An Overview of Deformation and Erosion Phenomena Related to Advancing Ice Sheets in Lowland Europe

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EXTERNAL REPORT SCK•CEN-ER-254 14/Kbe/P-11

An overview of deformation and erosion phenomena related to advancing ice sheets in lowland Europe

Koen Beerten

SCK•CEN Contract: CO-90-08-2214-00 NIRAS/ONDRAF contract: CCHO 2009-0940000 Research Plan Geosynthesis

February, 2014

SCK•CEN PAS Boeretang 200 BE-2400 Mol Belgium

EXTERNAL REPORT OF THE BELGIAN NUCLEAR RESEARCH CENTRE SCK•CEN-ER-254 14/Kbe/P-11

An overview of deformation and erosion phenomena related to advancing ice sheets in lowland Europe

Koen Beerten

SCK•CEN Contract: CO-90-08-2214-00 NIRAS/ONDRAF contract: CCHO 2009-0940000 Research Plan Geosynthesis

February, 2014 Status: Unclassified ISSN 1782-2335

SCK•CEN Boeretang 200 BE-2400 Mol Belgium

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Abstract The current distribution of northern hemisphere perennial ice (outside glaciated mountain areas) is limited to the Greenland ice sheet and polar sea ice. During the Pleistocene glaciations, the ice mass grew and gave rise to the Fennoscandian ice sheet which covered northern Europe and parts of western Europe. Although the Pleistocene ice sheets have never reached the present-day territory of Belgium and modelling exercises indicate that the ice will never reach Belgium during the next 1 Ma, an ice sheet advance can nevertheless not be excluded, given the uncertainty of the modelling. The geological record shows that ice sheets may drastically change the surface and subsurface environment with the creation of ice-marginal valleys, subglacial tunnel valleys, glacial basins, and glacio-tectonic deformation of the foreland (i.e., the area in front of an advancing ice sheet). In this report, a summary of deformation and erosion phenomena will be given that have been observed in relation to glacial advances over formerly glaciated areas that, as far as the subsurface geometry and constitution is concerned, show strong similarities with the Campine area, i.e., a relatively flat area with sand-dominated Tertiary and Quaternary deposits overlying one or several thick clay layers (such as the Boom Clay).

Table of Contents Abstract ...... 5 1 Introduction ...... 7 2 A short history of Pleistocene glaciations ...... 7 3 Glaciotectonic deformation ...... 13 4 Glacial basins ...... 18 5 Tunnel valleys ...... 21 6 Ice-marginal valleys and response to draining of proglacial lakes ...... 25 7 Conclusions ...... 27 8 References ...... 28

1 Introduction High-level and long-lived radioactive waste is a major radiological hazard to man and environment. ONDRAF/NIRAS, the Belgian agency for radwaste management, entitled to find a solution for this problem, currently investigates the safety and feasibility of geological disposal in poorly indurated plastic clays such as the Boom Clay and Ypresian clays. The Boom Clay is a thick and relatively homogeneous clay layer that is found in the outcrop and subcrop of a large part of north-eastern Belgium. Its thickness varies between several tens (outcrop) to more than 100 m (subcrop) while its top reaches depths of up to 200 m and more in the Campine area.

One of the most important requirements for geological formations hosting a repository for radioactive waste is sufficient depth to ensure isolation of the waste from potential receptors (humans, animals, plants) and to protect it from potential detrimental processes occurring at the surface, for a very long time period, up to 1 Ma. Over such long timescales, the repository depth and the thickness of the overburden may vary significantly due to various geodynamic processes. One of these is the advance of an ice sheet into north-western Europe, or even into northern Belgium (the Campine area), as a result of future glacial conditions similar to the ones experienced during the Pleistocene. In the past, the Pleistocene Fennoscandian ice sheet reached the territory of the Netherlands (offshore and/or onshore) at least three times, resulting in erosion and/or deformation of pre-existing sediments (Zagwijn, 1974; de Gans, 2007) – see also next section. Around 150 ka, the ice sheet reached the central part of the Netherlands, ca. 100 km north of Mol, but the ice did never reach the Belgian territory.

