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Glacial erosion by the Trift (Switzerland): Deciphering the development of riegels, rock basins and gorges

Author(s): Steinemann, Olivia; Ivy-Ochs, Susan; Hippe, Kristina; Christl, Marcus; Haghipour, Negar; Synal, Hans- Arno

Publication Date: 2021-02-15

Permanent Link: https://doi.org/10.3929/ethz-b-000458333

Originally published in: Geomorphology 375, http://doi.org/10.1016/j.geomorph.2020.107533

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ETH Library Geomorphology 375 (2021) 107533

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Geomorphology

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Glacial erosion by the Trift glacier (Switzerland): Deciphering the development of riegels, rock basins and gorges

Olivia Steinemann a,⁎, Susan Ivy-Ochs a, Kristina Hippe a, Marcus Christl a, Negar Haghipour b, Hans-Arno Synal a a Laboratory of Ion Beam Physics, ETH Zürich, Otto-Stern-Weg 5, 8093 Zürich, Switzerland b Institute of Geology, ETH Zürich, Sonneggstrasse 5, 8092 Zürich, Switzerland article info abstract

Article history: A long-lasting question in glacial geology is how and how fast were able to shape the distinctive land- Received 10 June 2020 scapes of the Alps. This study contributes to the understanding on the formation of overdeepened basins, espe- Received in revised form 18 November 2020 cially the processes and the amount of time involved. We examine the remarkably high (150 m) cross- Accepted 23 November 2020 bedrock riegel and the associated located in front of the Trift glacier in the central Swiss Alps. A Available online 26 November 2020 combined approach of field survey with measurements of two cosmogenic nuclides, 10Be and in-situ 14C, and a

Keywords: numerical model was used to determine the spatial glacial erosion patterns on the bedrock riegel. Ten samples fl Cosmogenic 10Be were taken along two transects; one perpendicular to the glacier ow direction, from outside of the Little Ice In-situ 14C Age (LIA) extent down to the centre of the riegel, and the other following the former ice-flow direction across Glacial erosion rates the riegel. Analysis of measured nuclide concentrations shows that the sample outside of the LIA was constantly Swiss Alps exposed since the retreat of the Egesen Trift glacier (~11.5 ka). The samples inside the LIA extent indicate a Overdeepening distinct trend of increasing glacial erosion rates from 0 mm/a near the LIA ice margin to high erosion all across the Inner gorge top of the riegel. The resulting minimum glacial erosion rates from samples on the riegel are 0.5–1.1 mm/a (10Be) and 0.6–>1.8 mm/a (in-situ 14C) which correspond to minimum erosion depths of 1.6–>3 m (10Be) and 1–>5 m (14C). The extremely low nuclide concentrations measured at the riegel highlight the substantial erosion (predominantly abrasion) of the bedrock surface during late Holocene glacier coverage. Field observations suggest that the formation of the overdeepening and, as a consequence, the riegel is due to a combination of valley shape, bedrock structures, glacier confluence and hydrology. We further hypothesise that the gorge is a key factor responsible for this impressive overdeepening, by lowering the threshold for the subglacial meltwater, effectively decoupling the height of the riegel from the depth of the overdeepening. © 2020 The Authors. Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).

1. Introduction main trunk valleys, in tributary valleys and in , overdeepend rock basins form because glaciers have the ability to erode their beds Glaciers and rivers shape valleys in completely different ways. This is below the fluvial profile (Penck, 1905). This leads to the distinctive gla- evident not only in the cross-profile, U- vs. V-shape, but in the longitu- cial longitudinal valley profiles which were already described by McGee dinal profile. While rivers tend towards a graded concave up profile, gla- in 1894 as: “… irregularly terraced – i.e., made up of a series of rude ciers uniquely and characteristically erode ‘unevenly’ into bedrock steps in variable form and dimension, - and some of the terraces are leaving cross-valley bedrock bars or riegels and rock basins in their so deeply excavated as to form rock-basins occupied by lakelets…” wake (Sugden and John, 1976; Evans, 2008, 2013). Glacially shaped (McGee, 1894). The rock basins, which are often tens to hundreds of are closed depressions found in the Alpine forelands meters deep, are usually followed downstream by a steep upward (Anselmetti et al., 2010; Brückl et al., 2010; Dehnert et al., 2012; Dürst (adverse) slope, which can transform into a cross-valley riegel if intense Stucki and Schlunegger, 2013; Buechi et al., 2018; Burschil et al., erosion has also occurred on the lee side of the bedrock ridge. Adverse 2019) where abundant drill core data (up to hundreds of meters slopes are often 10–30° steep, but even steeper adverse slopes have deep) and/or reflection seismic data have enabled an increased under- been reported (Röthlisberger, 1968; Hooke, 1991; Alley et al., 1997; standing. But overdeepenings are also frequent landforms of high Al- Creyts et al., 2013; Haeberli et al., 2016). In some cases, a narrow inci- pine areas (Pfiffner et al., 1997; Frey et al., 2010; Preusser et al., 2010; sion is eroded in the bedrock step or riegel, that can deepen with time Reitner et al., 2010; Linsbauer et al., 2012; Haeberli et al., 2016). In the to form a gorge across the riegel, also called an inner gorge. The origin and evolution of inner gorges is enigmatic (Montgomery and Korup, ⁎ Corresponding author. 2011; Dürst Stucki et al., 2012). Although several authors have attrib- E-mail address: [email protected] (O. Steinemann). uted their formation to predominantly fluvial processes during

