed interface, bridging the
gap between the analysis of visible structures at the surface
of the glacier and deformation within subglacial sediments.
AMS, therefore, has the potential to contribute to the highly
debated topic of glacier bed deformation.
[45] AMS has been used to calculate shear strains in de-
formed rocks and sediment [Borradaile, 1988, 1991]. The
link between the AMS fabric strength (based on the degree
of clustering of susceptibility axes) and strain has also been
investigated within subglacial sediments through experimen-
tal work with ring shear devices [Iverson et al., 2008]. This
study showed that fabric strength increases with increasing
shear strain, up to a point under which steady state fabrics
were reached. In the future, it may be possible to apply sim-
ilar experimental tests to the AMS of basal ice and thus inves-
tigate the link between fabric strength and strain. Also, in
contrast to ice-crystal fabric studies, which measure the caxis
orientation of ice crystals [e.g., Bader, 1951; Tison et al.,
1994; Wilson and Peternell, 2011], the AMS fabric is domi-
nated by the paramagnetic and ferromagnetic proportion of
detrital material in the basal ice. Therefore, the study of
AMS in conjunction with ice-crystal fabric analysis allows
the detrital portion of the ice to also be analyzed which, in
contrast to glacier ice, is not subject to recrystallization under
the pressure/temperature ranges encountered in glaciers.
[46] At this study site, although AMS has highlighted in-
teresting variation in the state of strain, the glacier ow direc-
tion was never in a doubt. The site was chosen intentionally
as visible structures such as the surface longitudinal foliation
measured from aerial photographs and the orientations of
folds and lineations at the outcrop scale provide a reference
frame for comparison with the AMS results. This has enabled
further interpretations to be made and shows that folding
style is dominated by sheath folds and lineations which form
parallel to stretching within the ice. However, one interesting
situation in which AMS could be applied is where the ow
direction or past strain history is not known or is poorly un-
derstood. For example, on large ice sheets where glacier ow
is slow and surface structures are absent or where ow direc-
tion is ambiguous [e.g., Conway et al., 2002], AMS of basal
ice collected from ice cores could potentially be analyzed to
provide insight into shear direction at the base of the ice
sheet. The AMS technique could also aid research into the
subject of massive ground ice, which is thought to originate
as the basal portion of pre-existing glaciers often dating
back to the Pleistocene, and is often buried and preserved
in permafrost regions [Fritz et al., 2011; Waller et al.,
2009]. Here little is known about paleoice ow directions,
and therefore, AMS could potentially provide considerable
paleoglaciological insight.
7. Conclusions
[47] The AMS fabrics of basal ice and their relationship
to deformation during the most recent surge of Tunabreen
have been investigated and number of conclusions can subse-
quently be drawn:
[48] 1. The AMS of basal ice can be measured, and the
three components of the susceptibility ellipsoid can rapidly
calculated, in the same way that is commonly done for sedi-
ment and rock.
[49] 2. Magnetic fabrics at the sections examined are con-
trolled predominantly by the preferred alignment of inclu-
sions of detrital sediment within the ice. The susceptibility
and anisotropy in this sediment is dominated by paramag-
netic minerals (presumably phyllosilicate clays). In some
samples, a high-coercivity phase, presumably hematite is
also present, possibly contributing to the fabric.
[50] 3. The folding style within the deformed basal ice is
highly noncylindrical. This is not unusual given the high shear
strains expected within the deforming ice and the perturbations
in ow that exist across the glacier prole. Within subglacial
glaciotectonites, since the deformation of underlying subgla-
cial sediments is largely controlled by the overlying ice mo-
tion, noncylindrical folding should be expected.
[51] 4. AMS lineations are parallel to, and independently
veried by, the macroscopic lineation given by the presence
of stretching lineations and the axes of sheath folds. The ori-
entation of stretching lineations in basal ice has the potential
to be used as a proxy for stretching direction within the strain
ellipse, in the same way that is used in structural geology.
[52] 5. Magnetic lineations at the NW section have been
affected by lateral shear, causing a minor amount of devia-
tion of the lineations away from being parallel to the mean
trend of the macroscopic foliation, reecting this noncoaxial
deformation. At the SE section, the irregular presurge cong-
uration of the contact between Tunabreen and Von Postbreen
has affected strain patterns and led to the anticlockwise rota-
tion of magnetic lineations, stretching lineations, and the
macroscopic foliation, resulting in a magnetic lineation ori-
entated subparallel to the dominant glacier ow direction.
[53]Acknowledgments. This work forms part of the NERC-funded
GAINS (Glacial Activity in Neoproterozoic Svalbard) grant (NE/H004963/1)
with a tied PhD studentship held by E.J.F. H.L. was also funded by a
NERC PhD studentship (NE/I528050/1). Additional funding for eldwork
was provided by the Queen Mary Postgraduate Research Fund and the
Research Council of Norway Arctic Field Grant program. ASTER satellite
images were acquired from the NASA Land Processes Distributed Active
Archive Center (LPDAAC) and aerial photographs were provided by the
NERC Earth Observation Data Centre (NEODC). We would like to thank
Bryn Hubbard for his help with the classication of the basal ice facies
encountered, the logistical staff at UNIS for their help in the collection of ice
cores and preparation (Gerd Irene Sigernes, Martin Indreiten, and Monica
Votvik), the other students on the UNIS AG-325 Glaciology course who were
also present during the initial inspection of the sections in 2011 and Kathrin
Naegeli and Philipp Schuppli for their assistance in the 2012 season.
Finally, we would like to thank the two reviewers, Neal Iverson and
Peter Knight, for their comments and suggestions and Jeremy Bassis for
his editorial assistance.
Alsop, G. I., and J. Carreras (2007), The structural evolution of sheath folds:
A case study from Cap de Creus, J. Struct. Geol.,29(12), 19151930,
Alsop, G. I., and R. E. Holdsworth (2004), The geometry and topology of
natural sheath folds: A new tool for structural analysis, J. Struct. Geol.,
26(9), 15611589, doi:10.1016/j.jsg.2004.01.009.