Ice sheet modelling, using a conservative climatic evolution scenario (A-scenarios; BIOCLIM, 2001), shows that an ice sheet would not reach Belgium during a future glaciation within the next 1 Ma (Huybrechts, 2010). The model that was used for calculation the ice sheet extent is calibrated for the Late Quaternary, spanning the Eemian interglacial (from ca. 130 ka onwards), the Weichselian glacial (from ca. 100 ka onwards) and the Holocene interglacial (from ca. 11 ka onwards) using palaeoclimate data and ice-sheet reconstructions. This exercise shows that the development of ice sheets over northern Belgium is not very likely, but the small calibration period (one interglacial-glacial-interglacial cycle; i.e., 130 ka) with respect to the prediction window (1000 ka or 1 Ma), and the subsequent fact that the ice-sheet configuration for only one glacial period is taken into account for calibration should urge towards careful interpretation of the results. This, together with the observation that the Saalian ice sheet, which originated in Scandinavia, almost (100 km north of) reached the Belgian territory might suggest that an ice sheet advance during the next 1 Ma cannot be ruled out completely. Therefore, it seems necessary to compile an overview of possible deformation and erosion phenomena that might occur during an ice sheet advance, based on the inspection of the geological record in the Netherlands, Germany and Denmark.

2 A short history of Pleistocene glaciations The Quaternary Period (ca. 2.6 Ma) is characterised by an enormous periodic expansion of the cryosphere (the part of the Earth's surface that is perennially frozen). Permanent and large Northern Hemisphere ice sheets developed during the Quaternary and their growth and decay were controlled by 104-105 year scale Milankovitch cyles (Miller et al., 2005). During the earlier part of the Quaternary (from 2.6 Ma up to about 400-800 ka ago), the warm to cool alternations presented a dominant periodicity of approximately 41 ka (Figure 1). Possibly already since 800 ka, and certainly during the last 400 ka, the situation becomes more complex. In general, the glacial periods during the last 80 ka occurred with a mean periodicity of about 100 ka (ranging between 50 ka to 130 ka), and displayed a slow (70-90 ka) and uneven cooling followed by a 7 rapid deglaciation (Hays et al., 1976; Berger, 1977; Imbrie et al., 1984). Sea level changes during the Quaternary (last 2.6 Ma) are typically between + 25 m and -125 m (Miller et al., 2005). During the Elsterian, Saalian and Weichselian glaciations, a large ice sheet originating in Scandinavia and Great Britain spread across northwestern Europe (Ehlers et al., 2011; Figure 2). During the Elsterian glaciation, the ice sheet reached the northern part of the Netherlands where a dense network of tunnel valleys was formed (Figure 3; Figure 4). During the Saalian glaciation, the ice sheet reached the Hoge Veluwe in the Netherlands, which is about 100 km north of the Belgian-Dutch border (Zagwijn, 1974; Lambeck et al., 1998). The rivers Meuse and Rhine were forced in a western course parallel to the southern limit of the ice sheet at that time. Glaciotectonic deformation resulted in the formation of ice-pushed ridges in the Netherlands (100 km north of the Campine area), up to 100 m high, and glacial basins, up to 150 m deep (Zagwijn, 1974). Examples of this glacial activity in the Netherlands are the ‘Utrechtse Heuvelrug’, the ‘Hoge Veluwe’ and the ‘Sallandse Heuvelrug’, ... (van den Berg and Beets, 1987; Berendsen, 2008). In northern Germany, similar phenomena have been recognized, such as the Itterbeck-Uelsen push moraine (van Gijssel, 1987; Kluiving, 1994; Reicherter et al., 2005).

Figure 1 ‐ Global sequences (MIS = Marine Isotope Stage; LGM = Last Glacial Maximum), sea‐level curve, marine oxygen isotope curve and northwestern European regional stages of the last 2.5 Ma (Hardenbol et al., 1998; Lisiecki and Raymo, 2005; Miller et al., 2005). Graph created using TimeScale Creater 5.0, available online https://engineering.purdue.edu/Stratigraphy/tscreator/index/index.php, consulted in December 2011. The Last Glacial Maximum (LGM, c. 20 ka) is the last ice volume culmination that took place in the Late Weichselian: about 50 × 106 km3 of ice spread over a huge territory in the Northern Hemisphere. In northern Europe, the most important feature was undoubtedly the development of the Fennoscandian ice sheet (Svendsen et al., 2004; Ehlers et al., 2011). Important to mention is that, during the Weichselian glacial stage, the extent of the Fennoscandian ice sheet was more