https://doi.org/10.1016/j.geomorph.2020.107533 0169-555X/© 2020 The Authors. Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/). O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 periods (Montgomery and Korup, 2011; Ziegler et al., 2013; abraded bedrock, of how much rock was removed by the glacier (Briner Leith et al., 2014) pressurised, sediment-charged, subglacial meltwater and Swanson, 1998; Fabel and Harbor, 1999; Goehring et al., 2011; unequivocally has the ability to cut deep gorges into bedrock (Creyts Wirsig et al., 2016, 2017; Young et al., 2016; Steinemann et al., 2020). et al., 2013; Jansen et al., 2014; Beaud et al., 2016; Werder, 2016; Cosmogenic nuclides are produced in the upper few meters of bedrock Blomdin and Harbor, 2017). surfaces as long as they are exposed to cosmic rays, whereas during gla- In the last decade, investigation of overdeepenings has gained new cial coverage the bedrock is completely shielded and no nuclides are attention, as improved understanding of their formation (Cook and produced (Dunai, 2010). Measurement of cosmogenic nuclides in bed- Swift, 2012) and distribution is sought (Patton et al., 2015); especially rock surfaces is commonly used to date the time of glacier retreat, in the context of investigation of suitable sites for deep geological repos- thus, the time of exposure after deglaciation. However, if a glacier did itories for nuclear waste (Fischer and Haeberli, 2012). The occurrence of not remove enough rock (at least 2–3 m) to re-zero the nuclide concen- glacially scoured rock basins is often linked to areas of glacier conflu- tration built up during previous exposures, those inherited nuclides re- ence (Anderson et al., 2006; MacGregor et al., 2009) and past equilib- sult in ‘too old’ exposure ages (Fabel et al., 2004). Fortuitously, this rium line altitude (ELA) positions (Hooke, 1991; Brocklehurst and excess of cosmogenic nuclides can be used to determine erosion depths Whipple, 2004), underlining the importance of ice thickness and ice- in the bedrock below a glacier, and if the time periods of glacial coverage flow velocity (Hooke, 1991; Alley et al., 1997). But overdeepenings are are known or can be inferred from independent data, glacial erosion also found at (past) terminus positions, including cirques, where forma- rates can be derived (Briner and Swanson, 1998). tion is attributed to upward flow towards the ablating ice surface (Alley The focus of this study is the polished bedrock inside the footprint of et al., 2003; Cook and Swift, 2012; Evans, 2013; Haeberli et al., 2016). the Little (LIA) extent of the Trift glacier (Fig. 1), especially the Several authors have emphasized the control of tectonic structures conspicuous cross-valley bedrock riegel. The methodological approach or weak lithologies, promoting intensified , on the location taken is a combination of geomorphological and geological observa- of overdeepenings (Sugden and John, 1976; Augustinus, 1995; tions, cosmogenic nuclide analysis and a numerical model to determine Dühnforth et al., 2010; Hooyer et al., 2012; Becker et al., 2014). Despite glacial erosion depths and rates (Wirsig et al., 2017; Steinemann et al., extensive research carried out in the last years, questions about the 2020). Measuring two nuclides, in this study 10Be and in-situ 14C, can temporal framework remain unanswered. Crucially, constraints on provide additional information about the duration of glacier coverage, how much a glacier can reasonably erode during a glacial period or dur- taking advantage of their different half-lives and depth profiles (Miller ing a single glaciation are rare (Preusser et al., 2010). et al., 2006; Hippe, 2017). We determine erosion rates along and across Recent studies have shown that cosmogenic nuclide concentrations the riegel to better understand the spatial erosion patterns, to learn measured on glacially polished surfaces allow quantification, directly on more about the development of the rock basin, to find out how long it

Fig. 1. Overview map of the study area. The simplified geological map is overlain on a multidirectional hillshade map (modified, reproduced with the authorisation of swisstopo (JA100120)). Glacier extent of the (1850) is based on Maisch et al. (2000) and Paul et al. (2008). Lateglacial extents are according to maps of Hantke (1980). Black rectangles show extent of Figs. 3,10 and 12. Inset map shows location within Switzerland. Abbreviations: (B) Location of the city of Bern, (T) Tsanfleuron glacier and (R) Rhône gla- cier. Coordinates in black are the metric swiss grid (CH1903/LV03), and in grey the World Geodetic System (WGS84).

2 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 might have taken the glacier to carve it, as well as to consider why the 2. Study site riegel persisted even through multiple glaciations. This allows us to ex- amine the following key questions: why do overdeepenings form The focus of this study is the cross-valley bedrock riegel in the where they form? and why are some overdeepenings apparently ‘so forefield of the Trift glacier (CH1903/LV03: 670170, 171860, WGS 84: deep’?, clearly surpassing limits that should be imposed by negative 46.694°N, 8.356°E). The area is located in a tributary valley of the feedbacks (Cook and Swift, 2012). Our results place at the forefront Gadmer Valley (Gadmertal), central Swiss Alps (Fig. 1). The bedrock the critical role of inner gorges that cut across the riegels. Results of riegel is approximately 500 m long, 300 m wide and ~150 m high this study will increase knowledge about the development and evolu- (~1750 m a.s.l.) measured from the floor of the overdeepening tion of longitudinal valley profiles of glaciated areas and help us to un- (Fig. 2). The average slope, from the bottom of the lake to the top of derstand where and how fast overdeepenings could form in the case of riegel, is nearly vertical (50–70°) (Fig. 2a), while the slope on the future glaciations. downvalley side of the riegel is 40–60° (Fig. 2b). The dominating

Fig. 2. Photographs of the cross-valley Trift riegel. (a) View in glacier flow direction (NNW) (Photo: Maxi Castrillejo Iridoy, 2019). (b) View up-glacier (SE). Note in the front the river that flows through the gorge and the suspension bridge highlighted as a brown line. (c) View from the riegel towards the Trift glacier (2018). (d) Profile across the overdeepening, note the remarkably high Trift riegel. Bedrock surface is a combination of the SwissALTI3D data, below the lake is based on bathymetry (Grischott et al., 2010) and below the glacier (dashed black line) is interpreted based on radar data of the overdeepening (Dalban Canassy, 2012). Dark blue dashed line across the riegel shows the drainage of the lake through the gorge. The ice surface height during the (LGM), shown in light blue, is based on Bini et al. (2009). LIA glacier is drawn in medium blue. The terminal position of the LIA is based on field observations, the elevation above the riegel on the elevation of the LIA . These two fix points are marked with a black dot. The LIA surface was drawn in comparison with the historic map (Dufour et al., 1864). Present-day glacier (2016) is shown in blue. Location and view direction of photographs and the location of the profile are shown in Fig. 10.

