Introduction
Valley glaciers on the High Arctic archipelago of Svalbard have typically experienced continuous recession and thinning since reaching their Neoglacial maximum positions towards the end of the Little Ice Age (LIA; AD ∼1900), except those that have surged during this time. This recession has largely been in response to the significant step-like increase in warming at the start of the 20th century which marked the termination of the LIA (Reference Hanssen-Bauer, Solås and SteffensenHanssen-Bauer and others, 1990), resulting in predominantly negative mass balances for many glaciers (Reference DowdeswellDowdeswell and others, 1997). The majority of small (<5 km long) land-terminating valley glaciers are now likely to be largely or entirely cold-based and frozen to their beds, although this is unlikely to have been the case when the glaciers were substantially thicker and more extensive at their LIA maxima (e.g. Reference Hodgkins, Hagen and HamranHodgkins and others, 1999; Reference Stuart, Murray, Gamble, Hayes and HodsonStuart and others, 2003; Reference HambreyHambrey and others, 2005; Reference Midgley, Tonkin, Cook and GrahamMidgley and others, 2013). Thus, many glaciers are currently in the process of undergoing, or have already undergone, a switch from a polythermal to a cold-based thermal regime (e.g. Reference Hodgkins, Hagen and HamranHodgkins and others, 1999; Reference HambreyHambrey and others, 2005). Reconstructing the timing and characteristics of these changes is important as it provides a robust link between glacier thermal regime and climate cycles, which also has implications for associated changes to flow dynamics (e.g. Reference HambreyHambrey and others, 2005), such as thermally controlled surge behaviour (e.g. Reference Fowler, Murray and NgFowler and others, 2001), and the existence (and mobility) of subglacial microbial life (e.g. Reference Tranter, Skidmore and WadhamTranter and others, 2005; Reference HodsonHodson and others, 2008). The difficulty lies in elucidating the exact nature of thermal regimes during the LIA, as this typically exists beyond observational and instrumented records (cf. Reference HambreyHambrey and others, 2005). Glacial geomorphology can be used to reconstruct the former dimensions and thickness of small valley glaciers but often provides few direct clues about past flow dynamics or thermal regime. This is because glacier forelands are typically dominated by thermo-erosion processes associated with the degradation of buried ice. These processes often negate the preservation of small-scale or poorly defined landforms that may be diagnostic of former warm-based flow (e.g. flutes) due to widespread sediment remobilization and meltwater action (Reference Etzelmüller, Hagen, Vatne, Ødegård and SollidEtzelmüller and others, 1996; Reference Lukas, Nicholson, Ross and HumlumLukas and others, 2005; Reference EvansEvans, 2009).
Evidence may, however, be preserved in basal ice sequences and the internal structural attributes of a glacier, as the formation and evolution of both is often linked to dynamic, warm-based ice flow (e.g. Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference Alley, Cuffey, Evenson, Strasser, Lawson and LarsonAlley and others, 1997; Reference HambreyHambrey and others, 2005; Reference Cook, Swift, Graham and MidgleyCook and others, 2011). Reference Sharp, Jouzel, Hubbard and LawsonSharp and others (1994) assessed the physical and isotopic characteristics of basal ice exposed at the margins of Variegated Glacier, Alaska, USA. They were able to link the formation of different facies to processes active during the dynamic ice flow of a surge. These processes included freeze-on of subglacial meltwater in close association with the debris-rich bed, apron overriding (cf. Reference EvansEvans, 1989), tectonic thickening due to folding and faulting of units, and metamorphism of glacier ice in close proximity to the bed. It has also been demonstrated that the three-dimensional (3-D) structural attributes of a glacier can be directly related to dominant strain patterns within the ice, and changing flow dynamics can be inferred from the sequential development of structures (e.g. Reference Lawson, Sharp, Hambrey, Maltman, Hubbard and HambreyLawson and others, 2000; Reference HambreyHambrey and others, 2005). At Midtre Lovénbreen, northwest Svalbard, the analysis of structural attributes and their evolution over time allowed Reference HambreyHambrey and others (2005) to determine that the formation of different structural elements was dependent on spatial and temporal variations in the flow regime of the glacier, characterized by simple shear close to flow unit boundaries (longitudinal foliation), longitudinal compression (arcuate fractures, interpreted as thrusts) and extensional flow (crevasse traces). It was concluded that the structures relating to compressive and extensional flow regimes, which were no longer actively forming, were most likely to have developed during the LIA, when the glacier was thicker, warm-based and more dynamic. Investigations of basal ice sequences and the internal structure of a glacier can therefore be extremely valuable for determining past flow dynamics when direct observations or diagnostic landform-sediment assemblages are absent.
The main aims of this paper are: (1) to describe and map the basal sequence and glaciological structures exposed at Tellbreen, a small cold-based valley glacier; (2) to reconstruct the formation of sediment/ice facies and the development of structural attributes; and (3) to assess evidence for changes to the flow dynamics and thermal regime of the glacier, and to consider the implications for the relationship between dynamic instabilities and climate change.
Study Site and Methods
Tellbreen is a 4 km long, 0.5 km wide land-terminating valley glacier in central Spitsbergen (78°15′N, 16°12′E), located ∼12 km east of Longyearbyen, the main settlement on Svalbard (Fig. 1). The glacier is fed by two high-elevation accumulation basins (up to 950 ma.s.l.) and terminates at ∼300 m a.s.l. in a ∼0.5 km2 area of ice-cored moraine. Local bedrock comprises sandstones, siltstones and shales of the Van Mijenfjorden and Adventdalen Groups (Reference Dallmann, Ohta, Elvevold and BlomeierDallmann and others, 2002). Ground-penetrating radar (GPR) data indicate that Tellbreen is currently almost entirely cold-based, with possibly only a small isolated area of warm ice below the thickest (∼100 m) part of the glacier (Reference Bælum and BennBælum and Benn, 2011). It was suggested by Reference Bælum and BennBælum and Benn (2011) that warm-based ice may have been more extensive at the LIA maximum extent of the glacier, but, based on the minimal evidence for valley erosion or modification, was unlikely to have been widespread or prolonged. Since the LIA, Tellbreen has undergone terminus retreat of ∼1 km (Fig. 1) and an estimated ∼60–70% loss of ice volume and >50% reduction in glacierized area. A long-term mass balance of −0.6 ± 0.2 m w.e. a−1 has been calculated over this period (Reference Bælum and BennBælum and Benn, 2011), similar to the average calculated for all Svalbard glaciers (−0.55 m a−1; Reference DowdeswellDowdeswell and others, 1997). Reference Bælum and BennBælum and Benn (2011) found no evidence to suggest Tellbreen had ever undergone surge behaviour.
Location map of Tellbreen in central Spitsbergen. The 2009 aerial photograph mosaic (S2009_1383500426 and 00428) of the glacier lower tongue and neighbouring Louis Careybreen (unofficial name) is published with the permission of Norsk Polarinstitutt (NPI). Red dots show location of SW, AC and NE cave entrances. Yellow dashed line shows LIA maximum position based on Reference Bælum and BennBælum and Benn (2011). White rectangle delimits area mapped in Figure 8. Inset shows location of Tellbreen (black dot) within Svalbard.
Tellbreen has an active drainage system characterized by a network of supra-, en- and subglacial conduits formed by cut-and-closure processes (incision followed by roof closure), which even within a cold glacier can route meltwater from the surface to the bed (Reference Bælum and BennBælum and Benn, 2011; Reference Naegeli, Lovell, Zemp and BennNaegeli and others, 2014). For the purposes of this study, the presence of active and abandoned conduits within the lower glacier tongue provided an accessible way to investigate ice facies and glaciological structures when they were largely free of water during spring.
Three conduits, or caves, were investigated, named the southwest (SW), active conduit (AC) and northeast (NE) caves (Figs 1 and 2). The SW cave (Fig. 2a) is an abandoned conduit located close to the indistinct transition between ice-cored moraine and debris-covered glacier; the AC cave (Fig. 2b; ‘Crack cave’ in Reference Naegeli, Lovell, Zemp and BennNaegeli and others, 2014) is the lowermost englacial section of a conduit that emerges from the glacier front at the cave entrance; and the NE cave (Fig. 2c) is an open, cavern-like area formed by meltout of the former northeast lateral channel (‘Feather cave’ in Reference Naegeli, Lovell, Zemp and BennNaegeli and others, 2014). The AC and NE caves gave access to the glacier bed, while the SW cave is located ∼5 m above the bed. The distributions of sediment and ice facies and glaciological structures were logged as two-dimensional (2-D) sections across the three caves, named the SW1, SW2, AC1 and NE1 sections (Figs 2 and 3). In places, the exposures of glacier ice are coated by thin layers of refrozen meltwater, which are shown where relevant but are not described in any detail in this paper.
