Rockglacier activity during the Last Glacial–Interglacial transition and Holocene spring snowmelting

ARTICLE IN PRESS

Quaternary Science Reviews 26 (2007) 793–807

Rockglacier activity during the Last Glacial–Interglacial transition and Holocene spring snowmelting Øyvind Paaschea,, Svein Olaf Dahla,b, Reidar Løvliec, Jostein Bakkea,b, Atle Nesjea,c a

Bjerknes Centre for Climate Research, Alle´gt. 55, N-5007, Bergen, Norway Department of Geography, University of Bergen, Fosswinckelsgt. 6, N-5020, Bergen, Norway c Department of Earth Science, University of Bergen, Alle´gt. 41, N-5007 Bergen, Norway

b

Received 7 November 2005; received in revised form 15 November 2006; accepted 16 November 2006

Abstract The environmental history of a talus-derived rockglacier located in northern Norway has been reconstructed through the Last Glacial–Interglacial transition based on two cores retrieved from an adjacent lake. The methods used to quantify sedimentary properties include rock magnetism, grain size analyses, loss-on-ignition (LOI) and bulk density, which when combined has enabled an unmixing of the various sediment components and their corresponding sources. Rockglaciers signify mean annual air temperatures (MAAT) of 4 1C or colder, but little is known about their dynamical response to changing thermal regimes. We document here for the first time that a permafrost regime did exist in northern Norway during the lateglacial period, and that it required a lowering equivalent of at least 7 1C compared to present-day MAAT. The lake sediments suggest that the rockglacier existed prior to the local deglaciation of the Fennoscandian Ice Sheet (414 800 cal yr BP), and continued its expansion until the end of the Younger Dryas whereupon it became fossil. The cool climate of the lateglacial was intersected by brief warming spells that caused a systematic release of sedimentladen meltwater from the rockglacier. During the Holocene the minerogenic influx to the lake was driven by spring snowmelting, which are related to the magnitude of winter precipitation. Three phases are recognised: (1) 9800–6500 cal yr BP when wet winters prevailed, (2) 6500–4000 cal yr BP with dry winters, and (3) the last 4000 cal yr BP with a return to wetter winters. r 2006 Elsevier Ltd. All rights reserved.

1. Introduction Rapid degrading permafrost followed by thaw subsidence represents a major challenge for policy makers in the coming century, a process that is escalating in the Arctic region (Overpeck et al., 1997; Moritz et al., 2002; Christensen et al., 2004; Smol et al., 2005). But such large-scale climate adjustments are not new, although for different reasons, and complementary analogues do exist in the lateglacial–Holocene record. During the lateglacial period, widespread areas above (i.e. nunataks) and outside the margins of the Fennoscandian Ice Sheet (FIS) were probably subjected to permafrost conditions (MAAT of 4 1C or colder), which most likely started to melt very rapidly at the onset of the Holocene as it did in northern

Corresponding author. Tel.: +47 55 583297; fax: +47 55 589803.

E-mail address: [email protected] (Ø. Paasche). 0277-3791/$ - see front matter r 2006 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2006.11.017

Siberia (see for instance Smith et al., 2004). Unfortunately, such scenarios have been difficult to observe in the Arctic due to the scarcity of relevant sites and suitable proxies. Lakes are one of the few robust terrestrial archives found in the Arctic that can record continuous environmental change over longer time periods (104 yrs) (e.g. Bradley et al., 1996), and is therefore well-suited for tracking sedimentary changes associated with degrading permafrost regimes. The connection between rockglaciers and lake sediments has only briefly been investigated (Zielinski, 1989) and current knowledge of how active rockglaciers interact with climate are primarily based on modern observations (see Humlum, 1997, 1998a). Talus-derived rockglaciers (referred to as rockglaciers hereafter) are unambiguous indicators of permafrost conditions (e.g. Barsch, 1996) and contain an inner body of ice that is carpeted by sheltering rocks. During summer heat penetrates the rocky surface and melts parts of the ice. Water-output from

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active rockglaciers are probably manifested through subsurface water currents, flowing along the upper ice surface in addition to subsurface discharge/seepage (Giardino et al., 1992) before joining downstream rivers and/or enter lakes. The source for such exit water is most likely balanced between melting of the rockglacier’s internal ice body, groundwater seepage and additional runoff from the catchment. Monitoring rockglacier runoff shows that peak discharges tend to co-vary with maximum air temperature during the summer season (Berger et al., 2005), demonstrating a close coupling between rockglacier activity and climate. A prevailing warmer climate will create disequilibrium between melting and refreezing and the rockglacier will lose its internal ice mass and eventually become fossil. Since it is mainly the MAAT that constrain rockglacier activity they may complement traditional summer temperature proxies (pollen, chironomides, plant macrofossils and so on) with additional paleoclimatic information (Haeberli, 1985; Humlum, 1998b; Sailer and Kerschner, 1999; Frauenfelder and Ka¨a¨b, 2000). Rockglaciers are lobe or tongue shaped landforms with steep front sides and near flat plateaus that typically continue into one or several parent talus sheets or cones (Wahrhaftig and Cox, 1957). The surface texture is usually dominated by angular material of varying size, which in some cases make up transverse ridges. The internal flow patterns of rockglaciers are complex with differential flowline trajectories along multiple shear planes that concentrate horizontal movement (Arenson et al., 2002). The presence of sediments and fluid water may affect shear stress (i.e. deformation) within the main ice body that can make up between 30% and 70% of the overall rockglacier volume (e.g. Barsch, 1996). Grain size analysis of sediments recovered from drilling rockglaciers shows a predominant mixture of sand and silt (Barsch et al., 1979; Arenson et al., 2002), whereas measures of sediment load in meltwater discharge from an Austrian rockglacier shows that during maximum discharge of 300 l s1 the sediment load is close to 1 g l1 (Krainer and Mostler, 2002). Discharge from rockglaciers is related to both total size and ice-volume and is known to vary from 1 to 100 l min1 (e.g. Giardino et al., 1992), but higher volumes have also been observed (Gardner and Bajewsky, 1987; Krainer and Mostler, 2002). The parent talus from which rockglaciers evolve constrains the sediment supply and production that is available for later reworking and meltwater transport. The sediment texture of taluses reflects local bedrock and most likely climate-related processes such as frequency of snow avalanches, debris flows, precipitation and temperature (Rapp, 1960; Ballantyne and Harris, 1994). Analyses of relict taluses in Scotland show that almost 30% of the weighted material is composed of sediments being o2 mm (Salt and Ballantyne, 1997; Hinchliffe et al., 1998) suggesting that fine grained sediments are available for entrainment by meltwater.