3 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 lithology of the riegel and its surroundings, is a banded, biotite-rich bedrock riegel became ice-free between 1948 and 1960. Rapid retreat gneiss (locally called Erstfeld gneiss), which is part of the pre-Variscan began again around 1996. In 2016/2017, the glacier retreated almost basement of the Aar Massif (Swisstopo, 2005). In higher areas around over the next bedrock step (Fig. 2d). Today, the Trift glacier has an the Diecherhorn (3387 m a.s.l.), granites and around Hinter Tierberg area of 14.9 km2 (GLAMOS, 2018) and is completely detached from a (3444 m a.s.l.) (Fig. 1) amphibolites crop out (Menkveld, 1995). Major partially debris-covered, remnant ice body located 300 m from the pres- fault structures follow the WSW-ENE trend of the geological contacts, ent glacier front (Fig. 2c). smaller faults show a SSW-NNE orientation (Pfiffner and Ramsay, Around the year 2000, Triftsee (Lake Trift) developed. The lake is 2011; Labhart et al., 2015). dammed by the riegel and lies in a ~0.9 km wide, ~1.5 km long rock Trift glacier is a confluence glacier composed of five catchments ex- basin (Fig. 2a, c, d). The lake has a surface area of 0.3 km2 and is, based tending between Steinhüshoren (3119 m a.s.l.), Diechterhorn (3387 m on bathymetric measurements (Grischott et al., 2010), around 40–50 a.s.l.) and Tieralplistock (3383 m a.s.l.) as the southwestern boundary m deep (1603 m a.s.l., Fig. 3), depending on the meltwater inflow. La- and from the Wysse Nollen (3398 m a.s.l.) to Hinter Tierberg (3444 m custrine sediment thickness of 2–6 m is suggested based on the geo- a.s.l.) as the northeastern boundary (Fig. 1). During the Last Glacial Max- physical data (Grischott et al., 2010). The lake level (1650 m a.s.l.) is imum (LGM), the Trift glacier flowed into the Gadmertal (Gadmer Val- controlled by the ~100 m-deep gorge that cuts through the cross- ley) as a tributary glacier of the Aare glacier. The latter extended into the valley riegel (Fig. 2a). Lake formation in front of the Trift glacier was foreland (~80 km) to join the Rhône glacier at Bern (Schlüchter, 1988; closely monitored due to the potential hazard of a sudden collapse of Bini et al., 2009). The LGM ice thickness is thought to have been on the glacier tongue and a subsequent glacial outburst flood (Dalban the order of 500–700 m in the area of the riegel (Fig. 2d), and only the Canassy et al., 2011; Geopraevent, 2019). highest peaks were reaching out of the ice as nunataks (Bini et al., 2009). During the first Lateglacial climate deterioration which triggered 3. Methods Gschnitz stadial glacier advances (17–16 ka (Ivy-Ochs, 2015)), Trift gla- cier likely extended down into and coalesced with the glacier in In this study, we combine the analysis of two cosmogenic nuclides Gadmertal. During the subsequent Egesen stadial re-advance (12.9– (10Be, in-situ 14C) with detailed geomorphological observations to de- 11.7 ka (Ivy-Ochs, 2015)), three right lateral moraine ridges near the vil- termine glacial erosion rates. The idea is to take advantage of the mark- lage of Schaftelen (1150 m a.s.l.) suggest that Trift glacier just reached edly different half-lives and the difference in the decrease of production down to Gadmertal (Fig. 1, red lines) (Kinzl, 1932; Hantke, 1980; with depth of the two nuclides. Thus, measured nuclide concentrations King, 1987). During the LIA, Trift glacier advanced all the way down to of surface samples can be used to show how much bedrock a glacier has Underi Trift (Figs. 1, 2d), at an elevation of 1340 m a.s.l. (Dufour et al., removed during recent periods of glacier coverage (Briner and 1864). From the maximal extent of the LIA (1850) onwards, photo- Swanson, 1998; Fabel and Harbor, 1999). Comparison of the two nu- graphs, topographic maps and almost continuous monitoring of the gla- clides could give further insight into the duration of the last glacier cov- cier position document its fluctuations, mainly ice surface lowering and erage. The half-life of 10Be is so long (1.4 Ma) that it will in essence not retreat of the snout (GLAMOS, 2018)(Fig. 3). The horizontal part of the decay over periods of glacier cover of only a few thousand years. In

Fig. 3. Sample location and apparent exposure age map (for location see Fig. 1). Encircled numbers show sample number and location. Two transects were sampled; one transect perpendicular to, and the other parallel to the glacier flow direction. Background map is a multidirectional hillshade map (SwissALTI3D) overlain by an aerial photo (Swissimage) (reproduced with the authorisation of swisstopo (JA100120)) combined with bathymetrical data (Grischott et al., 2010). Glacier extents of different years are shown as white and blue lines with year indicated. These indicate Trift glacier extent in the year given and are not glacier advance phases. Data taken from topographical maps (reproduced with the authorisation of swisstopo (JA100120)).