Ice caves within Tellbreen lower glacier tongue. (a) SW cave with SW1 and SW2 sections arrowed. (b) AC cave with AC1 section arrowed. (c) Down-glacier entrance of NE cave. (d) Matrix-supported diamict and sand and gravel overlain by dispersed facies ice within NE1 section.
Two-dimensional section logs: (a) SW1 section; (b) SW2 section; (c) AC1 section; (d) NE1 section. A5 and A6 in (c) refer to cave measurement stations within Reference Naegeli, Lovell, Zemp and BennNaegeli and others (2014), where it is named ‘Crack cave’.
Sediment and ice facies were identified and classified according to their physical characteristics following Reference Evans, Benn, Evans and BennEvans and Benn (2004) and Reference Hubbard, Cook and CoulsonHubbard and others (2009), respectively, including overall facies thickness, structure, debris content and bubble content (Table 1). A combination of both the sedimentological (Reference Evans, Benn, Evans and BennEvans and Benn, 2004) and basal ice facies (Reference Hubbard, Cook and CoulsonHubbard and others, 2009) classification schemes was necessary in order to fully capture the variation in sediment and ice facies present within the caves, particularly in relation to areas comprising debris-rich ice/ice-rich debris. Decimetre-scale blocks were removed with an ice axe and sampled for debris concentration, grain-size distribution and stable isotope analysis. Debris concentrations (% by volume) were calculated by melting the sample in a beaker, allowing the debris to settle and recording the total volume and volume of the debris, following Reference KnightKnight (1997) and Reference WallerWaller (1997). Similar to Reference WallerWaller (1997), a value of <1% was recorded for samples with a debris concentration too low to measure. Grain-size distributions of the debris were determined by dry sieving (−4.0 to 1.0φ) and laser granulometry (<1.0φ) and plotted using GRADISTAT (Reference Blott and PyeBlott and Pye, 2001). Distributions derived from laser sizing (percentage volume) were calculated as a percentage of the entire sample weight, and for graphical purposes were plotted together with the distributions derived from sieving (percentage weight) after standardization; this has been done to show the range of grain sizes present within different facies, and it is acknowledged that caution is essential when interpreting such datasets derived from different methods (cf. Reference Hoey, Evans and BennHoey, 2004). Clast-shape and -fabric data from the matrix-supported diamict were collected following Reference LukasLukas and others (2013) and Reference Benn, Evans and BennBenn (2004), respectively. Samples for stable isotope analysis were either taken from the entirely melted ice blocks following filtering or extracted directly from the sections using an ice screw. All samples were stored in 30 mL HDPE narrow-neck, screw-top bottles. Analysis was undertaken at the University of Birmingham, UK, on a GV Instruments Isoprime continuous-flow mass spectrometer.
Key characteristics of sediment and ice facies identified at Tellbreen
Planar glaciological structures were first identified within the ice caves and measured as strike and dip using a compass/clinometer. These field data were then integrated with the structures mapped on the glacier surface from aerial photographs and classified and coded from S0 upwards based on order of formation, in accordance with structural geology conventions (Reference HambreyHambrey and others, 2005; Reference Roberson and HubbardRoberson and Hubbard, 2010). The dip direction and dip (linear structure) of sheared englacial debris laminae, or ‘mineral stretching lineation’ (Reference FlemingFleming and others, 2013), were also recorded. All field-measured fabric data were plotted as equal-area stereographic projections using Stereo32 (Reference Röller and TrepmannRöller and Trepmann, 2008), which was also used to calculate fabric statistics (S1, S2 and S3 eigenvalues, not to be confused with structural glaciology notations) following Reference Benn, Evans and BennBenn (2004).
Results
Facies distribution
The geographical distribution of the main sediment and ice facies within the SW, AC and NE caves (Fig. 2) is briefly described first to provide some initial context, followed by a detailed description of the physical and isotopic characteristics of the facies (Table 1).
SW cave
Englacial facies ice is the dominant ice type throughout the SW cave, which is cross-cut by thin debris bands logged as S2 and S3 structures in the SW1 and SW2 sections (see Structural glaciology subsection below; Figs 2a and 3a and b). Some of the thicker S3 structures contain sorted sediments (Fig. 3a and b).
AC cave
Englacial facies ice is also the dominant ice type within the AC cave, as shown within the area logged as the AC1 section (Figs 2b and 3c). S2 and S3 structures cross-cut the englacial facies ice, and some of the thicker S3 structures contain sorted sediments (Fig. 3c). Matrix-supported diamict is found at the base of the cave immediately up-glacier of the AC1 section, and small areas of dispersed facies ice were also observed throughout the cave.
NE cave
The base of the NE cave consists of matrix-supported diamict (Figs 2d and 3d). This is overlain by dispersed facies ice, which extends to the cave roof in the NE1 section. There is a large area of sorted sediment close to the cave entrance, and several thinner bands of sorted sediment are found throughout the NE1 section (Fig. 3d).
Sediment facies
Matrix-supported diamict
This facies forms the lowermost unit at the NE1 section (Figs 2d and 3d), where it ranges in thickness from 2 to 4 m, and consists of frozen, poorly sorted diamict with interstitial ice and small, largely bubble-free clean ice lenses (Table 1; Fig. 4a and b). The diamict is matrix-supported and contains predominantly subangular clasts (Fig. 5a), ranging up to boulder size (0.5 m in diameter). The grain-size distribution of the facies is polymodal and displays distinct peaks within silt and sand (Fig. 5b). Debris concentrations vary with height across the thickness of the facies, grading over tens of centimetres from measured values of 52% close to the contact with overlying ice, in an area where numerous clean ice lenses are visible, to 80% at ∼1 m depth (Table 1; Fig. 4b). Intermediate debris concentrations of 65% and 70% were measured from the area between these two (Fig. 4b). The upper ice-rich area is observed to vary in thickness from ∼0.1 to 1 m and displays crude stratification highlighted by the presence of ice lenses. This stratification becomes less distinct with depth, and below ∼1 m the facies appears structureless. The debris in this lowermost area contains only interstitial ice and is effectively frozen diamict. The ice-rich parts of this facies could also be described as solid stratified facies ice according to the Reference Hubbard, Cook and CoulsonHubbard and others (2009) classification. Occasional thin (∼5 cm), horizontally aligned layers of sorted sands and gravel are found both within the diamict and in places at its interface with overlying dispersed facies ice (Fig. 4b). Two clast fabric samples recorded from either end of section NE1 (Fig. 3d) show mean lineation azimuths of 129° and 109°, and display moderate to weak clustering (S1 values of 0.65 and 0.56; Table 2; Fig. 6a and b). Stable isotope analysis of the clean ice lenses and interstitial ice within this facies returned mean values of −12.11‰ (δ18O), which is statistically distinct from the englacial facies mean values at the 95% confidence level (i.e. ±2σ), and −92.66‰ (δD), which shows slightly higher values than englacial ice but is not statistically distinguishable at the 95% confidence level (Table 3; Fig. 7). These data do not show significant linear relationships (e.g. freezing slopes) when plotted co-isotopically. Restricted exposures of similar matrix-supported diamict also occur in AC cave, although in most places the cave floor is obscured by the coarse, poorly sorted bed load of the meltwater stream.
Fabric statistics of S2 lineation and clast fabric data. Numbers in parentheses refer to numbered sampling sites in Figure 2d
Examples of sediment and ice facies. (a) Matrix-supported diamict at base of NE cave. Lens cap circled for scale. (b) Matrix-supported diamict (Dm) overlain by dispersed facies (D) in NE cave. Note clean ice lenses within matrix-supported diamict and thin layer of sands and gravels at contact with dispersed facies. Percentages in white are measured debris concentrations (% by volume). Lens cap circled for scale. (c) Dispersed facies overlying matrix-supported diamict at base of NE cave. (d) Dispersed facies overlying matrix-supported diamict within NE cave. Note areas of both bubble-free and bubble-rich ice. (e) Filament-like bubble structures within clear dispersed facies at NE cave. (f) Filament-like bubble structures within clear dispersed facies towards entrance of AC cave. (g) Englacial ice within SW cave. Note white appearance and evidence for horizontally aligned intercalated layers of bubble-rich and bubble-poor ice.