Geophysical methods have successfully revealed important aspects of the internal structure of rockglaciers (e.g. Isaksen et al., 2000), but knowledge of how they interact with changing climate conditions are poor due to few and short instrumental records. In this study, two cores with sediments partly originating from a rockglacier have been retrieved from Trollvatnet, a small lake in northern Norway. This setting has enabled the unique possibility to observe how these landforms react to changing climate conditions during periods with assumed permafrost, as well as the transition from a glacial to an interglacial climate. The Holocene sequence following the rockglacier fossilisation reveals a promising record of spring snowmelting, which seems to be closely related to the amount of winter precipitation. 2. Site and bedrock The study site is located in northern Lyngen, northern Norway (691380 N, 191530 E) on a north–south stretching peninsula with Ullsfjorden to the west and Lyngenfjorden to the east (Fig. 1). Alpine peaks, empty cirques and temperate glaciers dominate the landscape together with intersecting valleys with relict rockglaciers, moraines and large talus cones. The mean annual air temperature is 2.8 1C (Aune, 1993), whereas the mean annual precipitation is between 700 and 900 mm (Førland, 1993). The present climate may hence be defined as being semi-maritime where relatively cold and wet winters is balanced by mild summers. Polythermal plateau glaciers located in the mountains of southern Lyngen at 1800–1600 m altitude (Gordon et al., 1988) suggest a current local permafrost boundary in the vicinity of 1500 m. The upper western flank of the study area is dominated by gabbroic rocks that originate from layered opholites, while the lower flank is made up of different metasedimentary groups, typically with intrusions of quartzite and mica schist. Two bands of conglomeratic calciferous mica schist and conglomeratic quartzite runs right through Trollvatnet, whereas the taluses above the lake’s western side consists of gabbro intersected by thin layers of periodite (Zwaan et al., 1998). 3. Previous work and deglaciation history The basic morphological characteristics of the Trollvatnet rockglacier have previously been mapped and it was classified as an inactive/relict ‘valley-wall’ complex (Maclean, 1991). The rockglacier is positioned outside two established stages of the FIS in northern Norway, known as ‘Skarpnes’ (Older Dryas) 14 200 cal yr BP and ‘Tromsø-Lyngen’ (Younger Dryas) 12 500 cal yr BP (e.g. Andersen, 1968; Plassen and Vorren, 2003) (Fig. 1). Recent work document that most of the lower parts of northern Lyngen was deglaciated by the Bølling Interstade (Bakke et al., 2005) suggesting that large areas were ice free by the onset of the lateglacial. The physiographic setting of the

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Fig. 1. The overview map on the left-hand side shows the ice margins during the lateglacial in northern Norway (and in more detail on the right-hand side) represented by two prominent stages known as Skarpnes (14 200 cal yr BP) and Tromsø-Lyngen (12 500 cal yr BP), as suggested by Holms and Andersen, 1964; Andersen, 1968; Vorren and Elvsborg, 1979; Fimreite et al., 2002; and unpublished data of the authors. The map on the right-hand side shows that lake Trollvatnet and the investigated rockglacier lie just outside these two ice margins. The contour interval is 20 m (grey lines) between 0 and 200 m altitude and 100 m between 200 and 1400 m altitude.

low-relief landscape (o150 m altitude) where Trollvatnet and Fiskvatnet are located suggests that whenever a glacier occupied Ullsfjorden, meltwater was rerouted northwards (Fig. 1). This has been shown to be the case for the Skarpnes event, but not during the Tromsø-Lyngen stage (Paasche et al., 2004).

cores were disturbed during transport they were not sampled for any parameters. A Garmin GPS with build-in altimeter was used for the purpose of field mapping in combination with aerial photo interpretation (scale 1: 40,000). 4.2. Grain size distribution

4. Methods 4.1. Coring and mapping A piston corer (110 mm) (Nesje, 1992) operated from a mobile raft was used to retrieve two near 6 m long cores from Lake Trollvatnet (92 m altitude) during the summer of 2000. The bathymetric map of the lake is based on 8 transects made with continuous echo-sounding measurements (Fig. 2). The two cores (referred to as T-1 and T-2 hereafter) are taken some 100 m apart within the deepest part of the basin. Grain size analyses, magnetic susceptibility, loss-on-ignition (LOI) and bulk density was performed on core T-1, whereas standard rock magnetic parameters was obtained from core T-2. Material for radiocarbon dating was collected from both cores. Because the lowermost three metres of both

Grain size analyses were performed with 1 cm resolution for intervals between 0–127 and 240–266 cm depth, and with 0.5 cm resolution for 266.5–271.5 and 127.5–239.5 cm depth (n ¼ 380). All samples were wet sieved for the o63 mm fraction, mixed with a 0.05% calgon (sodium hexametaphosphate) dispersing solution and left to homogenise and deflocculate for 412 h before analysed. Considering that some material was lost during the sieving, and that the samples initially were wet when weighted, the final sample concentration is probably somewhat lower than 40 g L1 or 4% by volume. Subsequently, all samples were measured for grain sizes 65–1.5 mm on a Sedigraph 5100, being the optimal size range for this instrument (e.g. McCave and Syvitski, 1991). All statistical processing of the raw grain size data were analysed with GRADISTAT (Blott and Pye, 2001).