Alsop, G. I., R. E. Holdsworth, and K. J. W. McCaffrey (2007), Scale invari-
ant sheath folds in salt, sediments and shear zones, J. Struct. Geol.,29(10),
15851604, doi:10.1016/j.jsg.2007.07.012.
Bader, H. (1951), Introduction to ice petrofabrics, The Journal of Geology,
59(6), 519536.
Bamber, J. (1987), Internal reecting horizons in Spitsbergen glaciers, Ann.
Benn, D. I., and D. J. A. Evans (2010), Glaciers and Glaciation, 2nd ed.,
816 pp., Hodder Education.
Biedermann, A. R., W. Lowrie, and A. M. Hirt (2013), A method for improving
the measurement of low-eld magnetic susceptibility anisotropy in weak sam-
ples, J. Appl. Geophys.,88(0), 122130, doi:10.1016/j.jappgeo.2012.10.008.
Borradaile, G. J. (1988), Magnetic susceptibility, petrofabrics and strain,
Tectonophysics,156(1-2), 120, doi:10.1016/0040-1951(88)90279-X.
Borradaile, G. J. (1991), Correlation of strain with anisotropy of magnetic sus-
ceptibility (AMS), PAGEOPH,135(1), 1529, doi:10.1007/bf00877006.
Borradaile, G. J., and M. Jackson (Eds.) (2004), Anisotropy of Magnetic
Susceptibility (AMS): Magnetic Petrofabrics of Deformed Rocks,
299360 pp., Geological Society, London, Special Publications, doi:10.1144/
Borradaile, G. J., B. S. G. Almqvist, and I. Geneviciene (2012), Anisotropy
of magnetic susceptibility (AMS) and diamagnetic fabrics in the Durness
Limestone, NW Scotland, J. Struct. Geol.,34(0), 5460, doi:10.1016/j.
Callot, J. P., P. Robion, W. Sassi, M. L. E. Guiton, J. L. Faure, J. M. Daniel,
J. M. Mengus, and J. Schmitz (2010), Magnetic characterisation of folded
aeolian sandstones: Interpretation of magnetic fabrics in diamagnetic rocks,
Tectonophysics,495(34), 230245, doi:10.1016/j.tecto.2010.09.020.
Castelnau, O., H. Shoji, A. Mangeney, H. Milsch, P. Duval, A. Miyamoto,
K. Kawada, and O. Watanabe (1998), Anisotropic behavior of GRIP ices
and ow in Central Greenland, Earth Planet. Sci. Lett.,154(14), 307322,
Cifelli, F., M. Mattei, M. Chadima, A. M. Hirt, and A. Hansen (2005), The
origin of tectonic lineation in extensional basins: Combined neutron tex-
ture and magnetic analyses on undeformedclays, Earth Planet. Sci.
Lett.,235(12), 6278, doi:10.1016/j.epsl.2005.02.042.
Cifelli, F., M. Mattei, M. Chadima, S. Lenser, and A. M. Hirt (2009), The
magnetic fabric in undeformed clays: AMS and neutron texture analyses
from the Rif Chain (Morocco), Tectonophysics,466(12), 7988, doi:10.1016/
Cobbold, P. R., and H. Quinquis (1980), Development of sheath folds
in shear regimes, J. Struct. Geol.,2(12), 119126, doi:10.1016/0191-
Conway, H., G. Catania, C. Raymond, A. Gades, T. Scambos, and H. Engelhardt
(2002), Switch of ow direction in an Antarctic ice stream, Nature,419(6906),
465467, doi:10.1038/nature01081.
Cook, S. J., R. I. Waller, and P. G. Knight (2006), Glaciohydraulic supercooling:
The process and its signicance, Prog. Phys. Geogr.,30(5), 577588,
Cutbill, J. L., and A. Challinor (1965), Revision of the stratigraphical scheme
for the carboniferous and permian rocks of Spitsbergen and Bjørnøya,
Geol. Mag.,102(05), 418439, doi:10.1017/S0016756800053693.
Dallmann, W. K., K. Piepjohn, D. Blomeier, and S. Elvevold (2009),
Geological map of Svalbard 1:100,000, sheet C8G Billefjorden, Norsk
Polarinstitutt, Tromsø.
Dallmann,W. K., K. Piepjohn, G. P.Halverson, S. Elvevold, and D. Blomeier
(2011), Geological map Svalbard 1:100 000, sheet D6G Vaigattbogen,
Norsk Polarinstitutt, Tromsø.
de Wall, H., M. Bestmann, and K. Ullemeyer (2000), Anisotropy of diamag-
netic susceptibility in Thassos marble: A comparison between measured
and modeled data, J. Struct. Geol.,22(1112), 17611771, doi:10.1016/
Dunlop, D. J., and O. Özdemir (1997), Rock Magnetism: Fundamentals and
Frontiers, 596 pp., Cambridge Univ. Press, Cambridge.
Ellwood, B. B., and M. T. Ledbetter (1977), Antarctic bottom water uctua-
tions in the Vema Channel: Effects of velocity changes on particle align-
ment and size, Earth Planet. Sci. Lett.,35(2), 189198, doi:10.1016/
Evans, D. J. A., and D. I. Benn (2004), A Practical Guide to the Study of
Glacial Sediments, 266 pp., Arnold London, London.
Eyles, N., T. E. Day, and A. Gavican (1987), Depositional controls on the
magnetic characteristics of lodgement tills and other glacial diamict facies,
Can. J. Earth Sci.,34, 24362458.
Ferré, E. C. (2002), Theoretical models of intermediate and inverse AMS
fabrics, Geophys. Res. Lett.,29(7), 3131, doi:10.1029/2001GL014367.
Fleming, E. J., C. T. E. Stevenson, and M. S. Petronis (2013), New insights
into the deformation of a Middle Pleistocene glaciotectonised sequence in
Norfolk, England through magnetic and structural analysis, Proc. Geol.
Assoc.,124, doi:10.1016/j.pgeola.2012.11.004, in press.