Sedimentary characteristics of matrix-supported diamict and dispersed ice facies. (a) Clast shape data from matrix-supported diamict at NE1 section plotted on histograms (roundness) and ternary diagrams (shape) generated using the TRI-PLOT spreadsheet (Reference Graham and MidgleyGraham and Midgley, 2000). Each sample is of 50 sandstone clasts. Numbers in parentheses refer to sample numbers in Figure 3d. (b) Grain-size distributions of matrix-supported diamict (solid line; n = 5) and debris laminae within dispersed facies (dashed line; n = 3) from NE1 section. Troughs in the data are an artefact of graphically combining dry sieving (in shaded area) and laser granulometry methods.
Lower-hemisphere equal-area stereographic projections and rose diagrams of structural and fabric data within Tellbreen. (a, b) Clast fabric samples of matrix-supported diamict at NE1. Numbers in parentheses refer to numbered sampling sites in Figures 2d. (c)S0/S1 in SW1, SW2 and AC1. (d) S2 in SW1 and SW2. (e) S2 in AC1. (f) S2 in dispersed facies at NE1. (g) S2 lineation in SW1 and SW2. (h) S2 lineation in AC1. (i) S2 lineation in NE1. (j) S3 in SW1. (k) S3 in AC1. Structural planes are plotted as great circles while poles to bedding are plotted as solid squares. Black arrow shows centre-line ice flow direction at the terminus.
Box plots of stable isotope analysis of Tellbreen ice facies and glaciological structures showing maximum, upper quartile, median, mean (black diamond), lower quartile and minimum values. (a) δ18O composition. (b) dD composition.
Summary of stable isotope data
Sorted sediments
At the left-hand end of the NE1 section, the matrix-supported diamict is in contact with a ∼1–2 m thick band of sorted sediments that extends laterally into glacier ice. The sediments consist of layers of sorted fine sand, interbedded sands and gravels and clast-supported massive gravel (Fig. 3d). The sorted fine sand layers range from 10 to 50 cm in thickness and display laminations, which in places have been subjected to small-scale folding. These layers are overlain by interbedded coarse sand and fine gravel which dip down-channel at ∼30°. Clast-supported and imbricated coarse gravel layers are also present, the largest of which is ∼50 cm thick and extends for ∼6 m laterally. Similar sediment bands were also observed elsewhere in the walls of the NE cave, forming laterally extensive, gently sloping layers that cut across ice foliation and other structures. Examples also occur in the SW and AC caves (Fig. 3).
Ice facies
Dispersed facies
The dispersed facies overlies the matrix-supported diamict at the NE1 section (Fig. 3d) and within the AC cave, and is characterized by debris-poor ice with variable bubble content and character (Table 1; Fig. 4c and d). The thickness of the facies ranges from <0.1 m within the AC cave to ≥2 m at NE1 (Fig. 3d). The debris within the ice takes the form of suspended grains or clots of predominantly silt-sized fine sediment, occasionally forming thin laminae (<1 cm; Fig. 5b): Debris concentrations of <1% were measured at the NE1 section (Table 1). Suspensions of bright orange material were observed in places, and many of the laminae had a strong linear component, characterized by the strong alignment of grains or clots of fine sediment (Fig. 8c). Bubble content and character varies from clear areas, with no or very few bubbles (Fig. 4b), to areas containing dense, white clouds of bubbles and intercalated bubble-rich and bubble-poor layers (Fig. 4c and d). The bubble-poor areas are typically located close to the contact with the underlying frozen diamict (Figs 3d and 4b) and often contain filament-like structures of very fine bubbles (Fig. 4e and f). In places, these structures grade into thicker, ribbon-like features, several individual bubbles (∼1–5 mm diameter) and dense bubble clouds (Fig. 4c). In areas where they coincide, these bubble structures cut across intercalated layers of bubble-rich and bubble-poor ice within the dispersed facies. Stable isotope analysis returned mean compositions of −13.92‰ (δ18O) and −102.13‰ (δD), and the facies shows slightly heavier values of δ18O than englacial ice mean values but these are not statistically distinct at the 95% confidence level (Table 3; Fig. 7). These data do not display significant linear relationships (e.g. freezing slopes) when plotted co-isotopically.
Structural interpretation of Tellbreen lower tongue from 2009 imagery. Aerial photograph mosaic (S2009_13835 00426 and 00428) published with the permission of Norsk Polarinstitutt (NPI).
Englacial facies
The englacial facies has ubiquitous stratification at centimetre to decimetre scales, in the form of intercalated layers of bubble-rich and bubble-poor ice (Fig. 4e). This layering is visible on the glacier surface as slight colour changes. Apart from the thin debris bands which cross-cut this facies within the SW and AC caves (Figs 2a and b and3a–c), the ice contains only occasional suspended grains and small clots of fine sediment. The englacial facies returned mean stable isotope values of −14.89‰ (δ18O) and −101.26‰ (δD) (Table 3).
Structural glaciology
The planar glaciological structures described in this subsection (Table 4) occur in englacial ice, typically up to several metres above the bed. These structures are observed both within the ice caves (Fig. 3) and on the glacier surface (Fig. 8). The overall distribution of the structures (S0–S4; Table 4) on the glacier surface is as follows: Primary stratification (S0) is only identified in the upper part of the lower tongue of Tellbreen, towards the western margin of the glacier (Fig. 8). Longitudinal foliation (S1) is ubiquitous across the entire glacier surface and is generally aligned with the dominant ice flow direction. The only exception to this is close to the western margin, where there is a small area of S1 structures which are oriented obliquely to the main set. Arcuate fracture traces (S2) are generally oriented perpendicular to the dominant ice flow direction and are primarily found on the lowermost part of the glacier tongue, where they are distributed across the full width of the glacier. Fracture traces (S3) are also aligned perpendicular to flow, and are distributed throughout the lower tongue. Open fractures (S4) are only observed towards the eastern margin of the lower glacier tongue (Fig. 8) and high up in the accumulation basins.
Summary of principal glaciological structures within Tellbreen
Characteristics of each of these structures and their relationships to each other will now be described in more detail based on both the surface mapping (Fig. 8) and investigations within the ice caves (Fig. 3).
S0 – primary stratification
This structure is defined by the alternations of bubble-rich and bubble-poor ice within englacial ice (Figs 4e and 9a and d). This layering appears as slight colour changes on the glacier surface, and in places can be traced as irregular linear features, although these are not identifiable in the lowermost part of the tongue (Fig. 8). Stratification is crosscut by S2 and S3 structures (Fig. 9a and d).
Details of glaciological structures in Tellbreen. (a) S0/S1, S2 and S3 structures exposed in SW cave. Note how S3 structures cut across S2 structures. Lens cap circled for scale. (b) S1 (solid lines), S2 (dashed lines) and S3 (dotted lines) structures exposed on the surface of Tellbreen. Note longitudinal supraglacial ridges. (c) S3 structure exposed in side of a meltwater conduit. (d) S0/S1 and S2 structures within SW cave. Note linear smearing of sediment grains and clots within S2 structures. (e) S2 and S3 structures within SW cave. Note linear smearing of sediment grains and clots within S2 structures, which are cut across by S3 structures. (f) Poorly sorted sandy gravel within S3 structure in SW cave. (g) S3 structures cutting across S2 structures within SW cave. (h) Sharp-crested transverse supraglacial ridge composed of stratified sand and gravel on lower tongue. Note abrupt right-angled change in orientation.
S1 – longitudinal foliation
Linear, flow-parallel stripes on the glacier surface are identified as longitudinal foliation based on their similarity to features mapped as such in other studies (e.g. Reference Hambrey and DowdeswellHambrey and Dowdeswell, 1997; Reference HambreyHambrey and others, 2005; Reference Roberson and HubbardRoberson and Hubbard, 2010). Longitudinal foliation is closely associated with longitudinal supraglacial ridges mantled by thin (<5 cm) drapes of angular material (Figs 8 and 9b).
S2 – arcuate fracture traces
The glacier surface displays a number of linear stripes orientated perpendicular and sub-perpendicular to S1 structures and ice flow, identified as trace evidence of brittle deformation (Fig. 8). These are subdivided into arcuate fracture traces (S2) and fracture traces (S3). Arcuate fracture traces appear on the glacier surface as gently curving linear stripes, often traceable for tens to hundreds of metres (Figs 8 and 9b), and as one set of the thin (<2 cm) debris bands exposed within the ice caves (Figs 3 and 9). The correlation between the surface and englacial structures is based on their similar characteristics and relationships to other structural elements. Both on the glacier surface and in the ice caves, S2 structures (1) rarely cross-cut each other but do intersect on occasion (Figs 8 and 9a); (2) display similarly consistent orientations, with measurements from the SW and AC caves recording dominant orientations of ∼350–010°, and within the NE cave ∼040–060° (Fig. 6d–f); and (3) are cross-cut by S3 structures, which display more variable orientations (Figs 8 and 9a). In the ice caves, some S2 structures display centimetre-scale offsets of S0 stratification in a vertical direction. The debris within englacial S2 structures includes individual grains and small clots of fine material suspended within bubble-poor ice (Fig. 9d and e) with a strong peak in the silt size range (Fig. 5b). In almost all cases the debris displays a strong linear component, and fabric data from these show mean lineation azimuths of 316° (sections SW1 and SW2), 308° (AC1) and 131° (NE1), recording a northwest–southeast alignment (Table 2; Fig. 6g–i), and display strong clustering (S1 values of 0.93, 0.92 and 0.53; Table 2). Stable isotope analysis of ice within the S2 structures shows slightly heavier δ18O values than for englacial ice, and lighter values than for the ice lenses and interstitial ice within the diamict, but these values are not statistically distinct at the 95% confidence level (Table 3; Fig. 7).