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susceptibility (c) was determined on a KLY-2 induction bridge (sensitivity: 4  108 SI) on both wet and freezedried samples. The wet samples were corrected for the diamagnetic effect of water (0.9  108 m3 kg1). Frequency-dependant susceptibility was measured on a Bartington MS2B dual-frequency sensor. Anhysteretic remanent magnetisation (ARM) was imposed with a 0.1 mT DC field and a 100 mT AC-field in a 2G afdemagnetiser, whereas remanent coercivity was obtained by imposing selected samples to progressively higher magnetic fields. A solenoid was used to 175 mT succeeded by a pulse magnetiser (Redcliffe, maximum field 4.2 T). Imposed remanent magnetisation was measured either on a Digico spinner (noise level: 5  107 A m1) or a 3-axes Cryogenic magnetometer (CCL 350/450, noise level 3  108 A m1). Magnetic susceptibility was also measured on split cores from both T-1 and T-2 using a Bartington equipped with a MS2E surface sensor. Moreover, susceptibility on T-1 was measured on plastic bags for every 0.5 cm at the KLY-2 induction bridge with the exception of the dated intervals. 5. Results 5.1. Field mapping

Fig. 2. Shows the simple bathymetry of lake Trollvatnet, the coring sites (black stares) and the position of the front lobe of the rockglacier as well as the inflow and outflow. The area surrounding Trollvatnet include steep taluses, mountain ravines and barren bedrock. The mountain ravines high above the taluses store snowfields that discharge meltwater throughout the summer, relocating material within the talus and eventually bringing the finest fractions down to the lake.

4.3. Organic matter and radiocarbon dating LOI was obtained by extracting 1 cm3 of material with a plastic syringe at 0.5 cm interval. The samples were dried overnight at 105 1C before heated in a furnace to 550 1C for 1 h. LOI determined after heating to 550 1C gives an approximation of the total organic matter and is, if divided by a constant (2.1), indicative of the organic carbon present in the sample (Dean, 1974). Bulk density values are calculated based on the aforementioned approach. Plant macrofossils used for radiocarbon dating were identified, washed and purified with distilled water before dried overnight and stored in airtight glasses. The bulk samples were dried overnight at 100 1C and put in airtight plastic bags. Beta Analytic Inc. and the Poznan Radiocarbon Laboratory carried out the radiocarbon dating. All dates have been converted to calendar years using CALIB 4.1 (Stuiver et al., 1998).

The rockglacier above Trollvatnet is recognised by a series of diagnostic features typical for such a landform (Fig. 3). A concentric, steep front (35–401) delimits the lobe-shaped terminus (30–40 m high) that consists of large interlocked angular clasts overlying finer fractions as evident from exposures after minor fallouts. The lobe plateau is 600 m long and 500 m wide with a low mean surface slope angle (o51), and transverse ridges lying perpendicular to the general surface flow direction. The outer five ridges are crammed together, whereas the inner ridges are deflected or broken. The ridges are characterized by differentiated sorting. Circular depressions of several metres in both diameter and depth are present; the largest is located on the inner part of the lobe plateau. The parent coalescing talus cones show a profound fall sorting which causes the largest boulders to form a rim that runs along its lower part. The exact boundary between the inner plateau and the talus foot is therefore somewhat unclear. Perennial snowdrift accumulations are located in the mountain above Trollvatnet at some 1000 m altitude, and produce meltwater that enters the lake throughout the summer. Several of the adjacent cones inhabit distinct debris flow leve´es. A smaller rockglacier has partly been built onto and into the Trollvatnet rockglacier, and a third more immature rockglacier has formed on top of the northern side of the second complex.

4.4. Magnetic parameters 5.2. Core stratigraphy and intervariability Material used for magnetic analyses on core T-2 was sampled with cubic plastic boxes (2  2  1.6 cm3) for every 1 cm throughout the core (n ¼ 250). Initial magnetic

Visual interpretation of the two cores (T-1 and T-2) suggests that the lower three metres of the near 6-m long

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Fig. 3. Overview photography showing the rockglacier. An alluvial fan, the southern inflow and the outflow of lake Trollvatnet can also be observed. Lake Fiskvatnet (data presented in Paasche et al., 2004) lies in the upper left centre of the photography and Ullsfjorden can be seen in the upper panel. It is generally thought that the deformation of rockglaciers is very conservative and that the internal movement results in transverse ridges as seen on the plateau of this rockglacier.

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Depth (cm) Fig. 4. The intervariability between the two cores (T-1 and T-2) as evident from the w bulk records suggests that the sedimentation rate was similar in this part of the basin, but that there were significant differences in terms of magnitude and frequency. Note that the scales on the y-axis are not the same (T-1 should be read on the left y-axis and T-2 on the right y-axis), but that the values for the upper 150 cm as well as for those below 240 cm are approximately the same. The letters M and B denote the dated levels of the two cores, and the M with the asterisk mark a plant macrofossil that did not graphitise during laboratory analyses.

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5.3. Age-depth model A polynomial fit was used to construct the age-depth model combining five of altogether 11 radiocarbon dates (Table 1) obtained from either terrestrial plant macrofossils or bulk samples (Figs. 4 and 5) and on the assumption that the top of the cores represents present day (Fig. 5). No plant macrofossils were recovered from the lower organicdeficient section (150–270 cm), which unfortunately compelled dating on bulk sediment samples. Three bulk samples were also dated in relation to the aforementioned minerogenic band occupying the interval between 145 and 135 cm, in addition to two samples from the upper part of

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cores are completely dominated by grey homogenous silty clay interpreted to be of glacial origin. Layered blackish bands (0.5 cm) are present and increasing from 245 to 235 cm culminating in a blackish interval (223–219 cm) (Fig. 4). Clayey silt with no apparent structure dominates the interval between 219 and 170 cm, whereupon the blackish bands reappear and are present up to 165 cm depth. A gradual transition into a blackish silty gyttja follows from 165 cm, except for a return to a silty layer located between 140 and 130 cm. The remaining part of the core is characterized by a blackish (130–70 cm) to light yellow brown (70–0 cm) gyttja with sporadically stratified layering. Notable grey bands (0.2–0.8 cm) were identified at 3, 10, 19, 22, 25, 32, 90, 94, 101, 108 cm depth. The bulk susceptibility (wbulk) records for T-1 and T-2 are generally in good agreement, though significant differences are evident (Fig. 4). The lowermost section (260–240 cm) is characterized by values of 5  105 m3 kg1 and relatively low variability. A prominent drop in wbulk values is recorded at 240 cm in T-1, while T-1 does not show the same pattern. From 240 to 160 cm the two cores are different in terms of variability and strength. T-1 shows little variability and indicates a relative stable system, while T-2 shows a high variability pattern for the same interval. The lowest wbulk values in core T-2 are the same as those for T-1 indicating that they are recording the same background signal, but that T-2 receives material that escapes T-1 perhaps due to sediment focussing. A marked drop in the cbulk signal occurred synchronously in both cores at 160 cm and values of 3  107 m3 kg1 are reached at 150 cm. A pronounced peak in magnetic susceptibility shows up in both records between 145 and 135 cm. The initiation of this excursion occurs at exact the same depth, but the duration is slightly different. The subsequent phase from 130 to 80 cm shows relatively large amplitude changes in magnetic susceptibility that is partly out-of-phase. But for the remaining 80 cm of the core the wbulk signal of T-1 and T-2 is in-phase in terms of both strength and variability. The similarity between the two cores, with significant changes occurring at almost exactly the same depth, suggests a uniform sediment accumulation in this part of the basin.