Fritz, M., S. Wetterich, H. Meyer, L. Schirrmeister, H. Lantuit, and
W. H. Pollard (2011), Origin and characteristics of massive ground ice
on Herschel Island (western Canadian Arctic) as revealed by stable water
isotope and hydrochemical signatures, Permafrost and Periglacial Processes,
22(1), 2638, doi:10.1002/ppp.714.
Gentoso, M. J., E. B. Evenson, K. P. Kodama, N. R. Iverson, R. B. Alley,
C. Berti, and A. Kozlowski (2012), Exploring till bed kinematics
using AMS magnetic fabrics and pebble fabrics: The Weedsport drumlin
eld, New York State, USA, Boreas,41(1), 3141, doi:10.1111/j.1502-
Guerrero-Suarez, S., and F. Martín-Hernández (2012), Magnetic anisotropy
of hematite natural crystals: Increasing low-eld strength experiments, Int.
J. Earth Sci.,101(3), 625636, doi:10.1007/s00531-011-0666-y.
Hambrey, M. J., and N. F. Glasser (2003), The role of folding and foliation
development in the genesis of medial moraines: Examples from Svalbard
glaciers, The Journal of Geology,111(4), 471485.
Hambrey, M. J., and W. Lawson (Eds.) (2000), Structural Styles and
Deformation Fields in Glaciers: A Review,5983 pp., Geological Society
London Special Publications, doi:10.1144/GSL.SP.2000.176.01.06.
Hambrey, M. J., T. Murray, N. F. Glasser, A. Hubbard, B. Hubbard,
G. Stuart, S. Hansen, and J. Kohler (2005), Structure and changing dy-
namics of a polythermal valley glacier on a centennial timescale: Midre
Lovénbreen, Svalbard, J. Geophys. Res.,110, F01006, doi:10.1029/
Hargraves, R. B., D. Johnson, and C. Y. Chan (1991), Distribution anisot-
ropy: The cause of AMS in igneous rocks?, Geophys. Res. Lett.,18(12),
21932196, doi:10.1029/91GL01777.
Harland, W. B., L. M. Anderson, D. Manasrah, and N. J. Buttereld (1997),
The Geology of Svalbard, Geological Society Publishing House.
Hodgkins, R., and J. A. Dowdeswell (1994), Tectonic processes in Svalbard
tide-water glacier surges: Evidence from structural glaciology, J. Glaciol.,
40(136), 553560.
Hooyer, T. S., N. R. Iverson, F. Lagroix, and J. F. Thomason (2008),
Magnetic fabric of sheared till: A strain indicator for evaluating the bed de-
formation model of glacier ow, J. Geophys. Res.,113, F02002,
Hrouda, F. (Ed.) (2004), Problems in Interpreting AMS Parameters in
Diamagnetic Rocks,4959 pp., Geological Society, London, Special
Publications, doi:10.1144/gsl.sp.2004.238.01.05.
Hrouda, F., and A. Kapička (1986), The effect of quartz on the magnetic an-
isotropy of quartzite, Stud. Geophys. Geod.,30(1), 3945.
Hrouda, F., M. Chlupácová, and S. Mrázová (2006), Low-eld varia-
tion of magnetic susceptibility as a tool for magnetic mineralogy of
rocks, Phys.EarthPlanet.Inter.,154(3-4), 323336, doi:10.1016/j.
Hubbard, B. (1991), Freezing-rate effects on the physical characteristics of
basal ice formed by net adfreezing, J. Glaciol.,37(127).
Hubbard, B., and M. Sharp (1989), Basal ice formation and defor-
mation: A review, Prog. Phys. Geogr.,13(4), 529558, doi:10.1177/
Hubbard, B., and M. Sharp (1993), Weertman regelation, multiple refreezing
events and the isotopic evolution of the basal ice layer, J. Glaciol.,
39(132), 275291.
Hubbard, B., and M. Sharp (1995), Basal ice facies and their formation in the
western Alps, Arct. Alp. Res.,27, 301310.
Hubbard, B., J.-L. Tison, L. Janssens, and B. Spiro (2000), Ice-core evi-
dence of the thickness and character of clear-facies basal ice: Glacier
de Tsaneuron, Switzerland, J. Glaciol.,46(152), 140150, doi:10.3189/
Hubbard, B., S. Cook, and H. Coulson (2009), Basal ice facies: A review and
unifying approach, Quat. Sci. Rev.,28(1920), 19561969, doi:10.1016/j.
Iverson, N. R., T. S. Hooyer, J. F. Thomason, M. Graesch, and J. R. Shumway
(2008), The experimental basis for interpreting particle and magnetic fabrics
of sheared till, Earth Surf. Processes Landforms,33(4), 627645,
Jeffery, G. B. (1922), The motion of ellipsoidal particles immersed in a
viscous uid, in Proceedings of the Royal Society of London.Series
A,Containing Papers of a Mathematical and Physical Character,
102(715), 161-179.
Jelínek, V. (1981), Characterization of the magnetic fabric of rocks,
Tectonophysics,79(34), T63T67, doi:10.1016/0040-1951(81)90110-4.
Khan, M. A. (1962), The anisotropy of magnetic susceptibility of some igne-
ous and metamorphic rocks, J. Geophys. Res.,67(7), 28732885,
Knight, P. G. (1997), The basal ice layer of glaciers and ice sheets, Quat. Sci.
Rev.,16(9), 975993, doi:10.1016/s0277-3791(97)00033-4.
Lagroix, F., and S. K. Banerjee (2002), Paleowind directions from the mag-
netic fabric of loess proles in central Alaska, Earth Planet. Sci. Lett.,
195(1-2), 99112, doi:10.1016/S0012-821X(01)00564-7.
Lanci, L., D. V. Kent, P. E. Biscaye, and A. Bory (2001), Isothermal rema-
nent magnetization of Greenland ice: Preliminary results, Geophys. Res.
Lett.,28(8), 16391642, doi:10.1029/2000GL012594.
Larsen, N. K., C. Kronborg, J. C. Yde, and N. T. Knudsen (2010), Debris
entrainment by basal freeze-on and thrusting during the 19951998 surge
of Kuannersuit Glacier on Disko Island, west Greenland, Earth Surf.