9 restricted than during the Saalian glacial stage: the ice sheet never extended further south than Southern Denmark and Northern Germany, i.e., some 400 km north of the Belgian-Dutch border. Like most of northern Europe, the north-eastern part of Belgium experienced dry, tundra-like surface conditions during the LGM, together with permafrost, causing periglacial soil deformation (Gullentops et al., 1981; Huijzer and Vandenberghe, 1998). Our present-day climate started developing about 11.7 ka BP, when temperatures and sea-level rose again to values comparable to the present ones. North-western Europe became reforested again, first with birch and pine, later on with a mixed deciduous forest. At the same time, soils started developing and further contributed to the stabilisation of the landscape, together with a dense vegetation cover (Hoek, 2001).

Figure 2 – Extent of the Fennoscandian ice sheet during the Elsterian (~ 450 ka), Saalian (~ 140 ka), Late Weichselian (~ 20 ka) and Younger Dryas (~ 12 ka) (Ehlers et al., 2011). The Campine area is indicated by the rectangle.

Figure 3 – Palaeogeographical maps from the Praetiglian (ca. 2.5 Ma; cold), Late Tiglian (ca. 1.8 Ma; warm), Cromerian (ca. 0.7 Ma; mixed cold and warm) and Elsterian (ca. 0.45 Ma; warm) stages, compiled by de Gans (2007). The influence of the Fennoscandinavian ice sheet in the Netherlands is clearly visible during the Elsterian glaciations with the formation of tunnel valleys and subsequent glacigenic deposition (extent of ice sheet shown in red). 11

Figure 4 – Palaeogeographical maps from the Holsteinian (ca. 400 ka; warm), Saalian (ca. 140 ka; cold), Eemian (ca. 125 ka; warm) and Weichselian (ca. 100‐10 ka; cold) stages, compiled by de Gans (2007). The influence of the Fennoscandinavian ice sheet in the Netherlands is clearly visible during the Saalian glaciations with the formation of glacial basins and subsequent glacigenic deposition (extent of ice sheet shown in red). During the Weichselian glaciation, the ice sheet did not reach the Dutch offshore area.

3 Glaciotectonic deformation Glaciotectonic deformation (formerly 'cryotectonic' deformation) is the process by which complicated and deranged features and deposits are formed at glacier borders that consist of material that has been overturned, inverted, folded and transported by the shoving action of glaciers (Gary et al., 1977). Glaciotectonic deformation has been a widespread process during the Pleistocene in glaciated areas. The resulting sediments and landforms can be readily recognised in outcrops, borehole drillings and preserved landforms (Figure 5). Typical examples of glaciotectonic features include ice-pushed ridges that are bordered by glacial basins, now filled-up. Figure 5 shows the distribution of several ridges in the Netherlands and Germany that are associated with the Saalian ice-advance (note that individual ridges are not necessarily synchronous; they may belong to different re-advances within one glacial period). An interesting observation is that the southernmost ridges appear to be distributed along the north-eastern margin of the Roer Valley Graben.

Ice-pushed ridges are usually preserved as topographical highs (Figure 5). These ridges contain subsurface material that was eroded by glacial erosion (underneath the glacier) and brought up by tectonic forces induced by the pushing glacier (Figure 6; van der Wateren, 1995; Bennett, 2001). The deformation style may either be imbricated or more thrust-like in which slabs or nappes are detached right above the décollement horizon (or surface). The width and depth of the foreland wedge that is being deformed depends on variables such as the deformability of the ice front, nature and efficiency of the sediment-ice coupling, strength and deformability of the foreland wedge and frictional characteristics of the décollement horizon. Four ridge types can be distinguished then, with increasing compressive stress (Boulton et al., 1999; in Bennett, 2001):