4 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 contrast, the short half-life of 14C (5700 a) means that considerable Table 2 14 decay will take place in that same amount of time. In theory, if the deter- Additional in-situ Cdata. 14 10 13 14 mined exposure age based on in-situ C is younger than the Be expo- Sample ID Sample mass CO2 yield Fraction modern δ C C/Ctotal sure age, this is an indication of (glacier) coverage (Hippe, 2017). − g μgF14C ‰ 10 14

Trift-01 3.588 66.2 0.0111 ± 0.001 −15.9 1.30 ± 0.10 − 3.1. Fieldwork and GIS analysis Trift-02 3.030 22.5 0.7779 ± 0.007 15.4 91.00 ± 0.82 Trift-03 3.493 19.7 0.2489 ± 0.016 −53.6 26.90 ± 1.71 Trift-04 3.749 43.6 0.2448 ± 0.004 −17.9 28.49 ± 0.43 Geomorphological observations such as glacial striation directions, Trift-05 3.553 24.5 0.0656 ± 0.002 −4.9 7.84 ± 0.26 glacial polish, location and sedimentology of moraine deposits were Trift-06 3.724 21.1 0.0460 ± 0.002 −3.1 5.52 ± 0.19 − documented during fieldwork. Fieldwork was supported by digital car- Trift-07 3.102 33.5 0.7117 ± 0.007 9.6 84.23 ± 0.84 Trift-08 3.012 76.4 0.0163 ± 0.001 −15.6 1.91 ± 0.11 tographic data (aerial photographs, hillshade and slope maps) using the Trift-09 4.064 22.0 0.0664 ± 0.002 −12.2 7.82 ± 0.20 application Garafa GIS Pro (Garafa, LLC, Version 3.21.1) installed on a Trift-10 3.263 50.2 0.0594 ± 0.002 −12.5 6.98 ± 0.18 tablet. In two field campaigns (2017 and 2018), a total of ten bedrock sam- ples were taken along two transects on the west side of the gorge Cosmogenic in-situ 14C extraction was performed at the extraction (Table 1, Fig. 3). The first transect runs perpendicular to the glacier line at ETH Zürich (Lupker et al., 2019). The quartz was pre-heated for flow direction across the riegel crest, from just outside of the LIA extent 2 h at 500 °C to remove atmospheric 14C from the quartz surface, (Trift-07) at an elevation of 2020 m a.s.l. down to the lowest part of the followed by injection of a CO carrier gas and heating of the sample riegel at 1710 m a.s.l. (Trift-02, Trift-03, Trift-04, Trift-05, Trift-06). The 2 for 3 h at ~1670 °C to release the cosmogenic 14C from the quartz. After- second transect was sampled parallel to the ice-flow direction (Trift-06, wards, the CO gas was purified from other compounds by passing Trift-01, Trift-08, Trift-09, Trift-10). Samples were taken from the top- 2 through a Cu-Ag filter, a chemical water trap and cryogenic temperature most centimetres of glacially polished surfaces, which are located at traps. The clean CO gas sample was measured with a gas ion source slightly elevated positions compared to the surrounding bedrock, to ex- 2 connected to the MICADAS AMS system (Ruff et al., 2007; Synal et al., clude partial shielding of the bedrock from sediment or long-lasting 2007; Wacker et al., 2013). In-situ 14C data reduction was performed snow patches. Areas with any evidence of plucking were strictly as proposed by Hippe and Lifton (2014). Performance and reproducibil- avoided. Topographic shielding was measured at each sample location ity data for samples, intercomparison materials and blanks for the new with a compass and clinometer. ArcGIS and a digital surface model ETH the extraction line are given in Lupker et al. (2019). (DEM) with a resolution of 2 m (SwissALTI3D, reproduced with the au- For both nuclides apparent exposure ages were calculated with a thorisation of swisstopo (JA100120)) allowed further analysis of MATLAB code from the online calculator (v.2.3) (Balco et al., 2008)apply- the area. ing the scaling model by Lal (1991)/Stone (2000). For spallogenic pro- duction at sea level-high latitude (SLHL) a rate of 4.01 ± 0.33 at/g/a 10 3.2. 10Be and in-situ 14C sample preparation, AMS measurement and age (Phillips et al., 2016) was used for Be and 12.20 ± 0.98 at/g/a 14 calculation (Phillips et al., 2016) for in-situ C. Muonic production rates used are according to Heisinger et al. (2002a, 2002b). Detailed sample information In the laboratory, the rock samples were first crushed and sieved to a is shown in Tables 1 and 2. Measured cosmogenic nuclide concentrations, grain size of <0.8 mm. About 140 g of crushed rock was then leached ratios and calculated apparent exposure ages are summarised for both using HCl and weak HF solutions to obtain a pure quartz mineral sepa- nuclides in Table 3. rate. Additional magnetic separation was done if required. From the pure quartz samples, aliquots of 3–5gwereseparatedforin-situ14Cex- 3.3. Behaviour of 10Be, in-situ 14C and the 10Be/14C ratio with depth traction, the remaining 8–30 g were used for 10Be analysis. To extract 10Be from the quartz separates, the chemical preparation Fig. 4a depicts the main 10Be and in-situ 14C production mechanisms follows the procedure described in Kohl and Nishiizumi (1992) and Ivy- and their decreasing rates with increasing depth down into rock, calcu- Ochs et al. (2006). Samples were spiked with 200–250 μgof9Be and dis- lated based on sample information of Trift-10 (Table 1). Muon contribu- solved in concentrated HF. Two ion exchange columns precede the final tion to the total 10Be production is only a few percent even at the rock 10 9 precipitation of Be(OH)2. Be/ Be ratios were measured at the 600 kV surface (Dunai, 2010). Conversely, production by muon interactions Tandy accelerator mass spectrometry (AMS) system at ETH Zürich contributes a significant amount to total in-situ 14C production espe- (Christl et al., 2013). Measured ratios were corrected to a full chemistry cially at greater depth (>2.5 m) (Fig. 4a). For in-situ 14C, spallogenic long-term procedural blank of (3.1 ± 1.7) × 10−15. production at the surface is about 87%, production by muon interaction

Table 1 Sample information.