S3 – fracture traces
The second group of fracture traces (S3) are generally shorter and display more variable orientations than S2 structures on the glacier surface (Fig. 9b), and are observed to cross-cut S0/S1, S2 and other S3 structures (Fig. 8). These surface features are correlated to the second set of debris bands within the SW and AC caves, which clearly cross-cut both S0/S1 and S2 structures (Figs 3 and 9a, e and g) and in places can be directly traced to the surface (Fig. 9c). In addition to orientations, the sediment composition and thicknesses of the S3 structures are also more variable than S2, ranging from <1 cm thick bands of suspended grains and clots of well-sorted silt-sized material (Fig. 9e and g), to ∼5–10 cm thick bands of poorly sorted sandy gravel with up to 2–3 cm diameter clasts (Fig. 9f). This sediment composition is similar to that within transverse supraglacial ridges on the lower part of the glacier (Fig. 8), which are aligned perpendicular and sub-perpendicular to ice flow direction and in the field take the form of sharp-crested ridges composed of layers of cross-bedded sand and gravels up to cobble size (Fig. 9h). There is no evidence of mineral stretching lineations or shearing within S3 structures. The stable isotope analysis of ice within S3 structures returned slightly heavier values of δ18O compared to englacial ice, and lighter values compared to ice within the diamict, but these are not statistically distinct at the 95% confidence level (Table 3; Fig. 7).
S4 – open fractures
These have a very restricted distribution on the lower tongue and were typically only found in close association with the NE cave complex (Fig. 8). Open fractures also occur in steep areas of the upper accumulation basins (cf. Reference Bælum and BennBælum and Benn, 2011).
Interpretation
Sediment facies
Frozen subglacial traction till
This facies is interpreted as a frozen subglacial traction till (cf. Reference Evans, Phillips, Hiemstra and AutonEvans and others, 2006; Reference Benn and EvansBenn and Evans, 2010) based on its textural characteristics (poorly sorted, matrix-supported, predominance of subangular clasts), moderate to weakly clustered clast fabric aligned with ice flow (Table 2; Fig. 6a and b), and evidence for winnowing at the ice/bed interface in the form of sorted sand and gravel layers. The lowermost ∼0.5–1 m of the facies at the NE cave was observed to be almost entirely frozen diamict with very little ice content (Fig. 4b), and, in conjunction with the location of the cave close to the glacier terminus, it is concluded that the NE1 section is at the glacier bed. The ice lenses in the upper part of the diamict are interpreted as segregation lenses formed from the till pore water, and relatively low debris contents measured in these areas (∼50%) indicate that the till was highly saturated and probably dilated close to the ice/bed interface (Reference Evans, Phillips, Hiemstra and AutonEvans and others, 2006). The vertical gradation in debris content, which increases in a downwards direction, likely reflects a decrease in till dilatancy. It seems likely that freezing was initiated from the glacier ice above, leading to the progressive downwards migration of the freezing front (cf. Reference Christoffersen, Tulaczyk, Carsey and BeharChristoffersen and others, 2006) into the underlying diamict, forming ice lenses in the upper, more dilated parts of the facies and less-ice-rich frozen diamict towards the base. The results from stable isotope analysis demonstrate that the ice lenses are enhanced in heavy 18O isotopes relative to englacial facies ice (Table 3; Fig. 7), which is consistent with ice formation through refreezing of water in close association with a debris-rich bed (cf. Reference LawsonLawson, 1979; Reference Hubbard and SharpHubbard and Sharp, 1993; Reference Iverson and SouchezIverson and Souchez, 1996).
Glacifluvial sediments
The sorted sediments are interpreted as glacifluvial deposits. The cross-cutting relationship between the sorted sediment bands and the surrounding glacier ice indicates that they are superimposed upon, but not directly related to, the exposed sequence of ice facies and frozen diamict. One interpretation for the formation of the bands is as late-stage sediment infills of cut-and-closure conduits that progressively incised downwards and laterally through the glacier ice and into the bed, described in more detail by Reference Naegeli, Lovell, Zemp and BennNaegeli and others (2014). A second possible origin for the sediments is based on interpretations by Reference EvansEvans (1989) of similar discrete pockets of glacifluvial sediments within marginal ice at several glaciers on Ellesmere Island, Canadian High Arctic. At these sites, the presence of glacifluvial sediments was suggested to relate to the lateral incision of ice-marginal meltwater channels into glacier ice and the subsequent accumulation of sorted sediments within ponded areas. The lateral incision of the channels destabilizes the ice margin, resulting in block collapse and, if the glacier is advancing, apron entrainment (Reference EvansEvans, 1989). This interpretation is consistent with observations from the NE cave, where the northeast ice-marginal channel has incised laterally into the ice margin, leading to localized destabilization and progressive collapse of the cave roof.
In a few cases, thin bands of sand and gravel appear to have formed by other processes. These include the thin, conformable layers of sorted sediment at the contact between the frozen diamict and overlying dispersed facies ice at the NE1 section (Figs 3d and 4c), which are interpreted as evidence for winnowing at the ice/bed interface, and poorly sorted sand and gravel within some S3 glaciological structures, described below in the Structural glaciology subsection.
Ice facies
Dispersed facies: metamorphosed basal ice
Dispersed facies ice has been described from the basal ice sequences of several glaciers (Reference LawsonLawson, 1979; Reference Larsen, Kronborg, Yde and KnudsenLarsen and others, 2010; Reference Cook, Swift, Graham and MidgleyCook and others, 2011), and has also been referred to as ‘clotted facies’ ice (Reference KnightKnight, 1987; Reference Sugden, Clapperton, Gemmell and KnightSugden and others, 1987; Reference Knight and KnightKnight and Knight, 1994) and ‘clear facies’ ice (Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference Hubbard and SharpHubbard and Sharp, 1995; Reference Hubbard, Tison, Janssens and SpiroHubbard and others, 2000). Reference Cook, Swift, Graham and MidgleyCook and others (2011) highlighted that these facies are all descriptively similar, but are interpreted to have formed by a range of processes, with current interpretations favouring either a primarily sedimentary (e.g. Reference LawsonLawson, 1979; Reference KnightKnight, 1987; Reference Sugden, Clapperton, Gemmell and KnightSugden and others, 1987) or tectonic (e.g. Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference Hubbard, Tison, Janssens and SpiroHubbard and others, 2000; Reference Waller, Hart and KnightWaller and others, 2000; Reference Cook, Swift, Graham and MidgleyCook and others, 2011) origin. At Tellbreen, the evidence is most consistent with a tectonic interpretation for the dispersed facies, characterized by strain-induced metamorphism close to the bed (e.g. Reference KnightKnight, 1997; Reference Hubbard, Tison, Janssens and SpiroHubbard and others, 2000; Reference Waller, Hart and KnightWaller and others, 2000; Reference Cook, Swift, Graham and MidgleyCook and others, 2011). The bubble-poor areas are consistent with widespread melting and refreezing at grain boundaries and associated gas expulsion, caused by enhanced ice deformation and strain heating consistent with warm-based conditions (Reference Kamb and LaChapelleKamb and LaChapelle, 1964; Reference Blatter and HutterBlatter and Hutter, 1991; Reference Hubbard and SharpHubbard and Sharp, 1995; Reference Hubbard, Tison, Janssens and SpiroHubbard and others, 2000). This process provides an explanation for the bubble structures observed within the dispersed facies, which are characterized by fine bubble strands tracing the outline of ice crystals (Fig. 4e and f), consistent with a routing of gas around grain boundaries. Bubble-rich areas, in places displaying stratification not dissimilar to that within englacial ice (Fig. 4c), may represent preserved remnants of the latter facies where gas expulsion was incomplete (cf. Reference Hubbard and SharpHubbard and Sharp, 1995). The dispersed ice immediately overlying the frozen diamict, and therefore at the inferred ice/bed interface, typically has the lowest bubble content (Fig. 4b), indicating that this is where metamorphism was most effective. The fine-grained debris within the dispersed facies is inferred to represent the migration of muddy water within the vein network between ice crystals in the basal zone (Reference LliboutryLliboutry, 1993; Reference Knight and KnightKnight and Knight, 1994). The debris is likely to have been sourced from the underlying saturated till, and its elevation into the dispersed facies suggests that it was probably highly pressurized in order to be forced upwards into the vein network, indicating temperate conditions in the basal zone characterized by the availability and mobility of subglacial meltwater in close association with the debris-rich bed. The occasional suspensions of bright orange material within the clear ice are characteristically similar to precipitates of iron oxyhydroxide described at other sites on Svalbard (Reference HodsonHodson and others, 2008). The strong linear component to the debris laminae was also observed by Reference FlemingFleming and others (2013) and these were interpreted as stretching lineations, formed by the realignment of debris about an axis under high-strain conditions. Stretching lineations on the planar surface of laminae at the NE cave display strong evidence of shear in a direction sub-parallel to ice flow (Table 2; Fig. 6k), in agreement with clast fabric data from the underlying diamict (Table 2; Fig. 6a and b).