Processes Landforms,35(5), 561574, doi:10.1002/esp.1945.
Lawson, W. J., M. J. Sharp, and M. J. Hambrey (1994), The structural
geology of a surge-type glacier, J. Struct. Geol.,16(10), 14471462,
Liss, D., D. H. W. Hutton, and W. H. Owens (2002), Ropy ow structures: A
neglected indicator of magma-ow direction in sills and dikes, Geology,
30(8), 715, doi:10.1130/0091-7613(2002).
Maltman, A. J., B. Hubbard, and M. J. Hambrey (2000), Deformation of gla-
cial materials: Introduction and overview, Geological Society London
Special Publications,176(1), 1, doi:10.1144/GSL.SP.2000.176.01.01.
Mansell, D., A. Luckman, and T. Murray (2012), Dynamics of tidewater
surge-type glaciers in northwest Svalbard, J. Glaciol.,58(207), 110118,
March, A. (1932), Mathematical theory on regulation according to the parti-
cle shape and afne deformation, Z. Kristall.,81(3/4), 285297.
Murray, T., T. Strozzi, A. Luckman, H. Jiskoot, and P. Christakos (2003), Is
there a single surge mechanism? Contrasts in dynamics between glacier
surges in Svalbard and other regions, J. Geophys. Res.,108(B5), 2237,
Nagata, T. (1961), Rock magnetism, Maruzen, Tokyo.
Neves, S. P., J. M. R. da Silva, and G. Mariano (2005), Oblique lineations in
orthogneisses and supracrustal rocks: Vertical partitioning of strain in a hot
crust (eastern Borborema Province, NE Brazil), J. Struct. Geol.,27(8),
15131527, doi:10.1016/j.jsg.2005.02.002.
Owens, W. H., and E. H. Rutter (1978), The development of magnetic sus-
ceptibility anisotropy through crystallographic preferred orientation in a
calcite rock, Phys. Earth Planet. Inter.,16(3), 215222, doi:10.1016/
Parés, J. M., and B. A. van der Pluijm (2002), Evaluating magnetic lineations
(AMS) in deformed rocks, Tectonophysics,350(4), 283298, doi:10.1016/
Parés, J. M., B. A. van der Pluijm, and J. Dinarès-Turell (1999), Evolution
of magnetic fabrics during incipient deformation of mudrocks (Pyrenees,
northern Spain), Tectonophysics,307(1-2), 114, doi:10.1016/S0040-
Ramsay, J. G., and M. I. Huber (1983), The Techniques of Modern
Structural Geology, Volume 1: Strain Analysis, Academic Press,
Richter, C., and B. A. van der Pluijm (1994), Separation of paramagnetic and
ferrimagnetic susceptibilities using low temperature magnetic susceptibil-
ities and comparison with high eld methods, Phys. Earth Planet. Inter.,
82(2), 113123, doi:10.1016/0031-9201(94)90084-1.
Rigsby, G. P. (1958), Fabrics of glacier and laboratory deformed ice,
paper presented at Symposium on Physics of the Movement of the Ice,
Symposium of Chamonix.
Rochette, P. (1987), Magnetic susceptibility of the rock matrix related to
magnetic fabric studies, J. Struct. Geol.,9(8), 10151020, doi:10.1016/
Samyn, D., A. Svensson, and S. J. Fitzsimons (2008), Dynamic
implications of discontinuous recrystallization in cold basal ice:
Taylor Glacier, Antarctica, J. Geophys. Res.,113, F03S90,
Samyn, D., S. Fitzsimons, and R. Lorrain(2010), Rotating micro-structures in
Antarctic cold basal ice: Implicationsfor glacier ow and its interpretation,
Int. J. Earth Sci.,99(8), 18491857, doi:10.1007/s00531-009-0478-5.
Sharp, M., W. Lawson, and R. S. Anderson (1988), Tectonic processes in a
surge-type glacier, J. Struct. Geol.,10(5), 499515, doi:10.1016/0191-
Shumway, J. R., and N. R. Iverson (2009), Magnetic fabrics of the Douglas
Till of the Superior lobe: Exploring bed-deformation kinematics, Quat.
Sci. Rev.,28(1-2), 107119, doi:10.1016/j.quascirev.2008.09.020.
Souchez, R., G. Vandenschrick, R. Lorrain, and J. L. Tison (2000), Basal ice
formation and deformation in central Greenland: A review of existing and
new ice core data, Geological Society London Special Publications,
176(1), 13, doi:10.1144/GSL.SP.2000.176.01.02.
Tarling, D. H., and F. Hrouda (1993), The Magnetic Anisotropy of Rocks,
217 pp., Chapman & Hall, London.
Thomason, J. F., and N. R. Iverson (2006), Microfabric and microshear evo-
lution in deformed till, Quat. Sci. Rev.,25(9-10), 10271038.
Thomason, J. F., and N. R. Iverson (2009), Deformation of the Batestown till
of the Lake Michigan lobe, Laurentide ice sheet, J. Glaciol.,55(189),
131146, doi:10.3189/002214309788608877.
Tison, J. L., and R. Lorrain (1987), A mechanism of basal ice-layer forma-
tion involving major ice-fabric changes, J. Glaciol.,33(113), 4750.
Tison, J. L., T. Thorsteinsson, R. D. Lorrain, and J. Kipfstuhl (1994), Origin
and development of textures and fabrics in basal ice at Summit, Central
Greenland, Earth Planet. Sci. Lett.,125(14), 421437, doi:10.1016/
Twiss, R. J., and E. M. Moores (1992), Structural Geology, 532 pp., WH
Freeman, New York.
Waller, R. I., J. B. Murton, and P. G. Knight (2009), Basal glacier ice and mas-
sive ground ice: Different scientists, same science?, Geological Society,
London, Special Publications,320(1), 5769, doi:10.1144/sp320.5.
principles, and applications of automated ice fabric analyzers, Microsc.
Res. Tech.,62(1), 218, doi:10.1002/jemt.10380.