(1) Small (< 5 m high) push moraines, with a single crest orientated parallel to the ice margin. Deformation occurs close to the ice margin in a narrow zone, often as a consequence of seasonal ice-marginal fluctuations. These push moraines are normally referred to as seasonal, or annual push moraines. (2) Large (> 5 m high) push moraines, with a single crest orientated parallel to the ice margin, which result from a more sustained advance, usually due to a marked change in glacier mass balance. (3) Narrow, multi-crested, push moraines in which significant deformation has been transmitted horizontally for the order of 50 to 300 m beyond the glacier margin, and through a thickness of perhaps 10 to 20 m, giving an aspect ratio (depth:width) for the undeformed foreland wedge of between 1:5 and 1:20. The style of deformation may involve multiple folds, or fans of listric thrusts. (4) Wide, multi-crested, push moraines in which deformation has been transmitted in excess of 300 m beyond the glacier giving aspect ratios for the undeformed foreland wedge in excess of 1:20 and typically as much as 1:50. The style of deformation commonly involves either fans of imbricate thrusts, or superimposed sub-horizontal nappes produced by overthrusting.

Figure 5 – Ice‐pushed ridges from the Saalian glaciation, preserved as topographical highs in the landscape (dashed white line); (1): Itterbeck‐Uelsen ridge, (2): Hoge Veluwe ridge, (3): Utrecht ridge. Hatched area delimits the Rheno‐hercynian massif. Faults crosscutting the base of the Neogene are shown as black lines.

Figure 6 – The anatomy of a push moraine: (A) basic definitions and key variables and (B) geometry of imbricate push moraines compared to one built of nappes. From Van der Wateren (1995) in Bennett (2001).

Large-scale foreland deformation heavily depends on the presence and depth of a ductile horizon along which décollement occurs. Detachment horizons often involve Tertiary and Pleistocene clay layers that shallow out at depths of 100 to 200 m below the surface (Van der Wateren, 2005). Moreover, the formation of a décollement horizon, particularly in less favourable lithofacies, is strongly influenced by the hydrogeology of the foreland. A combination of local (hydro)geological factors favouring elevated pore-water pressures under loading and consequently reduced sliding friction would favour the thrusting process (Kluiving, 1994). The presence of pressurized groundwater or the occurrence of confined aquifers is particularly 15 critical (Bennett, 2001). The rheology of the foreland sediments (grain-size and facies) is also often thought to influence the style of deformation. Coarse-grained sediments favour the formation of overthrust nappes, while finer-grained, more ductile sediments favour the formation of imbricate stacks (Van der Wateren, 1995; Hart and Watts, 1997).

A relevant example of glaciotectonics is the Itterbeck-Uelsen push moraines in Germany, where fine-grained Pliocene and Miocene sands have been thrust over, a.o., coarse-grained Pleistocene sands (Kluiving, 1994). The thickness of the nappes ranges between 40-50 m (Uelsen, Jansen 1 and 2 gravel pit) and 50-75 m (Itterbeck, Warrink gravel pit; Figure 7). At some localities in the Uelsen push moraine, Oligocene heavy clay from the Rupel Formation has been thrust almost up to the surface, e.g., in the clay pit in Lemke. Pushing in the basin adjacent to the ridges took place where this clay is present in the subsurface. The location of at least the Uelsen push moraine is determined by the presence of the clay, offering perfect shearing and hydrogeological conditions (Figure 8). Ice-pushed ridges may reach altitudes up to 100 m, measuring 50 km by 12 km, as for example the eastern Hoge Veluwe in the central Netherlands (de Gans, 2007).

Figure 7 – Cross‐section of deformation phenomena in the Itterbeck push moraine as observed in extraction pits; a = green loamy fine sand; b = white micaceous fine sand; c = white coarse sand and gravel. Note the overthrusting deformation style bringing unit 'a' from a depth of 50 m below sea level to 50 m above sea level. Modified from Kluiving (1994).

Figure 8 – Isopach map of Rupelian clay in the eastern Netherlands and western Germany, and the extent of the Itterbeck‐ Uelsen push moraine. Dotted line is the state boundary, thin full lines are isopachs in meters and the grey area represents the extent of the Itterbeck‐Uelsen push moraine. The figure suggests that the presence of the clay in the subsurface at least partially controls the development of the push moraine. Modified from Kluiving (1994).

In Figure 9, a reconstruction of the development of an ice lobe and its glacio-tectonic effect is given for the Utrecht ridge. The ice lobe continues southwards through an old river valley and causes thrusting of subsurface layers in southern, western and eastern direction (southern direction not visible on E-W cross-section). Décollement is favoured by a clay layer at a depth of ca. 25 m. Thrusting continues as the ice lobe grows, and thrust nappes are being created that become 'younger' towards the centre of the ice lobe (towards the left in Figure 9e).