Stable isotope results for the dispersed facies show that values for δ18O are slightly heavier than those for englacial ice, but are not statistically distinct at the 95% confidence level (Table 3). These data do not form a freezing slope when plotted co-isotopically, indicating that the facies has not formed by the refreezing of parent water from a single source with specific isotopic composition (Reference Souchez, Lorrain, Tison and JouzelSouchez and others, 1988). This may instead reflect multiple melting and partial refreezing events at grain boundaries as part of the metamorphism process (Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994). Reference Souchez, Lorrain, Tison and JouzelSouchez and others (1988) suggested that this may produce high ranges in δ18O and δD values (e.g. 3.85‰ and 33.91‰, respectively, measured for the dispersed facies (Table 3)). The explanation for this is as follows: While no fractionation occurs upon initial melting of existing ice, fractionation occurs upon the refreezing of the water thereby produced, creating new ice that is initially heavy and subsequently progressively lighter than the composition of the initial liquid. Thus, a high range of sample compositions in δ18O and δD would be consistent with differential degrees of refreezing occurring at the sample scale within the dispersed facies. If a small amount of last-refrozen water is also lost from the facies then this can explain its slightly heavier isotopic composition relative to that of its proposed parent (englacial facies) ice.
Englacial facies: meteoric ice
This facies is interpreted as meteoric ice formed by the firnification of snow in the glacier’s accumulation area, with the alternating bubble-poor and bubble-rich layers reflecting seasonal melting and refreezing (cf. Reference HambreyHambrey, 1975; Reference Hambrey and DowdeswellHambrey and Dowdeswell, 1997). The 18O composition of englacial ice is similar to values for recently formed meteoric ice within the Lomonosovfonna ice core, located ∼60 km to the northeast (Reference DivineDivine and others, 2008). Aside from the thin debris bands (Fig. 3), the little scattered debris that exists within the ice probably originated as wind-blown dust (e.g. Reference LawsonLawson, 1979; Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference Hubbard and SharpHubbard and Sharp, 1995; Reference Larsen, Kronborg, Yde and KnudsenLarsen and others, 2010).
Structural glaciology
Primary stratification (S0) and longitudinal foliation (S1)
Although S0 and S1 structures have been mapped as separate features, in accordance with other structural glaciology studies (e.g. Reference Hambrey and DowdeswellHambrey and Dowdeswell, 1997; Reference Hambrey and GlasserHambrey and Glasser, 2003; Reference HambreyHambrey and others, 2005; Reference Roberson and HubbardRoberson and Hubbard, 2010), we suggest that these probably represent two end members of the same structure rather than distinct features. Primary stratification (S0) describes the layers of bubble-rich and bubble-poor ice formed through firnification processes in the accumulation area (cf. Reference HambreyHambrey, 1975; Reference Hambrey and DowdeswellHambrey and Dowdeswell, 1997) and is inferred to be the primary structure in the glacier (Table 4). When it initially forms, this layering is likely to be broadly horizontal and display only slight undulations related to the formation of seasonal snow layers, but will become progressively deformed and folded as the result of flow from the accumulation area. Folds increase in tightness as the ice converges in the narrow confines of the lower glacier tongue and is subjected to lateral compression. Eventually the original layering (S0) is folded to such an extent that the limbs become isoclinal and parallel to glacier flow. Where the hinge lines of the isoclinal folds intersect the glacier surface, they form linear stripes, mapped as longitudinal foliation or S1 structures (e.g. Reference Hambrey and DowdeswellHambrey and Dowdeswell, 1997; Reference Hambrey and GlasserHambrey and Glasser, 2003; Reference Roberson and HubbardRoberson and Hubbard, 2010). As with any continuum, some structural elements must exist between the two end members; at Tellbreen, this is represented by the bubble stratification within the englacial ice facies (Fig. 4e), which was measured to dip gently in various directions within the SW and AC caves (Fig. 6c). As this has clearly been subjected to some degree of folding during transport from the accumulation area to the glacier front, this layering was recorded as S1 structures, although in reality it is perhaps unnecessary to distinguish between the two as they represent the same structure.
Longitudinal supraglacial ridges, which are closely associated with longitudinal foliation, are interpreted as debris septa (cf. Reference Eyles and RogersonEyles and Rogerson, 1978) that have formed through the same process. The debris was probably entrained as rockfall layers within stratification (Reference Eyles and RogersonEyles and Rogerson, 1978; Reference Hambrey and GlasserHambrey and Glasser, 2003; Reference Roberson and HubbardRoberson and Hubbard, 2010), which has then been tightly folded and melted out at the glacier surface as flow-parallel ridges; this interpretation is supported by the predominantly angular debris within the debris septa.
Arcuate shear planes (S2)
Similar to Reference RobersonRoberson (2008) and Reference Roberson and HubbardRoberson and Hubbard (2010), S2 arcuate fracture traces are interpreted as shear planes based on the evidence that: (1) they cut across S0/S1 layering at low angles and in places displace it vertically, and (2) the fine material within S2 has a strong linear component (Fig. 9d), or stretching lineation (Reference FlemingFleming and others, 2013), indicative of shear in the direction of ice flow (Table 2; Fig. 9i–k). It has been suggested that the development of these types of structures may be facilitated by the presence of pre-existing planar structural weakness (e.g. Reference Hambrey and MüllerHambrey and Müller, 1978; Reference Evans and ReaEvans and Rea, 1999; Reference Rea and EvansRea and Evans, 2011), such as healed crevasses (e.g. S3 fracture traces); the similar character and sedimentological compositions of S2 and S3 structures indicate that this may be the case at Tellbreen. The heavier δ18O values than for englacial ice may reflect ice formed from water from a variety of sources and therefore with varying initial isotopic compositions (i.e. from an open hydrological system) which freezes in situ in a closed system (Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others, 2001). The δ18O values are also similar to the composition of dispersed ice, which is interpreted to result from strain-induced metamorphism of englacial ice and associated processes of recrystallization, partial melting and expulsion of gases (Reference Kamb and LaChapelleKamb and LaChapelle, 1964; Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference Hubbard and SharpHubbard and Sharp, 1995). It is possible that metamorphism of this type also occurred within S2 structures, which contain largely bubble-free ice and display evidence of shear. However, caution is advised because the width of the extracted samples (∼2 cm) is similar to that of the structures sampled, making it possible that ice from outside the structures was extracted in some cases, and so this inferred process may not apply to the entire dataset.