Wilson, C. J. L. (2000), Experimental work on the effect of pre-existing an-
isotropy on fabric development in glaciers, Geological Society London
Special Publications,176(1), 97, doi:10.1144/GSL.SP.2000.176.01.08.
Wilson, C. J. L., and M. Peternell (2011), Evaluating ice fabrics using fabric
analyser techniques in Sorsdal Glacier, East Antarctica, J. Glaciol.,
57(205), 881894, doi:10.3189/002214311798043744.
Wilson, C. J. L., and H. M. Sim (2002), The localization of strain and c-axis
evolution in anisotropic ice, J. Glaciol.,48(163), 601610, doi:10.3189/
... Evolving from this technique is the relatively novel technique of determining the "anisotropy of magnetic susceptibility," which involves measuring the orientation of magnetic particles in small drill cores, and plotting the data on stereographic projections. The technique is useful for examining the deformation processes associated with glacier motion, including the development of lineations and folds (Fleming et al., 2013). The importance of determining strain-rate tensors was recognized as important for understanding how some structures, such as foliation and crevasses, formed. ...
... During thrust-faulting, the maximum (σ 1 ) and intermediate (σ 2 ) principal stress tensors are in the horizontal plane, oriented perpendicular and parallel to the fracture respectively, with the minimum (σ 3 ) tensor oriented vertically ( Figure 12b). Movement is accommodated by a dip-slip relationship; however, it is worth noting that several studies have observed thrust-faults that cannot be solely explained by dip-slip movement, but also have a component of strike-slip, suggesting a transpressional stress regime (Fleming et al., 2013;Lovell, Fleming, Benn, Hubbard, Lukas, & Naegeli, 2015). ...
... Thrust-faulting has also been recognized in polythermal tidewater glaciers where a surge front has passed through a glacier, reflecting a progressive change in the thermal boundary conditions from cold-to wetbased (Fleming et al., 2013;King et al., 2016;Lovell, Fleming, Benn, Hubbard, Lukas, & Naegeli, 2015;Murray et al., 1997Murray et al., , 2000. ...
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The aims of this review are to: (i) describe and interpret structures in valley glaciers in relation to strain history; and (ii) to explore how these structures inform our understanding of the kinematics of large ice masses, and a wide range of other aspects of glaciology. Structures in glaciers give insight as to how ice deforms at the macroscopic and larger scale. Structures also provide information concerning the deformation history of ice masses over centuries and millennia. From a geological perspective, glaciers can be considered to be models of rock deformation, but with rates of change that are measurable on a human time-scale. However, structural assemblages in glaciers are commonly complex, and unravelling them to determine the deformation history is challenging; it thus requires the approach of the structural geologist. A wide range of structures are present in valley glaciers: (i) primary structures include sedimentary stratification and various veins; (ii) secondary structures that are the result of brittle and ductile deformation include crevasses, faults, crevasse traces, foliation, folds, and boudinage structures. Some of these structures, notably crevasses, relate well to measured strain-rates, but to explain ductile structures analysis of cumulative strain is required. Some structures occur in all glaciers irrespective of size, and they are therefore recognizable in ice streams and ice shelves. Structural approaches have wide (but as yet under-developed potential) application to other sub-disciplines of glaciology, notably glacier hydrology, debris entrainment and transfer, landform development, microbiological investigations, and in the interpretation of glacier-like features on Mars.
... Investigation of glaciological structures such as crevasses, shear planes, foliation and medial moraines have increased understanding of cumulative deformation and variable stress and strain regimes experienced during different phases of the surge cycle (e.g. Sharp and others, 1988;Hodgkins and Dowdeswell, 1994;others, 1994, 2000;Lawson, 1996Lawson, , 1997Hambrey and Dowdeswell, 1997;Fleming and others, 2013;Hudleston, 2015;King and others, 2016;Sevestre and others, 2018;Hambrey and Clarke, 2019;Young and others, 2022). ...
... Three-dimensional orientation measurements (strike and dip) of foliation and englacial debris-rich structures (debris layers) were collected using a compass clinometer. Measurements of linear structure (dip and dip direction) were collected from sheared debris laminae, or mineral stretching lineations, within the debris layers (Fleming and others, 2013). Structural data were corrected for magnetic deviation and plotted as equal-area stereographic projections using Stereo32 software (Röller and Trepmann, 2008). ...
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We assess the evolution of glaciological structures during the 2003–05 surge in the Paulabreen glacier system, Svalbard. Glaciological structures on the glacier surface were mapped using aerial photographs captured in the early stages of the surge (2003) and 5 years after surge termination (2011). Three-dimensional measurements of glaciological structures were collected at the tidewater front in 2013. These datasets document the physical changes during (1) the late quiescent phase; (2) the early phase of the surge as the surge front propagated down Skobreen and advanced into Paulabreen and (3) the final stages of the surge following the surge front reaching the glacier terminus. Crevasse patterns and clusters of arcuate shear planes record zones of compressive and extensional flow associated with the downglacier progression of the surge front. The transfer of surging ice from Skobreen into Paulabreen caused lateral displacement of the medial moraines to the northeast. At the ice front, this movement tilted glaciological structures in the same direction. Structures at the southwest margin record strike–slip faulting and the elevation of debris into the ice in a zone of compression and transpression. We summarise these observations in a schematic reconstruction of structural evolution during the surge.
... Many papers show that magnetic fabric provides a reliable strain indicator (Borradaile and Jackson 2004;Ferré et al. 2004;Hrouda 1979Hrouda , 2002. AMS characteristics have been investigated in Late Pleistocene-Holocene sediments in China (Liu et al. 2005), in Pleistocene tills of drumlin fields in the USA (Gentoso et al. 2012;Hopkins et al. 2016;Iverson 2017), and in basal ice of a surge-type glacier (Fleming et al. 2013). ...
... This can be the result of two stages of MSGL formation: relics of primary (regional) direction and glacial erosion by groove-ploughing. AMS provided specific useful information regarding the kinematics of deformation within subglacially deformed sediments (Fleming et al. 2013). Investigation of modern-day sedimentary environments on the glacier margin provides important information on the sediments and relief generated by retreating glaciers. ...