Sample ID Coordinates Coordinates Elevation Thickness Topographic shielding factor

WGS 84 CH1903/LV03

LAT LONG E N m. a.s.l. cm

Trift-01 46.6943 8.3563 670,188 171,865 1740 1.5 0.957 Trift-02 46.6937 8.3514 669,813 171,795 1910 1 0.979 Trift-03 46.6937 8.3522 669,881 171,799 1870 1.5 0.968 Trift-04 46.6939 8.3536 669,982 171,813 1850 1.5 0.979 Trift-05 46.6937 8.3554 670,126 171,800 1820 1.5 0.973 Trift-06 46.6937 8.3561 670,175 171,792 1750 1.5 0.970 Trift-07 46.6920 8.3488 669,620 171,602 2020 1.5 0.980 Trift-08 46.6947 8.3562 670,186 171,901 1730 1.5 0.956 Trift-09 46.6950 8.3562 670,188 171,937 1720 1.5 0.967 Trift-10 46.6950 8.3562 670,175 171,950 1710 1 0.947

5 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Table 3 AMS measured 10Be and in-situ 14C concentrations, their ratio, and calculated apparent exposure ages.

Sample ID 10Be concentrationa, b 14C concentrationc 10Be/14C 10Be/14C 10Be apparent exposure age 14C apparent exposure age

at/g at/g Normalised (SLHL) Years Years

Trift-01 2240 ± 1230 3099 ± 3181 0.72 ± 0.84 0.67 ± 0.66 140 ± 80 50 ± 50 Trift-02 186,040 ± 6760 328,310 ± 4714 0.57 ± 0.02 0.50 ± 0.02 9950 ± 360 7880 ± 110 Trift-03wm 35,990 ± 2050 67,010 ± 5769 0.54 ± 0.06 0.32 ± 0.05 2170 ± 120 1160 ± 100 Trift-04 82,690 ± 4260 157,655 ± 3828 0.52 ± 0.03 0.46 ± 0.03 4660 ± 240 3070 ± 70 Trift-05 6970 ± 1850 18,142 ± 3193 0.38 ± 0.12 0.34 ± 0.11 400 ± 110 310 ± 50 Trift-06 3980 ± 1450 7086 ± 2976 0.56 ± 0.31 0.51 ± 0.27 240 ± 90 120 ± 50 Trift-07 214,970 ± 6610 445,791 ± 5756 0.48 ± 0.02 0.42 ± 0.01 10,710 ± 330 12,600 ± 170 Trift-08 3520 ± 1300 13,655 ± 3871 0.26 ± 0.12 0.23 ± 0.11 220 ± 80 250 ± 70 Trift-09 4980 ± 1690 13,339 ± 2732 0.37 ± 0.15 0.33 ± 0.13 310 ± 100 240 ± 50 Trift-10 16,590 ± 3590 44,127 ± 3621 0.38 ± 0.09 0.34 ± 0.08 1050 ± 230 840 ± 70 wm 10Be concentration is the weighted mean of two measurements from aliquots of the same quartz mineral separate (38,030 ± 3500 at/g and 34,930 ± 2520 at/g). a Measured against an in house standard S2007N, which is calibrated relative to the 07KNSTD standard (Nishiizumi et al., 2007). b Given errors are at the 1σ level including analytical uncertainties and the error of the subtracted long time laboratory blank 10Be/9Be = (3.1 ± 1.7) × 10−15 (n = 25). c Given errors are at the 1σ level including analytical uncertainties and the error of the subtracted laboratory blank of (3.2 ± 1.1) × 104 at/g (n = 3).

contributes 13% to the total production; at 2.5 m depth spallogenic and (Fig. 4b, solid yellow and green lines). If additionally, a glacial erosion muogenic production are 18% and 82%, respectively (Fig. 4a). This leads rate of 1 mm/a is included during the time of glacier coverage, the to a complex depth-profile of the 10Be/14C ratio (Fig. 4a). At about 4 m upper 2 or 5 m are removed, respectively. Hence, the concentration at depth, the relative proportion of 14C production by muon interactions the surface corresponds to the nuclide concentration produced at this is highest but slowly decreases again with further depth. With the depth (Fig. 4b dashed yellow and green lines). In summary, compara- 10Be production being close to zero already at 4 m, the 10Be/14C ratio in- tively high 10Be/14C ratios at the surface (yellow and green solid lines creases in consequence of the decreasing 14C concentration. If the sur- in Fig. 4b), compared to the undisturbed (no burial and no erosion, face is eroded, 10Be/14C decreases with increasing erosion depth black line) 10Be/14C ratio, indicate surface coverage. Whereas lower (Fig. 4b). In order to show the effects of exposure time, coverage by 10Be/14C ratios (yellow and green dashed lines) suggest deep erosion. the glacier and glacial erosion on nuclide concentrations and nuclide ra- tios, we use the site information of Trift-10 considering start of exposure 3.4. Numerical modelling used to determine erosion depths and rates at 10.5 ka (Fig. 4b). The nuclide concentration depth profiles were calcu- lated with a MATLAB code described in more detail in Section 3.4. Fig. 4b The MECED model (Multi-nuclide Exposure, Coverage and Erosion reveals the change of the 10Be/14C ratio applying two different coverage Depth-profile), a MATLAB code, calculates theoretical nuclide concen- durations for the last 2000 (yellow) and 5000 years (green), respec- trations and the corresponding concentration-depth profiles (Wirsig, tively. If the surface is completely shielded (glacier coverage but no gla- 2015; Wirsig et al., 2016). The code implements the same production cial erosion) no nuclides are produced, the 14C concentration starts to rates and constants listed in Section 3.2. The essential input parameters decrease immediately due to its short half-life, whereas the10Be concen- are the glacier exposure/coverage history and an estimate for snow tration barely changes. Consequently, the ratio of 10Be/14Cincreases depth and duration of snow cover (number of months) at the study

a 10 14 b 10 14 production Betotal/ Ctotal Be/ C 0 0.1 0.2 0.3 0 0.2 0.4 0.6 0.8 1 0