Figure 9 – Outline of the development of the Utrecht ridge at the western margin of the Gelderse Vallei ice lobe, based on observations made in excavation pits. Figures a‐e show cross‐sections through the ice lobe. Décollement is favoured by a clay layer at a depth of 25 m (hatched layer). From van der Wateren (1985).

4 Glacial basins The distribution of ice-pushed ridges always seems to be associated with glacial basins (Figure 10). The existence of glacial basins can be explained by glaciotectonic processes, i.e. the removal of sediment/bedrock underneath an ice sheet when it becomes frozen and attached to this ice sheet (so-called sediment-ice coupling). In the previous section, we explained that a suitable ductile layer and hydrogeological conditions favour glacial foreland deformation with the creation of glacial basins and ice-pushed ridges. In Figure 11, the isohyps map of the top of the clayey Rupel Formation in the Netherlands is given together with the distribution and characteristics of glacial basins. The map may hint at the correlation between the depth of glacial basins and the depth of the Rupel Formation. Small and undeep basins (< 50 m) occur where the 18 clay is less than 200 m deep (below sea level) whereas basins up to 150 m deep occur where the clay is much deeper. Typically, such glacial basins become completely filled up by fluvio- lacustrine deposits after deglaciation, such that they are no longer visible in the landscape (Figure 12).

Figure 10 – Subglacial basins adjacent to ice‐pushed ridges in relatively coarse‐grained material. Vertical scale and basin size exaggerated. After van der Wateren (1995, 2005) in Huybrechts (2010).

Figure 11 – Depth of the top of Rupelian clay in the Netherlands (meters below sea level), and distribution of glacial basins (dark grey: continuous and up to 150 m deep; light grey: discontinuous small basins less than 50 m deep) and tunnel valleys (red lines). The 200 m depth contour line is accentuated by the hatched dark blue line. Isohyps map from Rijks Geologische Dienst (1996), distribution of tunnel valleys and glacial basins from Zagwijn (1975).

Figure 12 – Lithology and chronostratigraphy of the infill of the Amsterdam glacial basin (de Gans, 2007).

5 Tunnel valleys Tunnel valleys are large, elongated and irregular depressions cut beneath the margin of former ice sheets. They are widespread features in formerly glaciated areas. Some of them might still be visible in the morphology (open tunnel valleys; Figure 13), while others are filled and buried by younger sediment (buried tunnel valleys; Figure 14). In western and central Europe, tunnel valleys are found onshore in the Netherlands, Germany, Denmark and Poland, and also offshore in the North Sea.

Figure 13 – Open tunnel valley in Jutland (Denmark), ending at the Main Stationary Line of the Weichselian ice sheet. Coordinate system in meters. From Jorgensen and Sandersen (2006).

Quite a number of detailed investigations have been conducted on the dimensions and origin of tunnel valleys in Denmark (Figure 13 and Figure 14). Remarkably, the characteristics of open and buried tunnels are similar (Jorgensen and Sandersen, 2006). They are typically 0.5-1.5 km wide, but exceptionally can be up to 4 km wide, while the depth usually ranges between 20 m and 200 m with a maximum of 350 m. Tunnel valleys are usually U-shaped and show an anastomosing or parallel, rectilinear to slightly sinuous pattern. The internal structure is of the 'cut-and-fill' type with an overall average for tunnel valleys in Denmark of 50% infill by meltwater sands and gravels, 1/3rd by till and 1/6th by glaciolacustrine clay and silt. There are at least four generations (Weichselian, Saalian, Elsterian, pre-Elsterian) and they are observed in all kinds of substratum, ranging from sand to clay, till, limestone and chalk.

Figure 14 – Seismic sections of relatively wide buried tunnel valleys in Denmark and their subsequent interpretation with the buried valleys indicated in grey (from Jorgensen and Sandersen, 2006). Weichselian tunnel valleys in northern Denmark are interpreted to have been formed by subglacial meltwater erosion beneath the outermost part of the ice sheet during temporary standstills and minor re-advances in the overall Late Weichselian recession of the Fennoscandinavian ice sheet (Sandersen et al., 2009). All tunnel valleys were formed within a 23 time interval of a few thousand years, leaving only a few hundred years or less for the formation of the valleys at each of the 7 ice-marginal positions (see Figure 15B).