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Till macro-and microfab-rics of mega-scale glacial lineations of Mūša-Nemunėlis Lowland, north Lithuania. Baltica, 33 (1), 85-96. Vilnius. North Lithuania was chosen for a study of the "drumlinised" morainic surface produced during the Last Glaciation, typified as well-expressed mega-scale glacial lineations (MGSLs). The goal pursued in the present study was to investigate the morphology and macro-and microfabrics of some large glacial lineations to substantiate their formation mechanism. The geological structure of Quaternary strata of an area and the erosion depression of sub-Quaternary surface suggest favourable conditions for the glacier to rapidly fluctuate into the area during deglaciation of Late Glaciation. Investigations of Pleistocene tills observed in the MSGLs of the area preserved on the eastern and western margins of the study area show that these deposits are formed from the upper part of the Baltija Subformation-Middle Lithuanian till. According to two sets of grain sizes, MSGL tills are often notable for increased values of relative entropy. Therefore, morainic material deposited during the redeposition of the Baltija Subformation till was thoroughly mixed. The data on orientation and inclination of long axes of gravel and pebbles in the tills that form MSGLs, as well as the anisotropy of magnetic susceptibility (AMS) of microclast material suggest that the formation of MSGLs may have been influenced by directions of the local glacial stress that are different from the regional direction of glacial motion (about N-S). The change of macro-and microfabric of till confirms the formation of MSGLs during glacier erosion by groove-ploughing from the Baltija Subformation till. This occurred when basal ice carried over clast material to MSGL crests from interridge areas.
... Many papers show that magnetic fabric provides a reliable strain indicator (Borradaile and Jackson 2004;Ferré et al. 2004;Hrouda 1979Hrouda , 2002. AMS characteristics have been investigated in Late Pleistocene-Holocene sediments in China (Liu et al. 2005), in Pleistocene tills of drumlin fields in the USA (Gentoso et al. 2012;Hopkins et al. 2016;Iverson 2017), and in basal ice of a surge-type glacier (Fleming et al. 2013). ...
... This can be the result of two stages of MSGL formation: relics of primary (regional) direction and glacial erosion by groove-ploughing. AMS provided specific useful information regarding the kinematics of deformation within subglacially deformed sediments (Fleming et al. 2013). Investigation of modern-day sedimentary environments on the glacier margin provides important information on the sediments and relief generated by retreating glaciers. ...
Full-text available
Till macro-and microfab-rics of mega-scale glacial lineations of Mūša-Nemunėlis Lowland, north Lithuania. Baltica, 33 (1), 85-96. Vilnius. North Lithuania was chosen for a study of the "drumlinised" morainic surface produced during the Last Glaciation, typified as well-expressed mega-scale glacial lineations (MGSLs). The goal pursued in the present study was to investigate the morphology and macro-and microfabrics of some large glacial lineations to substantiate their formation mechanism. The geological structure of Quaternary strata of an area and the erosion depression of sub-Quaternary surface suggest favourable conditions for the glacier to rapidly fluctuate into the area during deglaciation of Late Glaciation. Investigations of Pleistocene tills observed in the MSGLs of the area preserved on the eastern and western margins of the study area show that these deposits are formed from the upper part of the Baltija Subformation-Middle Lithuanian till. According to two sets of grain sizes, MSGL tills are often notable for increased values of relative entropy. Therefore, morainic material deposited during the redeposition of the Baltija Subformation till was thoroughly mixed. The data on orientation and inclination of long axes of gravel and pebbles in the tills that form MSGLs, as well as the anisotropy of magnetic susceptibility (AMS) of microclast material suggest that the formation of MSGLs may have been influenced by directions of the local glacial stress that are different from the regional direction of glacial motion (about N-S). The change of macro-and microfabric of till confirms the formation of MSGLs during glacier erosion by groove-ploughing from the Baltija Subformation till. This occurred when basal ice carried over clast material to MSGL crests from interridge areas.
... The occurrence of this roughly east-west oriented vector is noteworthy as it also occurs in different subsamples from other outcrops as secondary overprint or ChRM (see supplement)). However, the cause of these conspicuous east-west directions is not conclusively analysed in this study nor in that of Graf (2019), as this would have required more detailed AMS measurements (e.g., Fleming et al., 2013), which would not have fitted into the scope of the two studies. Another aspect in which this study diverge from the interpretation of Graf (2019) is the origin of the haematite that is identified as an important carrier of the remanence in both studies. ...
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The Deckenschotter is a fluvial to glaciofluvial gravel unit in northern Switzerland and southern Germany. The depos- its are considered the oldest preserved glacial to interglacial Quaternary deposits in the northern Alpine foreland and are thus important geomorphological markers for landscape evolution. Nevertheless, the age of the deposits is only approximately known and subject to controversial debates. This study presents the results of an extensive palaeomagnetic investigation carried out on intercalated fine-grained sediments at 11 sites of the Höhere Deckens- chotter (HDS) and at 5 sites of the Tiefere Deckenschotter (TDS). The HDS show reversed and normal magnetisations, indicating deposition > 0.773 Ma, while the TDS exhibit only normal directions. Age constraints for the different sites are discussed in the light of evidence from other studies. The study therefore clearly supports the efforts to determine the age of the Deckenschotter. As data from previous palaeomagnetic studies on the HDS and TDS have not been published or preserved, this is in fact the only data-based palaeomagnetic study available.
... However, these fractures show small displacements, basal debris entrainment, and do not appear to be derived from pre-existing crevasses traces. They are therefore interpreted as thrust-faults in their own right, as documented in glaciers in the Canadian Arctic (Zdanowicz et al. 1996), Svalbard (Hambrey et al. 1999;Fleming et al. 2013;Lovell et al. 2015), ...