1

2

3

4

5

depth / m 6 10Be total 7 10Be spallation 10Be muons 8 14C total no coverage 14C spallation 2000 a coverage 9 14C muons 5000 a coverage 10 14 1 mm/a erosion Betotal/ Ctotal 10 0204060 production rate / at/g/a

Fig. 4. Behaviour of 10Be/14C ratio with depth into bedrock. (a) 10Be and 14C production rates with increasing depth for the different production mechanisms of the two nuclides. (b) 10Be/14C concentration ratio behaviour after 10,500 years (black line) of exposure. Additionally, two different scenarios are shown in yellow and green including (glacier) coverage for the last 2000 and 5000 years, both without (yellow and green solid lines) and with a glacial erosion rate of 1 mm/a (dashed yellow and green lines). Note that with increasing coverage duration the 10Be/14C ratio increases. Sample information of Trift-10 (Table 1) was used for the calculations.

6 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 site (Wirsig et al., 2017; Steinemann et al., 2020). The glacier exposure/ of the riegel. In addition to the record delivered by the historic map coverage history defines for each time increment (in this study 100 (Dufour et al., 1864), LIA moraine ridges of the Trift glacier are pre- years) if the sampled surface was covered by a glacier (no production) served at several locations. On the left valley side, several parallel mo- or if it was exposed (cosmogenic nuclide production). During the winter raine ridges can be followed from the LIA glacier front (1340 m a.s.l., months, a snow cover partially shields the sampled surface during Underi Trift (Fig. 1)) almost continuously up to the riegel. The frontal phases of exposure (no glacier coverage) which reduces the production moraines are difficult to recognise in the field but their height steadily of cosmogenic nuclides. Especially in high alpine and snow-rich areas increases upvalley where they are up to ~5 m high. Most parts of the this can significantly influence the results and should be incorporated moraines are vegetated. On the hiking trail, which mainly follows the in the calculations. The loss of cosmogenic nuclides due to radioactive moraine ridges, it is apparent that the moraine is composed of a diamictic decay is also considered, which has almost no effect on 10Be but a signif- matrix-supported sediment with clasts ranging from 0.1–1.0 m in size; icant effect on the in-situ 14C concentration. boulders larger than 2 m are rare. About 20 m below the Windegghütte With these inputs, the model executes a forward calculation starting (1880 m a.s.l. Fig. 3), four small, stacked moraine ridges were preserved at the defined time and a cosmogenic nuclide concentration of zero. For in between two bedrock highs. The blocky ridges are 0.5–1.5 m high each time interval, the model calculates a concentration-depth profile and about 100 m long. The clasts are between 5 and 50 cm in size, with considering if the location is covered by the glacier or exposed during littlematrix.Alongtheslopeabovetheriegel(1980ma.s.l.),againtwo that interval and how much snow there is during the ice-free intervals. parallel LIA moraine ridges were observed (Fig. 3 just east of sample 7 Then it reads the conditions of the next time increment, calculates its location). There they are about 1–2 m high. There are no lateral moraines concentration-depth profile, and adds it to the profile of the previous above the LIA lateral moraines. time step, and so forth until the final time step. Final model outputs are a sample specific theoretical surface nuclide concentration, a nuclide concentration vs. time plot and the final concentration-depth profile. 4.2. Cosmogenic nuclide data and apparent exposure ages To determine the erosion depth, the modelled nuclide concentration 10 14 is compared to the AMS-measured concentration of the individual sam- Measured Be and in-situ C concentrations, their ratio and ob- 10 14 ple (Fig. 5). The fundamental concept is that to allow glacial erosion, the tained apparent Be and in-situ C exposure ages assuming a simple modelled nuclide concentration has to be higher than the measured nu- exposure scenario are listed in Table 3 and plotted in Figs. 3 and 6.At clide concentration, or at least they have to match, which would be an this stage, no correction for snow coverage is included. Apparent expo- indication of no erosion. To determine quantitatively how much rock sure ages of the two nuclides cover similar age ranges from 10,710 ± 10 was removed by the glacier, the final concentration-depth profile is 330 a to 140 ± 80 a for Be and from 12,600 ± 170 a to 50 ± 50 a 14 10 used. From there the erosion depth can directly be extracted by for C(Fig. 3). Except for Trift-07 and Trift-08, Be ages are slightly 14 10 intersecting the depth-profile curve with the measured concentration older than the in-situ C ages (Fig. 6). For discussion, the Be values 14 (Fig. 5). This intersection point shows the depth that corresponds to are used; in-situ C values are given in parentheses. the present-day, measured surface concentration, or how much rock In Fig. 6, exposure ages of the samples with respect to their eleva- was removed, respectively. Dividing the depth (thickness of rock re- tion are shown. A clear trend from higher exposure ages at high ele- moved) by the total duration of glacial cover allows the calculation of vation to lower ages at low elevations is seen. For better average glacial erosion rates. visualization, the young samples are plotted with a logarithmic scale in the figure inset (Fig. 6). Trift-07, the sample at highest alti- tude and the only one located outside of the LIA ice extent (Fig. 3, 4. Results and interpretation white line), has an exposure age of 10,710 ± 330 a (12,600 ± 170 a). This suggests that the site was last glacier covered and deeply 4.1. Geomorphology of the recently exposed ice-free forefield eroded during the Egesen stadial (12.9–11.7ka)andwascontinu- ously exposed thereafter. All of the other samples are located inside The bedrock riegel was intensely glacially scoured, as its overall ap- pearance is rounded. Abundant striations (general NW-NNW direction) and glacial polish can be observed across the entire riegel. Areas of plucking were frequently observed but predominantly on the lee side 2050 10Be ages Trift-07 14C ages 0 2000 50 1950 100 Trift-02 150 1900 200 Trift-03 Trift 05 250 1850 Trift-04 1800 300 elevation / m a.s.l. Trift-05 depth / cm 06 350 1800 01 08 400 09 10 450 1750 Trift-10 Trift-10 1700 500 01000100 0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 1700 nuclide concentration / at/g x 105 0 2000 44000000 6000 8000 10000 12000 age / a Fig. 5. Erosion depth determination. Curves show the modelled 10Be (red) and in-situ 14C (blue) concentration with depth; vertical lines measured nuclide concentration. Fig. 6. Plot of sample elevation against apparent exposure age. 10Be (red squares) and 14C Intersection of the curves and the corresponding measured nuclide concentrations with (blue triangles) exposure ages including 1σ uncertainties (simple exposure, no snow the uncertainties allows determination of the amount of rock removed on the y-axis correction). Inset shows the magnification of the young samples on a logarithmic scale (maximum and minimum amount). so that the uncertainties are well visible.