Crevasse traces (S3)
Fracture traces are interpreted as healed fractures or crevasses which opened in response to extensional flow within the glacier (cf. Reference HambreyHambrey, 1976; Reference HambreyHambrey and others, 2005). S3 fracture traces cross-cut S2 shear planes both on the surface and in section, indicating that their formation postdates that of S2 structures. Fracture traces display variable orientations with a predominant alignment perpendicular and sub-perpendicular to ice flow (Figs 6j and k and 8), which is consistent with extensional transverse crevassing, and dip at angles varying from 008° to 055°. The sediment content of these structures can be explained as injections of meltwater of varying turbidities into fractures and crevasses under conditions of extensional flow and high basal water pressures, as has been observed at Matanuska Glacier, Alaska (Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others, 2001), and described from Skeiðarárjökull, Iceland (Reference Bennett, Huddart and WallerBennett and others, 2000). Similar to Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others (2001), it appears that many of these crevasses could be better described as narrow cracks or fractures, which may not open much wider than millimetres, into which thin films or sheets of pressurized turbid water are injected. The preservation of the debris within the structures is indicative of in situ freezing, perhaps associated with conductive cooling from surrounding colder ice and/or fracture closure (cf. Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others, 2001). The relatively low dip angles of these structures are inconsistent with the vertical crevasses/fractures traditionally associated with extensional flow (cf. Reference Rea and EvansRea and Evans, 2011), which indicates they may have been reoriented since formation (Reference Evans and ReaEvans and Rea, 1999; Reference Rea and EvansRea and Evans, 2011). The slightly heavier values of δ18O compared to englacial ice and lighter values compared to the frozen diamict (Table 3; Fig. 7) are similar to the findings of Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others (2001) for debris bands at Matanuska Glacier. This may indicate formation of ice from initial parent waters from an open hydrological system by in situ closed system freezing (Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others, 2001), although caution is advisable as this is based on limited sampling (n = 2). In addition, as with the S2 structures, the similar widths of the extracted samples and the structures mean it is also possible that some ice from outside the S3 structures was sampled. The thicker bands of sand and gravel in the SW cave are less common but are interpreted to be formed through the same process. No features similar to the coarser S3 bands were described by Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others (2001), but they share sedimentary characteristics (poorly sorted sandy gravel) with features interpreted as upwardly injected hydrofracture structures at the bed of Kuannersuit Glacier, a surge-type glacier in West Greenland (Reference Roberts, Yde, Knudsen, Long and LloydRoberts and others, 2009). The transverse supraglacial ridges composed of sand and gravel are suggested to result from the melting-out of infilled fractures, although no direct link could be made between the two apart from similar sediment compositions and ridge orientations. The sharp-crested transverse ridge displaying a prominent right-angle change of direction (Fig. 9h) is consistent with pressurized basal meltwater exploiting intersecting crevasses at a higher level in the glacier, as suggested by Reference Evans and ReaEvans and Rea (1999) for the formation of concertina or zigzag eskers associated with surge-type glaciers.
Open crevasses (S4)
Open fractures are extensional crevasses and on the lower glacier tongue are closely associated with the NE cave complex, where the ice margin has been gradually undercut by the northeastern lateral conduit (Fig. 8). Crevasses observed in the steepest parts of the upper basins represent bergschrund-type fracturing (cf. Reference Bælum and BennBælum and Benn, 2011).
Discussion
Frozen till/basal ice formation
The formation of the basal sequence at Tellbreen, comprising frozen matrix-supported diamict overlain by dispersed facies ice, can be directly linked to the flow of warm-based ice. The frozen diamict is inferred to have originated as a saturated subglacial traction till, demonstrated by the high ice content in the form of both interstitial ice and small ice lenses. The ice content reflects the freezing of subglacial meltwater stored within the pore spaces of the basal till (Reference Christoffersen, Tulaczyk, Carsey and BeharChristoffersen and others, 2006) and the development of segregation ice lenses (Reference Waller, Hart and KnightWaller and others, 2000; Reference Christoffersen and TulaczykChristoffersen and Tulaczyk, 2003). The availability of water is also demonstrated by the thin layers of sand and gravel (glacifluvial sediments) at the boundary between the diamict and dispersed facies ice, reflecting the movement of water at the ice/bed interface. These observations are consistent with the bed being at the pressure-melting point, as basal melting under warm-based conditions would allow meltwater to saturate the underlying permeable bed. Freezing of the saturated till is likely to have occurred as a cold wave or ‘freezing front’ propagated downwards through the thinning lower tongue (e.g. Reference BjørnssonBjørnsson and others, 1996). The vertical gradation in ice/debris content within the facies (Fig. 4b) may be the result of progressive freeze-on as the freezing front extends downwards into the till, forming segregation ice lenses close to the ice/bed interface (e.g. Reference Christoffersen and TulaczykChristoffersen and Tulaczyk, 2003). The gradation could also be the product of decreasing efficiency of meltwater percolation with depth, perhaps as distance from the source increases and as the sediment itself changes (e.g. in degree of consolidation or pore development and connectivity (cf. Reference Christoffersen, Tulaczyk, Carsey and BeharChristoffersen and others, 2006)). This would produce a more saturated basal till layer (i.e. higher pore-water content) closer to the ice/bed interface, creating a frozen till with higher ice content towards the contact with overlying ice (Fig. 4b).
The stable isotope data from ice lenses within the diamict show that it has been enhanced in heavy 18O isotopes relative to englacial facies ice (Table 3; Fig. 7), which is interpreted as evidence that it formed by the refreezing of highly saturated debris (cf. Reference LawsonLawson, 1979; Reference Hubbard and SharpHubbard and Sharp, 1993; Reference Iverson and SouchezIverson and Souchez, 1996). It has been demonstrated that ice formed from refreezing of a localized parent water source, or closed-system freezing (Reference KnightKnight, 1997; Reference Cook, Robinson, Fairchild, Knight, Waller and BoomerCook and others, 2010), should produce a freezing slope with a lower gradient than englacial ice when plotted on a co-isotopic graph (Reference Souchez and JouzelSouchez and Jouzel, 1984); however, this is not the case for the ice lenses within the frozen diamict at Tellbreen. An explanation for this is that the parent water was not from a single, localized source, but reflects a range of different sources with variable original isotopic compositions (Reference Hubbard and SharpHubbard and Sharp, 1993, Reference Hubbard and Sharp1995; Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference KnightKnight, 1997). This situation is consistent with the production of water through net basal melting across large parts of the bed that are at the pressure-melting point, saturating subglacial debris with water from various sources. The frozen diamict samples plot at the heavier end of the isotopic spectrum (Fig. 7), which indicates that the facies, or at least the parts of the facies sampled, formed during the early stages of the freezing process (Reference Sharp, Jouzel, Hubbard and LawsonSharp and others, 1994; Reference KnightKnight, 1997).
One of the main characteristics of dispersed facies ice is its largely bubble-poor nature, with some areas entirely devoid of bubbles. These areas are typically adjacent to the contact with the frozen till, which underlies dispersed facies ice in most cases (Fig. 4b). This is interpreted as the result of strain-induced metamorphism of englacial ice close to the bed (Reference Hubbard, Tison, Janssens and SpiroHubbard and others, 2000) due to intense and variable shear, resulting in melting and refreezing at grain boundaries and the expulsion of gas (Reference LliboutryLliboutry, 1993) away from the shear zone. Evidence for the expulsion of gas takes the form of thin filaments of bubbles routed along crystal boundaries which extend from the area of bubble-free dispersed ice (Fig. 4d). Intense and variable shear within this zone would promote strain heating of the ice, consistent with the presence of temperate ice at the bed (Reference Blatter and HutterBlatter and Hutter, 1991), and therefore warm-based conditions. The second main characteristic of dispersed facies ice is the presence of individual grains or small clots of typically silt-sized material; this is inferred to be thin films of muddy water sourced from adjacent saturated subglacial debris which have been squeezed along the intercrystalline vein network (Reference Knight and KnightKnight and Knight, 1994) in a manner similar to the expelled gases. The melting and refreezing events and movement of muddy water and gases are envisaged to effectively occur in a wet (temperate) ‘mushy’ zone (cf. Reference FowlerFowler, 1984) at the base of the englacial ice, where gas and pressurized muddy water are allowed to migrate through the vein network. This zone is likely to have partly developed due to strain heating close to the ice/bed interface, increasing the temperature of the ice to the pressure-melting point and providing an additional internal source of meltwater (Reference Blatter and HutterBlatter and Hutter, 1991).
The preserved basal sequence provides compelling support for Tellbreen having undergone a thermal switch from a polythermal regime, when large parts of the bed were at the pressure-melting point, to a cold-based thermal regime, characterized by extremely low flow velocities (personal communication from C. Bøggild, 2012) and large parts (if not all) of the glacier being frozen to its bed (cf. Reference Bælum and BennBælum and Benn, 2011).
Development of glaciological structures
Evidence for former dynamic flow activity of Tellbreen is also provided by the sequential development of glaciological structures, illustrated schematically in Figure 10, which records the passage of a ‘parcel’ of ice from the accumulation area to the glacier front.
Schematic diagram of structural formation of Tellbreen and associated geomorphology. A = longitudinal ridge, B = transverse ridge, C = sand/gravel piles. Parts of this figure are adapted from Reference Hambrey and GlasserHambrey and Glasser (2003).