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Using satellite remote sensing, this study aims to assess the validity of upscaling ground‐based structural observations of small valley glaciers, to larger‐scale ice masses that are too vast or inaccessible for field‐study or ground‐truthing. Focusing on four adjacent valley glaciers on Bylot Island, Nunavut, Arctic Canada, we establish that ground‐based structural observations from two smaller (Stagnation and Fountain glaciers) can be used to interpret the structures visible in optical satellite imagery in two much larger glaciers (Aktineq and Sermilik glaciers). All the glaciers investigated have prominent longitudinal lineations, which are interpreted from ground observations to be longitudinal foliation. Other structures that were identified include primary stratification, crevasses, crevasse traces, and thrust‐faults. Strong longitudinal foliation is concentrated at flow‐unit boundaries, with differential ablation of ice facies commonly resulting in a ridge‐and‐furrow supraglacial topography that controls supraglacial streams and debris concentrations. Consequently, areas of strong foliation appear darker than areas of weak foliation in satellite imagery. As coarser resolution imagery is utilised to map large‐scale ice masses, sub‐pixel structural information is lost. Individual lineations mapped in coarser resolution imagery therefore probably comprise groups of clustered foliation at the sub‐pixel scale. Lateral narrowing measurements and calculated one‐dimensional strain across zones of longitudinal foliation are assessed as a tool for identifying large‐scale surface strain patterns, in particular large‐scale pure shear regimes. These one‐dimensional strain measurements suggest that flow‐unit boundaries are areas that undergo considerable cumulative strains. The upscaling approach used here can be applied to the largest ice masses, notably the Antarctic Ice Sheet.
... Between 1990 and 2010, the growth of the bulge led to an increase in surface slope, with the largest increase occurring near the bulge front. Surface steepening at the glacier front is a Dowdeswell and others (1991); b Rolstad and others (1997); c Luckman and others (2002); d Strozzi and others (2002); e Mansell and others (2012); f Murray and others (2003b); g Murray and others (2003a); h Murray and others (2012); i Dowdeswell and Benham (2003); j Fleming and others (2013); k Flink and others (2015); l Sevestre and others (2018); m Sund and others (2014); n Nuth and others (2019); o Dunse and others (2015); p Strozzi and others (2017); q Benn and others (2019); r Ottesen and others (2017); s Ottesen and others (2008); t Burton and others (2016). ...
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Negribreen, a tidewater glacier located in central eastern Svalbard, began actively surging after it experienced an initial collapse in summer 2016. The surge resulted in horizontal surface velocities of more than 25 m d ⁻¹ , making it one of the fastest-flowing glaciers in the archipelago. The last surge of Negribreen likely occurred in the 1930s, but due to a long quiescent phase, investigations of this glacier have been limited. As Negribreen is part of the Negribreen Glacier System, one of the largest glacier systems in Svalbard, investigating its current surge event provides important information on surge behaviour among tidewater glaciers within the region. Here, we demonstrate the surge development and discuss triggering mechanisms using time series of digital elevation models (1969–2018), surface velocities (1995–2018), crevasse patterns and glacier extents from various data sources. We find that the active surge results from a four-stage process. Stage 1 (quiescent phase) involves a long-term, gradual geometry change due to high subglacial friction towards the terminus. These changes allow the onset of Stage 2, an accelerating frontal destabilization, which ultimately results in the collapse (Stage 3) and active surge (Stage 4).
Glacier surging provides a window into the processes responsible for some of the fastest ice flow on Earth. Surge-type glaciers are festooned with fold trains—many visible from space—which encode the history of polyphase deformation associated with this form of episodic fast flow. We conduct the first investigation of the kinematic evolution of these kilometre-scale folds using a full-Stokes numerical ice-flow model. We model the folds through multiple surge cycles within a set of synthetic glacier confluence configurations, and identify how differences in glacier flow regimes imprint themselves on three-dimensional fold geometry. Based on simulation results across parameter space, we present an archetype of kinematic evolution that describes the transition from cylindrical vertically plunging gentle folds emplaced during the surge phase, to complex depth-varying folds following multiple cycles of surging and quiescent flow. A detailed examination of the surface trace of these fold highlights the links between glacier flow regime and folding. The initial fold geometry is controlled by longitudinal and lateral shear stress regimes during surging, while fold evolution is governed primarily by lateral shearing after emplacement. This reflects the influence of valley geometry and glacier dynamics on the variability of flow regimes during both surging and quiescent flow. Finally, we illustrate the potential of our approach to reconstruct more complex fold geometries as observed in nature.
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This study examines how longitudinal foliation develops in glaciers and ice sheets in a wide range of topographic, climatic, and dynamic settings, at a variety of spatial scales. Study locations include four valley glaciers in Svalbard (Austre Brøggerbreen, Midtre Lovénbreen, Austre Lovénbreen, and Pedersenbreen), a valley glacier in Canada (Sermilik Glacier), and seven outlet glaciers in Antarctica (Hatherton Glacier, Taylor Glacier, Ferrar Glacier, Lambert Glacier, Recovery Glacier, Byrd Glacier, and Pine Island Glacier). Detailed structural mapping of the valley glaciers from satellite imagery and field-based measurements were used to document the formation of longitudinal foliation in small-scale ice masses. These findings were ‘up-scaled’ and applied to much larger glaciers and ice streams. Longitudinal foliation develops in concentrated bands at flow unit boundaries as a result of enhanced simple shear. However, longitudinal foliation is not directly observable from satellite imagery at the surface of larger-scale valley glaciers. The longitudinal structures visible at the surface of larger-scale glaciers form at flow-unit boundaries and are composed of bands of steeply dipping longitudinal foliation; however, they appear as individual linear features on satellite imagery as a result of the comparatively low spatial resolution of the imagery. The persistence of flowlines in the Antarctic Ice Sheet through areas of crevassing and net ablation (blue-ice areas) suggests that they are the surface representation of a three-dimensional structure. Flowlines are therefore inferred to be the surface expression of flow-unit boundaries composed of bands of steeply dipping longitudinal foliation. The survival and deformation of flowlines in areas of ice flow stagnation indicates that flowlines form in their initiation zones and not along their entire length. Furthermore, these ice stagnation areas indicate that flowlines record past ice dynamics and switches in ice flow.