7 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 of the LIA extent of ~1850 CE. Therefore, the true time of exposure is The samples of the longitudinal transect, downvalley along the around 150 a or less for every one of the samples from inside of the riegel, all have very young exposure ages and overlap with each other LIA extent. Following the perpendicular transect from Trift-07 down- within their 1σ uncertainties (Fig. 6 inset); Trift-06 at 240 ± 90 a slope, the obtained exposure ages of the samples in marginal posi- (120 ± 50 a), Trift-01 at 140 ± 80 a (50 ± 50 a), Trift-08 at 220 ± 80 a tion to the LIA extent but inside of it are, with decreasing elevation: (250 ± 70 a), Trift-09 at 310 ± 100 a (249 ± 50 a), except for Trift-10 Trift-02 at 9950 ± 360 a (7880 ± 110 a), Trift-03 at 2170 ± 120 a which has a slightly older exposure age of 1050 ± 230 a (840 ± 70 a). (1160 ± 100 a) and Trift-04 at 4660 ± 340 a (3070 ± 70 a). These The concentrations for these samples are all near the detection limit ages are all distinctly older than 150 years, the maximum amount (few 10310Be at/g, Table 3). The young exposure ages are a first indication the bedrock has been exposed since onset of retreat after the LIA. of substantial glacial erosion at those locations. This suggests that the exposure ages are not “true” but are actually “apparent” exposure ages and shows the presence of significant con- 4.3. Interpretation of 10Be/14C ratios centrations of cosmogenic nuclides inherited from previous ice-free periods. Trift-05 is located right at the riegel (about 25 m below Plotting the 10Be/14C against the 14C concentration allows a first as- Trift-04) and its age of 400 ± 110 a (310 ± 50 a) is only slightly sessment of whether the samples experienced a simple or a complex older than its true exposure age (~60 years, Fig. 3). exposure history (Fig. 7a). Note that all, except one (Trift-07), of our

a field of complex 1155 1100 5 21 exposure kkaa ofof glacierglacie rcoverage coverage

Trift-01 Trift-03 Trift-02 0.50 Trift-06 Trift-04 C

14 Trift-09 Trift-05 Trift-07

Be/ Trift-10 10 constant Trift-08 exposure line

forbidden field 0.05

b field of complex 1155 1100 5 21 exposure kkaa ofof glacierglacie rcoverage coverage

0.50 C 14 11000000 aa coveragecoverage Be/ 10 constant exposure line 2000 a Glacial erosion 0 mm/a 0.1 mm/a 3000 a forbidden 0.5 mm/a field 1.0 mm/a 5000 a 4000 a 0.05 0.1 1 10 100 14C x 103 (at/g at SLHL)

Fig. 7. Plot of the 10Be/14C ratio against the in-situ 14C concentration. Note sample concentrations are normalised to sea level high latitude (SLHL). (a) Solid black line represents the constant exposure line (no coverage). A sample with a simple exposure history would plot on this line. With increasing exposure time samples would move along the line towards higher 14C concentrations (to the right). Below and to the right of the line of constant exposure is the forbidden field, meaning that theoretically no (uneroded) sample should plot in this field. Above and to the left of the line of constant exposure is the field of complex exposure. Dashed lines in this field are showing different periods of coverage during the last several thousand years. With increasing coverage duration, the 10Be/14C ratio increases (due to radioactive decay of 14C). Grey diamonds are sample data points. (b) Plot to highlight the effect of glacial erosion (coloured lines) using the site information of Trift 10. The red squares were calculated with the MECED model, by starting exposure 10,500 years ago and including a continuous glacier coverage at the end of the 10,500-year period, calculated in 1000-year steps. The same calculation was repeated by implementing glacial erosion rates of 0.1 mm/a (green curve), 0.5 mm/a (blue curve) and 1.0 mm/a (yellow curve). Both glacier coverage and glacial erosion are included in the calculation. Glacial erosion modifies the ratio of 10Be/14C ratio significantly and alters the burial signal. Note that 1 mm/a is a commonly quoted erosion rate value (Hallet et al., 1996).