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Fig. 5. A polynomial fit was used to establish a shared age-depth model for the two cores retrieved from Trollvatnet (T-1 and T-2) (closed circles). The three youngest dates were obtained from plant macrofossils, while the two oldest dates was obtained from bulk samples (Table 1). Open squares represent rejected dates obtained from bulk samples. The difference between a polynomial fit and linear interpolation illustrates the effect this choice can have on the sedimentation rate (values can be read from the second y-axis).

the core. The two oldest bulk dates (12 930790 and 12 700790 14C yr BP) taken 11 cm apart are slightly inversed. They were extracted from a sequence with layered sedimentary structures suggesting no deformation due to coring or transport. The oldest date is therefore rejected (accepting it would have little bearing on the age-depth model). These old dates are located within a well-known 14 C plateau centring on 13 000–12 500 14C yr BP (e.g. Stuiver et al., 1998). The three bulk dates collected above and below the minerogenic band located between 145 and 135 cm yield approximately the same age of 11 430790 14C yr BP suggesting a near-constant contamination of old carbon. If this contamination was solely due to a hard-water effect then a larger spread of ages might be expected, since the initial age of the organic matter would add to the final age. One explanation for this phenomenon might be related to the melting of rockglacier ice containing very old CO2 that would saturate the upper sediment column (e.g. Bjo¨rck and Wohlfarth, 2001). The difference between the younger bulk sediment samples and plant macrofossils confirms a Holocene hard-water error, but here the initial age of the organic matter seems to be included (Fig. 5). Groundwater seepage through the calcareous bedrock within the watershed was possibly restricted during the time of the older bulk dates due to the presence of permafrost, which may explain why the lateglacial bulk dates appears to be more or less correct. The similarity between the two cores,

ARTICLE IN PRESS Ø. Paasche et al. / Quaternary Science Reviews 26 (2007) 793–807 χARM (m3 kg-1) 10 -8

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Fig. 6. Down core variations (T-2) in various concentration dependant magnetic parameters (wbulk, wAMR, sSIRM), magnetic mineralogy (IRM100mT/ IRM3T, referred to the S-ratio in the text), magnetic grain size indicator (sARM/sSIRM), organic matter proxy (LOI) and bulk density measurements together with the corresponding ages. Notice the pronounced jump in bulk density values between 230 and 240 cm depth. The two melt-out-phases (MOP) are highlighted in grey.

as discussed above, permits the construction of a combined age-depth model. 5.4. Parameterization of lacustrine sediments An overview of various magnetic parameters, LOI (%) and bulk density (g cm3) are given in Fig. 6, whereas grain size data are given in Fig. 7. All parameters reproduce the same major sedimentary changes that occurred at 14 800 cal yr BP, 11 600 cal yr BP and 11 100 cal yr BP. The magnetic grain size, as inferred from the wARM/ sSIRM ratio (e.g. Moskowitz et al., 1993) and hysteresis ratios (Hcr/Hc values from 1.68 to 8.62 and Mrs/Mr values from 0.06 to 0.30 (see Paasche et al., 2004 for data)) define two main groups: the lateglacial sequence which indicates the presence of pseudo-single domain (PSD), and sediments spanning the last 10 000 cal yr BP where singledomain (SD) appears to dominate (Fig. 8). The relationship between LOI (%) and wbulk can similarly be divided into two main populations and expressed statistically as shown in Fig. 9. The magnetic susceptibility signal probably reflects the degree of detrital influx where high wbulk suggests a dominance of minerogenic matter. A reduced influence of minerogenic influx over LOI (%) combined with more favourable conditions for lake

productivity during the last 10 000 cal yr BP is evident from Fig. 9. After 4000 cal yr BP LOI values tend to weaken towards the present. The magnetic mineralogy, as seen in the S-ratio values (Stober and Thompson, 1979), indicates a non-magnetite pre-Holocene fraction (S-ratio of o0.9) and a magnetite (S-ratio of 40.9) dominated Holocene sequence (Fig. 6). The latter phase shows low amplitudes and low frequency, as compared to high frequency and high amplitudes in the former sequence. The grain size distribution is generally poorly sorted with a polymodal spread and a mean ranging from very fine-tofine silt (fine skewed). Variations in clay and very fine silt dominate the mean grain size (r2 ¼ 0.9), though with less influence during the Holocene where the tendency shifts towards coarser fractions. Two prominent events occur at 14 800 and 10 500 cal yr BP, where medium silt prevails together with coarse silt. The initiation of the former event at 15 500 cal yr BP also marks the termination of a gradual reduction in the mean grain size. Internal correlations of the different grain sizes reveal a positive relationship of clay (o2 mm) versus very fine silt (2–4 mm) for the entire record (r2 ¼ 0.81). Medium silt (8–16 mm) versus coarse silt (16–31 mm) shows a positive relationship during the lateglacial period (r2 ¼ 0.70) and a negative correlation

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Fig. 7. Down core variations in percentages of the different grain size fractions (very coarse silt 63–31 mm, coarse silt 31–16 mm, medium silt 16–8 mm, fine silt 8–4 mm, very fine silt 4–2 mm, and clay o2 mm) and mean (mm) together with the corresponding ages. A more detailed evolvement of the two MOPs is presented in Fig. 12.