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The analysis of low‐field anisotropy of magnetic susceptibility (AMS) was used to reconstruct the subglacial deposition conditions during the Main Stadial of the Odranian Glaciation (MIS 6) in till deposits from a site in Dębe (central Poland). Based on the AMS parameters, six till beds were identified (intervals 1–6). The declination of the maximum magnetic susceptibility axis (k1) indicates that the ice sheet was moving in from the northwest. The obtained results confirm the thesis about the preferred direction of ice‐sheet transgression during the Odranian Glaciation (MIS 6) in this part of Poland. This interpretation is also confirmed by data obtained from measurements of the long axis of clasts, which agree with the orientation of k1. Based on the AMS results, a significant part of the profile was deformed through simple shear and direct interaction of the ice sheet with the underlying sediment (beds 2–5). The lowest part of the till (bed 6) may have been deposited on a southeast‐trending slope or post‐depositional deformed by uneven loading of the ice cover. The upper part of the profile (especially in interval 1) could be deposited with an impact of pore water.
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A simple model is developed to simulate the isotopic fractionation which accompanies Weertman regelation at the bed of temperate-based glaciers. The fractionation equations of Jouzel and Souchez (1982) are applied to multiple refreezing events over measured glacier-bed profiles, and mass balance is maintained as the basal ice and meltwater produced at one bedrock hummock enter the next. Simulation results indicate that undeformed regelation ice layers are on the order of millimetres to centimetres thick, often being completely melted at the stoss face of certain hummocks and exceptionally reaching a thickness in excess of 10 cm. Neither the internal morphology nor the isotopic composition of these layers is constant, but both vary down-glacier in accordance with bedrock configuration. A glacier-wide fractionation process is identified whereby heavy isotopes are preferentially removed from the basal meltwater film and incorporated into the basal ice. This process might go some way to explaining the anomalously “light” isotopic composition measured in base-flow waters leaving some glaciers. Vertical isotope profiles through undeformed basal ice layers are reconstructed and show that significant isotopic excursions can occur at a scale of millimeteres, while the range of isotopic compositions within such multi-layered regelation ice is greater than that which would occur in ice produced by a single refreezing event. In circumstances where the regelation system is disrupted by removal of film waters into a network of linked cavities, it is found that the remaining basal ice may be significantly enriched in heavy isotopes relative to the composition of the initial mass inputs to the system. Heavy isotope enrichment of this magnitude and consideration of the thickness of the basal ice layers concerned may explain the absence of recorded basal ice samples heavy enough to have been formed in equilibrium with subglacial precipitates sampled at one of the sites (Glacier de Tsanfleuron) and reported in an earlier paper.
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A number of theoretical and empirical studies have indicated that many individual characteristics of ice formed by processes of net basal adfreezing may be sensitive to the rate of propagation of the freezing front through the reservoir concerned. The effects of freezing rate on the the stable-isotope chemistry and crystallography of ice, in addition to the disposition and character of included debris and gas are reported. Unidirectional freezing through a cylindrical reservoir containing various water–sediment mixtures has been conducted in the laboratory and the resulting cores analysed for debris and gas disposition and ice-crystal size and fabric. The data lend support to inferences drawn from studies concerned with specific ice properties and an idealized suite of characteristics is developed which may be diagnostic of basal ice formed by net adfreezing.
Rock Magnetism, first published in 1997, is a comprehensive treatment of fine particle magnetism and the magnetic properties of rocks. Starting from atomic magnetism and magnetostatic principles, the authors explain why domains and micromagnetic structures form in ferromagnetic crystals and how these lead to magnetic memory in the form of thermal, chemical and other remanent magnetizations. The phenomenal stability of these magnetizations, providing a record of plate tectonic motions over millions of years, is explained by thermal activation theory. One chapter is devoted to practical tests of domain state and paleomagnetic stability; another deals with pseudo-single-domain magnetism. The final four chapters place magnetism in the context of igneous, sedimentary, metamorphic, and extraterrestrial rocks. This book will be of great value to graduate students and researchers in geophysics and geology, particularly in paleomagnetism and rock magnetism, as well as physicists and electrical engineers interested in fine-particle magnetism and magnetic recording.
A single pronounced internal reflecting horizon has been observed on radio echo-sounding from over 30 glaciers in Spitsbergen. They are often present along the entire length of the glacier, remaining at a fairly constant depth (100–200 m) below the ice surface. Echo-strength data from radio echo-sounding have been used to obtain reflection coefficients, for these horizons, of between -15 and -25 dB. Combined with results of ice-core studies, the possible causes of this internal layer are investigated. The presence of water is found to be the most likely explanation, indicating the existence, at depth, of a layer of temperate ice.
An opportunity to analyse the mechanism for the formation of a basal ice layer was provided by Glacier de Tsanfleuron in the Swiss Alps. This glacier has some lateral subglacial cavities in which a basal ice layer forms on the bedrock floor and is subsequently incorporated at the base of the glacier. Petrographic and crystallographic analyses of the different types of ice have provided a means of investigating the successive stages of the process. These analyses show major structural changes which provide field confirmation of the deformation mechanisms studied by several authors in laboratory experiments.
A single pronounced internal reflecting horizon has been observed on radio echo-sounding from over 30 glaciers in Spitsbergen. They are often present along the entire length of the glacier, remaining at a fairly constant depth (100–200 m) below the ice surface. Echo-strength data from radio echo-sounding have been used to obtain reflection coefficients, for these horizons, of between -15 and -25 dB. Combined with results of ice-core studies, the possible causes of this internal layer are investigated. The presence of water is found to be the most likely explanation, indicating the existence, at depth, of a layer of temperate ice.
This work presents the detailed geology of Svalbard. It arises from about 50 years of research in many aspects of Svalbard geology by the author, with many colleagues and collaborators. The work is divided into four parts, the first being introductory, setting the scene and outlining the main geological features and the principal geological conventions used throughout. Part two divides Svalbard into eight regions, describing each in a separate chapter. Part three looks at historical events and environments, and part four comprises a summary of the economic aspects of Svalbard geology, plus indexes and reference lists.