Figure 15 – Ice‐margin positions deduced from buried valleys and the topography, northernmost part of Denmark. Topography is shown in grey scale (darkest colour at + 136m). A: NW‐SE trending ice margins from an older glaciation than those in (B). B: N‐S trending ice‐margins related to the main advance in the Weichselian and subsequent re‐advances. From Sandersen et al. (2009).

There appears to be a strong correlation between substratum type and tunnel valley formation in onshore Denmark (Sandersen and Jorgensen, 2012). Overall, the Danish substratum can be subdivided into 4 zones (Figure 16): (1) Quaternary clays, (2) Cretaceous and Palaeocene chalk and limestone, (3) Palaeogene and lower Neogene clays, and (4) Miocene clays and sands. As expected, the highest density of buried valleys is found where thick and relatively shallow (within ca. first 100 m of stratigraphical column) successions of Palaeogene, Neogene or Quaternary clays were present at the time of their formation. In areas dominated by Miocene sand, Cretaceous chalk or Danian limestone, the number of tunnel valleys is limited. These observations can clearly be linked with the permeability of the substratum. The high subglacial water pressures caused by the inability of the substrata to drain subglacial meltwater favoured channelized erosion and tunnel-valley formation. Once created and filled, the tunnel valleys caused a change in the hydraulic properties of the substratum if filled with coarser sediment.

During a subsequent glaciation, they would act as hydraulic conduits for subglacial meltwater. Once the pressure is high enough, such tunnel valleys would then become reactivated.

Figure 16 – Distribution of tunnel valleys (red lines) and subcrop lithology; (1): Quaternary clays, (2): Cretaceous and Palaeogene chalk and limestone, (3): Palaeogene and lower Neogene clays, (4): Miocene clays and sands. From Sandersen and Jorgensen (2012).

6 Ice-marginal valleys and response to draining of proglacial lakes During the Saalian ice advance over the Netherlands, the rivers Rhine and Meuse were pushed to the south, forming an ice-marginal valley directly south of the ice sheet that was draining towards a large proglacial lake (Figure 17). During disintegration and retreat of the ice sheet, the Rhine reoccupied its northern course (Figure 18) while both rivers (Rhine and Meuse) started to incise as a result of the catastrophic draining of the Saalian proglacial lake through the English Channel (see also Toucanne et al., 2009). Draining of the proglacial lake caused significant lowering of the base level which resulted in fluvial incision. The deposits that were filling the

25 valleys are part of the Kreftenheye Formation (Busschers and Weerts, 2003), and reach thicknesses up to 25 m in the valleys. This value is considered to reflect the amount of incision prior to sedimentation. Note that this formation has been reported to have thicknesses up to 100 m in glacial basins that are re-occupied by rivers after glacial retreat.

Figure 17 – Position of the Rhine and Meuse during the maximum ice advance of the Saalian glaciation. Modified by van Balen and Busschers (2010) after Busschers et al. (2008).

Figure 18 – Position of the Rhine and Meuse during ice sheet retreat following maximum glaciation. Modified by van Balen and Busschers (2010) after Busschers et al. (2008). Glacial basins and ice‐pushed ridges are indicated (see Figure 17 for legend).

7 Conclusions  Clay layers play an important role in the formation and distribution of glacial deformation and erosion phenomena such as ice-pushed ridges, glacial basins and tunnel valleys.  Glacial basins in the eastern Netherlands seem to be more shallow if the clay substrate is closer to the surface; while the maximum depth of glacial basins where the first clay layer is much deeper but never exceeds 150 m.  Dense distributions of subglacial tunnel valleys are expected to occur where impermeable clay layers are within the first ca. 100-150 m of the stratigraphical column; in sandy substrate they are much less common.  Tunnel valleys might be as deep as ca. 350 m and as wide as ca. 4 km, as can be observed in Denmark; typical values are 20-200 m depth and 500-1500 m width.  The dimensions of ice-marginal valleys are limited in depth, up to ca. 25 m as can be observed from the geological record in the Netherlands.

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