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Fig. 8. wARM plotted against sSIRM in a bi-logarithmic plot showing two distinct populations. The one with the steepest gradient is interpreted to represent single domain magnetite, whereas the second population probably represents a cluster of pseudo-single domain carriers.

during the last 7000 cal yr BP (r2 ¼ 0.57). Fine silt and very coarse silt does not correspond significantly against any of the other fractions. A consistent negative correlation exists

6.1. The deglaciation transition The retreat of the FIS from the Skarpnes stage in Ullsfjorden (see Fig. 1) routed sedimentladen meltwater northwards that was partly deposited in Fiskvatnet (Paasche et al., 2004), but also in Trollvatnet, as evident from the lower 3 m of homogeneous sediments. The signature of these glacigenic deposits is significantly different from later sediments with specifically high wbulk values (Fig. 4), high bulk density values, very stable sARM/sSIRM values (Figs. 6 and 11), high content of very coarse silt and very low content of clay values (Fig. 7). The transition from a deglaciation and to a sedimentary regime dominated by the rockglacier occurred abruptly, indicating that the fjordglacier retreated rapidly after releasing huge quantities of meltwater, and that it did not readvance to a similar position during the remaining period of the lateglacial. Articulated black ‘laminations’ (2–5 mm) are present from 15 700 cal yr BP and they culminate in a 3 cm thick black coloured section with peak LOI-values of 4.5% dated to 14 800 cal yr BP (Fig. 6 and Table 1). The

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-8

10

-7

10

-6

25

25 LOI:125-260 cm LOI:0-125 cm Power (L OI:125-260 cm)

20

20

Power (L OI:0-125 cm)

15

15

LOI (%)

y = 0.0002x -0.6777 r

2

Holocene

10 y = 2E-05x -0.7738 r 2 = 0.85

5

5

0

0

10

-8

-7

10 χbulk (m3 kg-1)

10

-6

Fig. 9. A semi-logarithmic plot of LOI plotted against wbulk. The correspondence between the two parameters indicates to what degree the minerogenic influx suppresses the production of organic material. Two populations can be statistically expressed, representing the lateglacial and the Holocene fractions.

0

20

40

60

80

35

35

0-271 cm

y = -1.0767x + 42.788

30

30

Anomaly-1

r 2 = 0.7644

Clay (%)

Anomaly-2

25

25

20

20

15

15 y = 1084.3x -1.2376 r 2 = 0.8038

10

10

5

5

0

0 0

20

40 Medium silt%

probably overestimates the age of this transition and that an age close to the Bølling Interstade seems more reasonable. An early age of the local deglaciation is, however, in accordance with the reconstructions of a glacier chronology less than 10 km north of Trollvatnet (Bakke et al., 2005). 6.2. Rockglacier activity

= 0.6471

10

Lateglacial

801

60

80

Fig. 10. A consistent negative relationship between medium silt and clay exists throughout the core, with the exception of anomaly-1 which represents short-lived events in the late Holocene and anomaly-2 which corresponds to the two melt-out phases (MOP-1 and MOP-2). Higher amount of clay contra medium silt probably indicates a reduced capacity of the sediment carrying meltwater.

comparison between magnetic parameters and d18O values from NGRIP (Johnsen et al., 2001) showed in Fig. 11 (and explained below in Section 6.2) suggests that the age-depth model is accurate from the Bølling Interstade and onwards, but that the exact timing before this event is difficult. The black laminations and the increased LOI values do not agree with a deglaciation scenario when large meltwater pulses would flush the lake, still the age-depth model

Because the age-depth model performs poorly before the Bølling Interstade it is hard to estimate the actual inception of the rockglacier, but considering that it responds to the Bølling warming, recorded as the first melt-out-pulse (MOP-1) (Fig. 7), it must have had some time to develop prior to this period. The sedimentary response defined as MOP-1 is represented by a domination of medium-and coarse silt subsequent to a period with layered bands that marks the first large-scale sedimentary imprint of an active rockglacier in the lateglacial (Figs. 10 and 11). Fine-grained sediments, mostly clay and very fine silt, define the sedimentary feeding from the rockglacier after MOP-1, which indicates a stable sedimentary regime. Superimposed on this stable mode are short-lived sediment-pulses (Fig. 11). This is particular evident from the concentration parameters (Fig. 11) and the wbulk record of T-1 and T-2 (Fig. 4). The wbulk signal shows how these peaks repeatedly reach similar high values suggesting that T-2 perhaps is somewhat closer to the sedimentary source than T-1. Moreover, these shifts are not reproduced by the grain size data, LOI or bulk density, which suggests that even though the wbulk signal is significant, it does not signify a ‘dramatic’ hydrological change, i.e. a major shift of the sedimentation rate. The peaks identified by the magnetic parameters is interpreted to represent sediment carrying meltwater pulses originating from the rockglacier; driven by warmer than average summers, as documented by the NGRIP temperature reconstruction (Fig. 11). Enhanced melting of the internal ice body in combination with drainage from the plateau ponds (see Section 6.3 below) could possibly produce enough meltwater to explain these pulses. In Fig. 11, these short-lived pulses show higher than average sSIRM and sARM values, and importantly, that the sARM/sSIRM ratio shows coarser magnetic grain size being evident of an increased carrying capacity of the meltwater. When the sediment pulses entered the lake they probably caused ‘‘sediment focusing’’, thus explaining the incongruity between core T-1 and T-2. The recurrence interval of the most prominent peaks in wbulk from T-2 ranges from 100 to 220 yrs with higher counts at 160–180 yrs and 200–220 yrs, whereas the duration of these events rests between 90 and 240 yrs. The relationship between clay and very fine silt changes at 12 200 cal yr BP (Fig. 7) indicating a modest shift towards coarser fractions, which could indicate somewhat higher meltwater energy. A pollen reconstruction from

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802 Table 1 Radiocarbon dates from Trollvatnet Depth (cm)

Lab. no.

Material

14

T-1 21 45 70 126 133 140 167 221 232

Poz-452 Poz-3195 Poz-453 Poz-454 Beta-154200 Beta-154201 Beta-154202 Beta-154203 Beta-154204

Bulk sample Macrofossil Bulk sample Bulk sample Bulk sample Bulk sample Bulk sample Bulk sample Bulk sample

3890745 3600730 6980755 11 470780 11 480790 11 340790 10 220790 12 930+90 12 700790

T-2 27 82

Poz-3196 Poz-3197

Macrofossil Macrofossil

2525730 5670735

a

C-age BP

Age cal yr BP

1 sd range

4330b 3890a 7770b 13 450a 13 450a 13 260b 11 940a 15 600a 15 420a

4414–4246 3963–3841 7919–7737 13 778–13 195 13 510–13 290 13 470–13 140 12 340–11 690 15 870–15 020 15 650–14 430

2710a 6420b

2738–2507 6491–6408

The intercept age obtained by CALIB 4.1 (Stuiver et al., 1998). Average of one sigma values are used.