Primary stratification and longitudinal foliation
The development of primary stratification (S0) occurs in the accumulation area as englacial ice forms due to seasonal firnification processes (Fig. 10a; Reference HambreyHambrey and others, 2005). There is considerable evidence of the deformation of stratification, typically in the form of (1) folding accompanied by simple shear due to lateral compression, which eventually leads to the formation of S1 longitudinal foliation as stratification is at first gently (Fig. 10b) and then more tightly (Fig. 10c) folded in response to compressive stresses caused by a change in flow geometry (e.g. from accumulation area to topographically confined tongue); (2) offset layers of S0/S1 due to faulting associated with shear plane development (Fig. 10d); and (3) the erasing of layering due to intense shearing and strain-induced metamorphism of englacial ice close to the glacier bed, which causes the expulsion of gases (e.g. from bubble-rich layers) along grain boundaries and forms dispersed facies ice (Fig. 10e). Rockfall debris buried within primary stratification in the accumulation area (Fig. 10a), which is transferred to the lower tongue and breaches the surface as a result of folding, forms longitudinal supraglacial ridges or debris septa (Figs 8, 9a and10c). While this transfer of mass remains active (i.e. ice is flowing, however slowly), the folding of stratification and development of foliation will also continue and, therefore, some of these structures may, for a little while longer at least, be actively forming (cf. Reference HambreyHambrey and others, 2005).
Arcuate shear planes
Arcuate shear planes (S2) are located on the lower tongue towards the glacier front (Fig. 8), and in the ice caves are observed to offset primary stratification/longitudinal foliation (S0/S1) but not crevasse traces (S3), indicating that formation of S2 pre-dates that of S3 (Fig. 10d). The development of shear planes within polythermal valley glaciers is consistent with longitudinal compression at the margin leading to brittle failure and thrusting (Reference HambreyHambrey and others, 2005; Reference Roberson and HubbardRoberson and Hubbard, 2010), and at Tellbreen this is supported by the vertical displacement of S0/S1 across S2 planes, indicating high compressive stresses parallel to ice flow direction (Reference SouchezSouchez, 1967; Reference Rees and ArnoldRees and Arnold, 2007). However, the predominant orientation of S2 structures measured at the SW sections shows an approximate offset of 20° relative to ice flow direction (Fig 6d), rather than the perpendicular relationship that might be expected in a purely compressional zone. The presence of sub-horizontal stretching lineations along these fractures (Fig. 6g) reveals that, rather than purely through dip–slip movement, strain along these fractures has been accommodated through a component of strike–slip (e.g. Reference FlemingFleming and others, 2013). Thus, the orientation of these structures suggests that the marginal areas of the glacier have experienced a transpressional stress regime (cf. Reference Twiss and MooresTwiss and Moores, 2007; Reference FlemingFleming and others, 2013). The development of shear planes may be facilitated by the presence of preexisting structural weaknesses, such as fractures relating to reoriented healed crevasses (Reference NyeNye, 1952; Reference Hambrey and MüllerHambrey and Müller, 1978; Reference Rea and EvansRea and Evans, 2011), which are ubiquitous on the lower tongue of Tellbreen (S3 fracture traces; Fig. 8). The similar character and sediment composition of both S2 and S3 structures appears to provide support for this, particularly as S2 structures display evidence for shear and S3 structures do not, indicating that the former may have developed through the reorientation and shearing of the latter (Fig. 10d). However, in the caves, S3 crevasse traces cut across and are not offset by S2 shear planes, demonstrating that S3 formed after S2 structures. The implications of this are that either: (1) S2 shear planes represent an entirely new generation of structures unrelated to any inherited weaknesses (e.g. Reference Glasser, Hambrey, Etienne, Jansson and PetterssonGlasser and others, 2003; Reference Roberson and HubbardRoberson and Hubbard, 2010) or (2) some S3 crevasse traces (although not those exposed in the ice caves) are inherited from an earlier phase of extensional flow, pre-dating S2 formation.
It seems very unlikely that shear planes are still active within Tellbreen given its current cold-based and near-stagnant flow conditions and the high compressive/transpressive stresses which are necessary to form them. The presence of S2 therefore provides compelling evidence that the glacier has been much more dynamic and subjected to high strain rates in the past (cf. Reference HambreyHambrey and others, 2005; Reference Bælum and BennBælum and Benn, 2011). The formation of shear planes in polythermal glaciers is often suggested to occur at the thermal transition between active warm-based ice and inactive cold-based ice at the margin (Reference RippinRippin and others, 2003; Reference King, Smith, Murray and StuartKing and others, 2008), where compressive stresses might be expected to be highest. However, the validity of this was questioned by Reference Moore, Iverson and CohenMoore and others (2010, Reference Moore, Iverson, Brugger, Cohen, Hooyer and Jansson2011), who found no evidence for a sharp slip/no-slip boundary at the thermal transition within Storglaciären, Sweden, and suggested that longitudinal compression at this boundary may be insufficient to generate compressive fractures. Reference Moore, Iverson and CohenMoore and others (2010) proposed that the required conditions for compressive fracturing are likely to be met only by thin (but actively flowing) glaciers subjected to high compressive stresses and with an abundance of pre-existing weaknesses, which characterizes a surge-type glacier in its active phase. The necessary high compressive stresses would be generated at the transition between areas of extremely weak bed, facilitated by high subglacial water pressures and where ice is temperate and actively surging, and inactive, cold, non-surging ice towards the glacier margin. The presence of silt-sized debris consistent with thin films of waterborne sediment within shear planes supports the suggestion that their development requires hydraulic communication with highly pressurized water (Reference Moore, Iverson and CohenMoore and others, 2010). This provides further evidence that they formed at a time when Tellbreen was experiencing dynamic, warm-based ice flow and elevated subglacial water pressures, in stark contrast to the conditions at the bed of the current cold-based and largely inactive glacier.
Extensional crevassing
The prevalence of S3 fracture traces, interpreted as healed crevasses, within the lower tongue of Tellbreen indicates that the glacier has also experienced significant longitudinal extensional stresses, and the relationship between S2 shear planes and S3 crevasse traces observed in the caves (and outlined above) shows that the main phase of extensional flow post-dates the formation of these shear planes (Fig. 10d). The dense population of crevasse traces on the lower tongue suggests that the glacier has been heavily crevassed, since when the crevasses have closed and been transported passively down-glacier as fracture traces (e.g. Reference HambreyHambrey and others, 2005). Extensional crevasses are clearly not forming within the currently cold-based and inactive glacier, and therefore the crevasse traces must have originally formed as open crevasses at a time when the glacier was experiencing far more dynamic ice flow, resulting in significant longitudinal stretching of the ice. Such conditions are consistent with enhanced velocities associated with warm-based ice flow within a thicker glacier at its LIA maximum (Reference HambreyHambrey and others, 2005).
The relationship between S2 and S3 structures indicates that the ice underwent compressive flow followed by extension, consistent with the down-glacier passage of a surge front or kinematic wave (cf. Reference Lawson, Sharp, Hambrey, Maltman, Hubbard and HambreyLawson and others, 2000). The presence of waterborne sediments within crevasse traces, ranging from well-sorted silt-sized material up to sand and gravel, also indicates that the crevasse network was exploited by highly pressurized water, possibly sourced from the temperate bed (Fig. 10e; Reference Evans and ReaEvans and Rea, 1999; Reference Ensminger, Alley, Evenson, Lawson and LarsonEnsminger and others, 2001); these crevasse infills can subsequently melt out on the glacier surface (Fig. 9h). Although there are open fractures (S4) within Tellbreen (Fig. 8), these are either related to localized collapse due to margin undercutting by meltwater action, or bergschrund-type crevassing on steep slopes in the accumulation area (Reference Bælum and BennBælum and Benn, 2011), and so do not reflect current flow dynamics.