b

-35

10000 10 1

-37

10 -1

11000

12000

13000

14000

15000

16000

0.00 0.01 0.02

-41

-43

0.03 0.04

10 -3

0.05 10 -4

ARM/SIRM

-39

σ SIRM- σ ARM (Am2 kg-1)

δ18O (NGRIP)

10 -2

0.06 0.07

10 -5

0.08 δ18O (NGRIP) σ SIRM σ ARM ARM/SIRM

10 -6

Rockglacier fossilisation

-45

10 -7 10000

11000

12000

13000

14000

15000

0.09

16000

Cal yr BP Fig. 11. The concentration parameters (sARM-sSIRM), the sARM/sSIRM ratio and the d18O from NGRIP (Johnsen et al., 2001) are shown for the lateglacial to Holocene transition on independent age-depth models. The grey bars denote where inferred short-lived meltwater events took place. As evident from the figure they coincide with peaks in ARM and SIRM, as well as with troughs in the sARM/sSIRM ratio. The d18O record from NGRIP indicates that warm anomalies coincide with meltwater pulses, indicating a possible climate forcing mechanism. This relationship is especially good from the onset of Bølling and to the onset of Allerød (13 250 cal yr BP). During Allerød the similarity between the two records is poor, but improves again in the Younger Dryas and the early Holocene. Notice how the SIRM-ARM values lag the prominent warming (d18O) associated with the onset of the Holocene with nearly 400 yrs, indicating perhaps a delay in the thawing of the permafrost (cf. Haeberli, 1984).

Tromsø shows a warming that commenced at the same time and prevailed until 11 450 cal yr BP, whereupon it cooled again (Fimreite et al., 2002). It is, however, not until 11 100 cal yr BP that grain size data shows dramatic

changes and the initiation of the second melt-out-peak (MOP-2) is observed, with average LOI (%) values of 3–5% being comparable to values obtained from proglacial lakes (Souch, 1994).

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During MOP-2 coarse and medium silt reached values close to 80%, bulk density increased, LOI decreased, magnetic susceptibility and concentration parameters increased. The similarity between the grain size structure of MOP-1 and MOP-2 is clearly evident from Fig. 12. But whereas the rockglacier returns to more shallow active layers after MOP-1 the opposite is true for MOP-2 when the Holocene warming prevails and melts the permafrost altogether. The final part of MOP-2 therefore represents the final fossilisation of the rockglacier. The other parameters indicate a somewhat faster response to the external forcing (warming) than the grain size distribution during MOP-2. The S-ratio (sIRM100mT/ sIRM3T) dropped rapidly between 11 525 cal yr BP (S ¼ 0.86) to 11 350 (S ¼ 0.68) yr BP, whereas sSIRM and wbulk falls significantly from 11 350 cal yr BP (1.53  102 A m2 kg1 and 3.36  107 m3 kg1) to 11 150 cal yr BP (1.54  104 A m2 kg1 and 7.88  108 m3 kg1), respectively. Since the process in question is the transition from permafrost to a more temperate regime, the above mentioned lag in the grain size data, may suggest that some ice was still present within the rockglacier even after the Holocene warming commenced. The low-altitude parts of the watershed had, by that time, most likely begun to thaw. The fact that summer temperature reconstructions show a regional cold setback, commonly referred to at the ‘‘Preboreal oscillation’’ (PBO) (Bjo¨rk et al., 1997), occurring between 11 450 and 11 100 cal yr BP, was perhaps strong enough to counterbalance the melting that was

0

200

400

600

800

1000

1200

1400 80

80 Start

70

End

70

MOP2-Coarse silt MOP2-Medium silt MOP1-Coarse silt

60

60

% of total

MOP1-Medium silt

50

50

40

40

30

30

20

20

10

10

0 0

200

400

600 800 Time (years)

1000

1200

0 1400

Fig. 12. Coarse- and medium-silt values for MOP-1 and MOP-2 as a function of time, indicating that the sedimentary response to a warming perturbation is comparable in terms of grain size distribution. MOP-1 reverts to lower values after some 400 yrs, whereas MOP-2 lasts for 1000 yrs. During the latter phase of MOP-2 coarse- and medium silt dominates the grain size distribution completely with around 90% of the total distribution (cf. Barsch et al., 1979). This grain size pattern is probably driven by a melting of the internal ice body of the rockglacier, increasing the sediment-carrying capacity of the meltwater, but perhaps also by higher availability of sediments.

803

initiated at the onset of Holocene. Thus, this intermediate period could have been one with discontinuous permafrost conditions. 6.3. Inferred rockglacier hydrology Based on a series of transects across the rockglacier surface the total volume of the fossil rockglacier was calculated to be 2.11  106 and 1.69–1.58  106 m3 when void space (20–25%) were accounted for (Maclean, 1991). If the interstitial ice represented 30–70% of the total rockglacier volume when active (Arenson et al., 2002) such numbers would imply a hydrological reservoir of anything between 1 477 000 and 633 000 m3 of ice. Considering the relatively few depressions on the plateau, it seems likely that the ice volume in this particular rockglacier did not exceed 40%. In sum, this makes the rockglacier a major aquifer in the catchment during the lateglacial when the winter precipitation was 40–60% lower than today (Bakke et al., 2005), hence reducing the importance of additional runoff. In comparison, the current volume of water in Trollvatnet is approximately 2  106 m3. Depressions located on the lobe plateau act as sinkholes for subsurface drainage, and can direct drainage as they expand (Johnson, 1978; Ka¨a¨b and Haeberli, 2001). Transient meltwater ponds filling such semi-circular depressions are previously reported (Chandler, 1973; Johnson, 1978; Swett et al., 1980) and substantiate a mechanism for the rockglacier to release pulses of meltwater large enough to entrain sediments that can be deposited in the lake. There is also a possibility that entrapped liquid water stored within the rockglacier can be drained instantaneously due to melting of damming ice. This is supported indirectly by tracer experiments demonstrating that the residence time for meltwater within rockglaciers may be considerable (Giardino et al., 1992; Tenthorey, 1992; Harris et al., 1994). The concept of transient sediment storage and reentrainment due to variable runoff capacity is well known from glacier hydrology (e.g. Richards and Moore, 2003), and could be relevant to a rockglacier system as well. The trapping of fine-grained sediments in cavities and clefts in the lower part of the plateau would be prone to redeposition by flowing meltwater. The snout of the rockglacier terminates today well within the lake and it is possible that parts of the discharge network reached the subsurface of the lake creating local depocenters, which can explain some of the difference between T-1 and T-2 (see above). 6.4. Lateglacial paleoclimate Temperature reconstructions from the region such as Tromsø (Fimreite et al., 2002) and Hammerfest (Birks et al., 2005) show mean July temperature estimates of 6–7 1C for the lateglacial, which is comparable to present day July temperatures found in the High Arctic. Since the

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804

rockglacier lies only some 20 m above the maximum lateglacial sea level the permafrost boundary must have been nearly 1400 m lower than today in order to ensure an active rockglacier. Using an environmental lapse rate of 0.6 1C/100 m (see Lie et al., 2003) such a lowering would have caused an 8 1C depression of the MAAT. Subtracting the present day MAAT of 2.8 1C this would produce a lateglacial MAAT of 5 1C, which is in accordance with the regional summer temperature reconstructions (Fimreite et al., 2002; Birks et al., 2005). Such cool temperatures suggest that most of the local glaciers, having equilibriumline-altitudes (ELAs) around 400 m or lower (Bakke et al., 2005) probably were polythermal during this period. A similar altitudinal difference between polythermal glaciers and rockglaciers has also been documented for western Greenland (Humlum, 1988). 6.5. Holocene spring snowmelting and winter precipitation During Holocene, when the rockglacier no longer influenced the sedimentary budget of the lake, the competing magnetic carriers in Trollvatnet were probably of biogenic (internal lake production) and terrigenous (detrital influx) origin (Fig. 8). The presence of bacterial magnetite over the last 10 000 cal yr BP seems to be in line with evidence for SD magnetite (see also Paasche et al.,

0

2000

4000

6000

2004). The domination of two magnetic components with dissimilar qualities facilitates the possibility of an unmixing of the signal through the Holocene by analysing the difference between sARM/sSIRM and the wARM/wbulk (Fig. 13). This is possible as bacterial magnetite (biogenic) yields low magnetic susceptibility and high remanence (e.g. Frankel et al., 1998), whereas the paramagnetic contribution has the opposite quality and may consequently be attributed to influx of detrital minerogenic material, which most likely is controlled by the amount of spring snowmelting. The sARM/sSIRM ratio peaks rapidly just after 10 000 cal yr BP with high amplitude oscillations around 0.1 (Fig. 13). The wARM/wbulk shows somewhat higher values after 10 000 cal yr BP, but it is not until 7000 cal yr BP that it peaks and obtain values higher than 25 (Fig. 13). The lag between the two ratios suggests that the minerogenic influx (i.e. paramagnetic contribution) is relatively high between 10 000 and 7000 cal yr BP, as high wbulk values suppress the wARM values. Between 7000–4000 cal yr BP the difference between wARM/wbulk and sARM/sSIRM become small suggesting a reduced paramagnetic contribution from the catchment. From 4000 cal yr BP and up to present day the difference between the two ratios increases again. The sARM/sSIRM are stable, but a gradual coarsening of the magnetic minerals is

8000

10000

250

0

0.00

Winter precipitation Snowmelting

150 15 100 20

0.04 0.06 0.08

ARM/SIRM

Winter precipitation (%)

10

0.02 Spring snowmelting (χARM/ bulk)

5

ARM/SIRM

200

0.10 0.12

50

25 Dry winters (DW)

Wet winters (WW)

Wet winters

DW

0.14

WW 30

0 0

2000

4000

6000

8000

10000

0.16

Cal yr BP Fig. 13. Reconstructed winter precipitation from northern Lyngen (the 1961–1990 mean in winter precipitation is set as 100% (Bakke et al., 2005)) is shown in comparison with the inferred spring snowmelting signal (wARM/wbulk). During the Holocene the competition between internally produced biogenic magnetite by magnetotactic bacteria and externally influx of detrital matter can be read as the relative difference between wARM/wbulk and the sARM/sSIRM values. The covariance between the spring snowmelting signal and the winter precipitation record is especially good during dry winters, and both reconstructions shows a steady increase in winter precipitation after 5000 cal yr BP.

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evident. The wARM/wbulk values become lower and suggest an increase in influx of minerogenic material from the catchment, especially after 2000 cal yr BP. Comparing the inferred Holocene spring snowmelting with the reconstructed winter precipitation from northern Lyngen (Bakke et al., 2005) it seems that the two is positively correlated—particularly after 7000 cal yr BP (Fig. 13).

4.

7. Conclusions 1. A multi-proxy approach has for the first time been able to connect rockglacier activity to suspended sediments retrieved from an adjacent lake. The sediment record consisting of two near 6 m long piston cores demonstrate how the rockglacier has interacted with changing climate conditions in northern Norway throughout the lateglacial period before becoming fossil at the onset of the Holocene. The sedimentary signature originating from the talus-derived rockglacier is significantly different from the glacigenic sediments flushing the lake during the deglaciation of the FIS, and also from the sediments that dominate the Holocene record. The transition from deglaciation to a rockglacier sedimentary regime occurs sometime before Bølling Interstade (14 800 cal yr BP), but the age-depth control is poor in this part of the record. A sedimentladen melt-waterpulse (MOP-1) dominated by coarse to medium silt is released from the rockglacier in response to the Bølling warming, and marks the first robust sedimentary signal from the rockglacier. 2. Between 14 800 and 11 100 cal yr BP the rockglacier steadily releases suspended sediments and two sedimentary modes are observed: (1) a stable regime defined by the dominance of fine grained sediments (clay to fine silt), and (2) a high-frequent, high-amplitude regime being superimposed on the former system. These two systems are particularly evident in the rock magnetic parameters, where peak values of SIRM-ARM and magnetic susceptibility denotes periods with coarser magnetic grains. The fact that the physical grain size data do not pick up this signal indicates that no major change in the sedimentation rate was associated with these brief pulses. The recurrence interval of these events ranges from 100 to 220 yrs with an approximately similar duration. 3. The frequent pulses punctuating the stable sedimentary mode provides evidence for rockglacier interaction with climate through release of meltwater that can transport suspended sediments. This occurred whenever there were warmer summers than average causing a more intense melting of the internal ice body and hence improving the sediment entrainment. Additional melting of snowbanks in the catchment and draining of temporary meltwater ponds coexisting on the plateau of the rockglacier may have strengthened this signal. If the meltwater from the rockglacier was released as subsurface currents, as documented in modern systems,

5.