LIA thermal switches and changes to glacier dynamics
The basal ice and till facies and structural evidence at Tellbreen provides a clear indication that this currently thin and cold-based glacier was much more dynamic in the past. Although direct dating evidence is unavailable, it is most likely that warm-based dynamic flow occurred at the Neoglacial maximum of the glacier, towards the end of the LIA, when Tellbreen was more extensive, contained ∼60–70% more volume and may have been up to 200 m thick (Reference Bælum and BennBælum and Benn, 2011). The LIA represents the culmination of Neoglacial glacier expansion, which also typically records the most extensive position reached by glaciers during the Holocene (Reference Humlum, Elberling, Hormes, Fjordheim, Hansen and HeinemeierHumlum and others, 2005). Prior to this, conditions during the early and mid-Holocene are suggested to have been warmer than the present-day climate (Reference Salvigsen, Forman and MillerSalvigsen and others, 1992; Reference SalvigsenSalvigsen, 2002), and most current valley glaciers are likely to have been considerably smaller (Reference Svendsen and MangerudSvendsen and Mangerud, 1997; Reference Forwick and VorrenForwick and Vorren, 2007; Reference Mangerud and LandvikMangerud and Landvik, 2007) or even non-existent (Reference Ingólfsson, Martini, French and Perez-AlbertiIngólfsson, 2011) at this time. In the latter case, valley glacier build-up during the early part of the LIA would have been characterized by snow accumulation on permafrost initially leading to the development of cold-based ice; these conditions have persisted beneath some Svalbard glaciers throughout the LIA, as evidenced by the preservation of in situ plants in a subglacial position at Longyearbreen (Reference Humlum, Elberling, Hormes, Fjordheim, Hansen and HeinemeierHumlum and others, 2005). Thickening and steepening associated with glacier growth would have insulated the ice/bed interface and encouraged strain heating of ice in the basal zone (Reference Blatter and HutterBlatter and Hutter, 1991), creating areas at the pressure-melting-point, which in turn would lead to the generation of meltwater at the bed, facilitating related increases in dynamism in the form of sliding/subglacial deformation and higher ice velocities (Reference Blatter and HutterBlatter and Hutter, 1991; Reference HambreyHambrey and others, 2005; Reference Bælum and BennBælum and Benn, 2011). This build-up of mass and associated conditions at the bed is likely to have allowed Tellbreen to advance to its LIA maximum position, delimited by latero-frontal ice-cored moraines (Fig. 1b). The build-up effectively records a switch from cold-based to warm-based conditions during the LIA. Saturation and deformation of the subglacial till and strain-induced metamorphism of englacial ice close to the bed occurred as a result of this switch, in the latter case leading to the formation of the dispersed ice facies. The period of enhanced flow velocities and warm-based conditions also facilitated the formation of shear planes (S2) and extensional crevasses (precursor to S3). Following the LIA, the step-like increase in warming at the start of the 20th century (Reference Hanssen-Bauer, Solås and SteffensenHanssen-Bauer and others, 1990) initiated consistently negative mass balances, extensive thinning, and retreat of the lower tongue, instigating a thermal transition from a glacier with areas of warm-based ice back to one which is almost entirely frozen to its bed and in a state of low glacier activity. The saturated subglacial traction till froze to the bed as the thermal regime switched back to cold-based, and the extensional crevasses closed up to form crevasse traces (S3).
In simple terms, this represents a switch from a warm-based glacier in a cold period (e.g. LIA) to a cold-based glacier in a warmer period (e.g. today). Similar inferences have been made for a handful of small valley glaciers on Svalbard, either where the switch to an entirely cold-based regime is thought to be complete (e.g. Reference Hodgkins, Hagen and HamranHodgkins and others, 1999; Reference Stuart, Murray, Gamble, Hayes and HodsonStuart and others, 2003; Reference Lukas, Nicholson, Ross and HumlumLukas and others, 2005) or is ongoing (e.g. Reference HambreyHambrey and others, 2005). These examples could be part of a broader trend exhibited by small High Arctic valley glaciers of a post-LIA shift from polythermal to cold-based conditions in a warming climate, which is also likely to have wider implications, such as for hydrogeological systems in Arctic regions (e.g. reduction in groundwater recharge; Reference Haldorsen and HeimHaldorsen and Heim, 1999; Reference HaldorsenHaldorsen and others, 2010; Reference Scheidegger, Bense and GrasbyScheidegger and others, 2012).
This suggested sequence of thermal switches experienced by Tellbreen from the beginning of the LIA to the present day shares many similarities with the model proposed by Reference Fowler, Murray and NgFowler and others (2001) to explain surging of polythermal glaciers. According to that model, build-up of mass in the accumulation area of a largely cold-based glacier at the beginning of its surge cycle results in parts of the bed being raised to the pressure-melting point. This switch to warm-based conditions leads to the production of meltwater and increases subglacial water pressures and pore-water pressures in underlying sediments, resulting in a dramatic shift in glacier dynamics, associated positive feedbacks, and typically glacier advance during the active phase of the surge. Post-surge, the over-extended glacier thins rapidly and begins to freeze to its bed, eventually switching back to largely cold-based conditions. The similarities in the thermal changes and associated flow dynamics experienced by polythermal surge-type glaciers and LIA advances of small valley glaciers on Svalbard have led to some discussion as to whether the latter describe surges or simply ‘normal’ polythermal glacier behaviour. An example of this is the case of Midtre Lovénbreen, which has been classified as both surge-type (Reference LiestølLiestøl, 1988; Reference HansenHansen, 2003) and non-surge-type (Reference HambreyHambrey and others, 2005), and similar discussions have evolved in relation to Austre Lovénbreen (Reference Jiskoot, Murray and BoyleJiskoot and others, 2000; Reference Midgley, Tonkin, Cook and GrahamMidgley and others, 2013) and Marthabreen (Reference Hagen, Liestøl, Roland and JørgensenHagen and others, 1993; Reference Glasser, Bennett and HuddartGlasser and others, 1999). This difference of opinion, however, is largely semantic. The key point is that a major change in glacier dynamics, controlled by changes to the thermal structure of the glacier, occurs during climate cycles in both cases. Some glaciers (e.g. Tellbreen) appear to have experienced a single dynamic cycle during the LIA, whereas other (typically larger) glaciers have undergone several dynamic cycles, in the form of surges. This suggests an underlying dynamical similarity, expressed in different ways by different glaciers.
Conclusions
Evidence for former flow dynamics and changes to the thermal regime of Tellbreen, a cold-based, land-terminating valley glacier in central Spitsbergen, are recorded by glaciological structures and within its basal ice sequence. In common with many other small valley glaciers on Svalbard, Tellbreen has retreated steadily since its LIA maximum position, significantly reducing the volume and areal extent of the glacier (Reference Bælum and BennBælum and Benn, 2011). The following conclusions are based on a combination of investigations within meltwater conduits at the glacier front and structural mapping of the surface from aerial photographs.
The basal sequence consists of a frozen matrix-supported diamict overlain by debris-poor dispersed facies ice. The matrix-supported diamict is interpreted as a frozen sub-glacial traction till which has been highly saturated. The overlying dispersed facies has a tectonic origin, relating to strain-induced metamorphism of englacial ice due to shearing close to the bed. The formation of both facies is consistent with a warm-based thermal regime and the availability of subglacial meltwater.
The sequential development of structures within Tellbreen has been determined, and consists of: (1) the formation of primary stratification through firnification processes in the accumulation area; (2) the folding of primary stratification leading to the development of longitudinal foliation; (3) the formation of arcuate shear planes in the lower glacier tongue in response to compressional and, at the lateral margins, transpressional stress regimes; and (4) the opening of extensional fractures and crevasses and the injection of pressurized meltwater into these, followed by the healing of fractures. Of these, (3) and (4) are strongly indicative of former dynamic, warm-based flow (while none is inconsistent with it).
Both the basal sequence and glaciological structures are consistent with Tellbreen having experienced more-dynamic ice flow in the past, characterized by warm-based conditions, tectonic deformation and the availability of pressurized subglacial meltwater. It is likely that these conditions coincided with the LIA maximum extent of Tellbreen, when it was significantly larger and thicker than today.
This evidence records switches in the thermal regime of Tellbreen, from a small early-LIA cold-based glacier to a polythermal glacier with extensive areas of warm-based ice at the LIA maximum, before returning to a glacier which is almost entirely frozen to its bed post-LIA. This sequence is similar to the thermal switch proposed for polythermal surge-type glaciers, suggesting an underlying dynamical similarity. It is likely that many more small valley glaciers in Svalbard have also experienced, or are currently undergoing, a similar switch in response to climatic warming and consistently negative mass balances since their LIA maxima.
Acknowledgements
H.L. was funded by a UK Natural Environment Research Council (NERC) PhD studentship (NE/I528050/1), the Queen Mary Postgraduate Research Fund, and an Arctic Field Grant from the Research Council of Norway. E.J.F. was funded by a NERC PhD studentship as part of the GAINS (Glacial Activity in Neoproterozoic Svalbard) grant (NE/H004963/1). K.N. was funded by an Arctic Field Grant, the Swiss Society for Speleology, and the travel grant commission of the Swiss Academy of Science. Landsat satellite images were provided by US Geological Survey (USGS) Earth Explorer, and aerial photographs were acquired from Norsk Polarinstitutt. We thank Ian Boomer for isotope analyses, UNIS logistics (particularly Martin Indreiten, Jukka Pekka Ikonen and Monika Votvik) and Philipp Schuppli for fieldwork support, and all members of the UNIS AG-325 Glaciology course who were present when we first visited Tellbreen in 2011. Andy Hodson provided helpful advice on various aspects of the work. The paper also benefited significantly from thorough and constructive reviews by Dave Evans and Richard Waller. We thank the International Glaciological Society for support with the publication of this paper.