6.

7.

805

it could explain the observed sediment focussing in Trollvatnet. The close covariance between warm summers, as seen from the d18O reconstruction of lateglacial temperatures in the NGRIP core, and the corresponding meltwater pulses released from the rockglacier is particularly evident between the Bølling warming and the onset of Allerød, as well as during the late Younger Dryas. This yields strong support to the notion that rockglaciers are capable to interact with climate during cool permafrost conditions. It also corroborates the age-depth model during a period when dating is most difficult. The pronounced temperature increase following the onset of the Holocene causes the rockglacier to become fossil and it remains so up to the present. The details of this process are recorded in MOP-2 when medium to coarse silt almost completely dominates the grain size data. The initiation of MOP-2 lags the onset of the Holocene with nearly 500 yrs as observed in most of the parameters and its duration is perhaps somewhat shorter than the grain size data indicate. The pronounced similarity between MOP-1 and MOP-2 underscores that the warming perturbation of the Bølling reached northern Norway, though returning to a cool climate after 300–400 yrs. The interval between the two MOPs (14 800– 11 600 cal yr BP) was probably dominated by permafrost conditions with MAATs fluctuating around 4 1C. Compared to the present-day climate of Lyngen this would require a lowering of at least 7 1C (cf. Frauenfelder et al., 2001). This corresponds to a lowering of the permafrost boundary with 1400 m, which is in good agreement with the assumed MAAT of 4 1C, but also with the mean summer temperature estimates of 6–7 1C (Fimreite et al., 2002; Birks et al., 2005). With a permafrost boundary close to sea level during the lateglacial it is suggested that most of the local glaciers probably were polythermal during this period. Even though the thermal prerequisites for a permafrost regime ceased to exist with the onset of the Holocene it seems that a period of discontinuous permafrost prevailed until 11 100 cal yr BP, which can be explained by the Preboreal cooling (PBO) lasting from 11 450 to 11 100 cal yr BP. The final thawing of the permafrost in the lower areas of northern Lyngen commenced rapidly after the PBO and is in-line with the early Holocene thermal optimum observed for the sub-arctic region (Kaufman et al., 2004; Smith et al., 2004; Paasche et al., 2004). During the Holocene detrital influx was most likely transported to the lake with the spring snowmelting, which is related to the amount of winter precipitation, and complemented by later melting of high-lying perennial snowfields. By using the wARM/wbulk versus the sARM/sSIRM the paramagnetic contribution from the catchment has been isolated from internal production of magnetic carriers. The reconstruction of the

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spring snowmelting shows that wet winters prevailed in the early Holocene (9800–6500 cal yr BP), followed by a period (6500–4000 cal yr BP) with a somewhat drier climate, whereas the last 4000 yrs shows a return to wetter winters. Acknowledgment ØP would like to thank Joachim Riis Simonsen for assistance in field, Dr. Wenche Eide for picking macrofossils and Dr. Øyvind Lie for introducing the Sedigraph5100. The paper benefited heavily from discussion with NORPEC members as well as from MAPECS forum meetings, and also from insight offered by Dr. Ann Hirt and Prof. Ole Humlum. Finally, we would like to thank the reviewers Wilfred Haeberli and Brian Whalley for helpful and constructive comments. This work was supported by NORPEC, an NFR funded Strategic University Programme (SUP) at the University of Bergen, which was chaired by Prof. J.B.H. Birks. This is publication No. A 153 of the Bjerknes Centre for Climate Research. References Andersen, B.G., 1968. Glacial Geology of Western Troms, North Norway, Norwegian Geological Survey, vol. 256, 160pp. Arenson, L., Hoelzle, M., Springman, S., 2002. Borehole deformation measurements and internal structure of some rock glaciers in Switzerland. Permafrost and Periglacial Processes 13, 117–135. Aune, B., 1993. Air temperature normal, 1961–1990. Report 2, The Norwegian Meteorological Institute. Bakke, J., Dahl, S.O., Paasche, Ø., Løvlie, R., Nesje, A., 2005. Glacier fluctuations, equilibrium-line altitudes and paleoclimate in Lyngen, northern Norway, during the Lateglacial and Holocene. The Holocene 15, 518–540. Ballantyne, C.K., Harris, C., 1994. The Periglaciation of Great Britain. Cambridge University Press, Cambridge. Barsch, D., 1996. Rock glaciers. Indicators for the Present and the Former Geoecology in High Mountain Environments. Springer, Berlin, p. 331. Barsch, D., Fierz, H., Haeberli, W., 1979. Shallow core drilling and borehole measurements in permafrost of an active rock glacier near the Grubengletscher, Wallis, Swiss Alps. Arctic, Antarctitc and Alpine Research 11, 215–228. Berger, J., Krainer, K., Mostler, W., 2005. Dynamics of active rock glacier (O¨tzel Alps, Austria). Quaternary Research 62, 233–242. Birks, H.H., Kiltgaard Kristensen, D.K., Dokken, T.M., Andersson, C., 2005. Explanatory comparisons of quantitative temperature estimates over the last deglaciation in Norway and the Norwegian Sea. In: Drange, H., Dokken, T.M., Furevik, T., Gerdes, R., Berger, W. (Eds.), The Nordic Seas; An Integrated Perspective; Oceanography, Climatology, Biogeochemistry, and Modeling. Geophysical Monograph, vol. 158. AGU, Washington, DC, pp. 341–355. Bjo¨rk, S., Rundgren, M., Ingolfsson, O., Funder, S., 1997. The Preboreal oscillation around the Nordic Seas: terrestrial and lacustrine responses. Journal of Quaternary Science 12, 455–465. Bjo¨rck, S., Wohlfarth, B., 2001. 14C chronostratigraphic techniques in paleolimnology. In: Last, W.M., Smol, J.P. (Eds.), Tracking Environmental Change Using Lake Sediments, Basin Analysis, Coring, and Chronological Techniques, vol. 1. Kluwer Academic Publishers, Dordrecht, pp. 205–245. Blott, S., Pye, K., 2001. Gradistat: a grain size distribution and statistics package for the analysis of unconsolidated sediments. Earth Surface Processes and Landforms 26, 1237–1248.

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