8 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 samples plot within uncertainties in the field of complex exposure 07 (horizontal line Fig. 8a). This suggests that the implemented snow (Fig. 7a) suggesting that they were covered by a glacier for a certain pe- values are realistic. riod of time, likely during the Holocene. Trift-07 plots slightly outside of A first constraint for defining the glacier history is given by the ap- the line of constant exposure, which shows that this sample did not ex- parent exposure age of Trift-07 (10,710 ± 330 a; 12,600 ± 170 a), perience any coverage after it was exposed about 10.7 ka ago (Table 3). which pinpoints the moment of first re-exposure of the bedrock after In areas where there is no or very limited glacial erosion (e.g. cold- retreat of the Egesen stadial glacier (11.5 ka). It further shows that the based glaciers or slow-moving temperate glaciers) these plots would Egesen glacier eroded deep enough to remove all nuclides from previ- provide an estimation of how long the samples were last glacier cov- ous ice-free phases. Numerous studies investigated Holocene glacier ered, assuming the coverage occurred continuously and recently. For fluctuations in the Alps (Holzhauser, 2007; Nussbaumer et al., 2011; the Trift glacier, however, the very young exposure ages at the centre Schimmelpfennig et al., 2014; Le Roy et al., 2015), these data were of the riegel indicate high erosion rates that significantly influenced used to constrain the glacier coverage intervals for the sample sites in- the 10Be/14Cratios(cf.Fig. 4b). The combined effect of coverage by the side the LIA extent. Careful comparison of these studies led to imple- glacier (complete shielding) and glacial erosion on the 10Be/14C ratio is mentation of two glacier fluctuation histories that depend on the shown in Fig. 7b. With 10,500 years of continuous exposure the data location of the bedrock sampled: (i) marginal but inside the LIA extent point plots on the black solid line (Fig. 7b, black square). If the sample (Fig. 3, Trift-02, Trift-03, Trift-04) and (ii) at the horizontal part or experienced hundreds to several thousand years of coverage but no trough of the riegel (Trift-01, Trift-05, Trift-06, Trift-08, Trift-09, Trift- erosion, it would follow the red line. However, if a glacier is eroding at 10). The samples located marginally but inside of the LIA extent 0.5 mm/a (blue line, Fig. 7b) for 3000 years (the third square) the 10Be (Fig. 8c) (Trift-02, Trift-03, Trift-04) became ice-free at 11.5 ka and and 14C data plot near the 1000-years coverage line. When a glacier is were not covered by the Trift glacier until climate deteriorated mark- eroding its bed rather rapidly, non-unique solutions can be obtained edly in the late Holocene, when the samples were repeatedly glacier for the combined 10Be-14C data. This complex effect of glacial erosion covered during the intervals 3.6–3.5 ka, 2.8–2.5 ka, 1.7–1.6 ka, 1.5–1.2 on 10Be/14C ratios especially on samples with low concentrations was ka, 0.9–0.5 ka and 0.3–0 ka. The samples at the trough (lowest elevation also pointed out in other studies (Goehring et al., 2011; Beel et al., part) of the riegel became ice-free at 10.5 ka (Fig. 8e), were covered 3.7– 2015). Another fact revealed by the modelled data in Fig. 7bisthat 3.5 ka and after advance of the Trift glacier at 2.8 ka. Thereafter, the gla- with glacial erosion rates of 1 mm/a (commonly cited value, e.g. Hallet cier covered the riegel continuously until it became ice-free again be- et al., 1996), 10Be/14C ratios can even plot in the ‘forbidden’ field (yellow tween 1948 and 1960. The concentration evolution curves in Fig. 8c, e curve). At our study site, due to the high erosion rates, we were not able show that during phases of glacier coverage, no 10Be nuclides are pro- to determine the duration of the latest glacier coverage with in-situ 14C. duced and very few decay. In stark contrast, the concentration of 14Cde- Nevertheless, by implementing the periods of glacier coverage gleaned creases noticeably during glacier coverage due to radioactive decay. The from published Alpine Holocene glacier studies in the MECED model, above calculations (Fig. 8a, c) were made in a first step considering no we are still able to quantify glacial erosion rates at Trift glacier. glacial erosion as this is the sought parameter. To aid visualization of the effect of glacial erosion on nuclide concen- 4.4. MECED model trations, hypothetical erosion rates were implemented in the code (Fig. 8e). This allows a qualitative estimation of the impact of glacial ero- 4.4.1. Definition of the input parameters for MECED analysis sion rates of different magnitude on concentrations of both 10Be and in- Thorough, prudent definition of the two main input parameters, the situ 14C. The use of the hypothetical erosion rates provides as well a glacier coverage/exposure history and snow coverage, is important for glimpse of the effects of all parameters combined on nuclide concentra- reliable results. Snow depth data measured at two nearby (<7 km) tions. In the next step, the MECED model is used to calculate depths of weather stations: Gadmen/Gschletteregg (SLFGAD; CH1903/LV03; erosion and erosion rates for each sample. 673,270, 177,465, 2060 m a.s.l.) and Guttannen/Homad (SLFGU2; CH1903/LV03; 665,100, 170,100, 2110 m a.s.l.) recorded average snow 4.4.2. Determined erosion depths and erosion rates from the MECED model depths of about 100 cm and 150 cm during 6 months of the year for Calculated erosion rates and depths are shown in Table 4 and Fig. 9. the period 1999–2019. Due to the frequent strong winds at the riegel Erosion depths determined based on 10Be vary from 0 cm (Trift-02) to (name of the hut “Windegg”, means windy corner), we consider a thin- 321 cm (Trift-01). The latter depth is close to the limit of erosion rates ner snow cover as more realistic. We assume a snow depth of 50 cm resolvable with 10Be. This is because from about 3 m downwards the during 6 months of the year for the early and middle Holocene and production depth profile (Figs. 4 and 5) is nearly vertical and model- 100 cm from 2.8 ka onward when several sources suggest that climate calculated 10Be concentrations cannot be differentiated within the became colder (Wanner et al., 2011; Solomina et al., 2015). However, given uncertainties from the AMS-measured 10Be concentration. If the to accommodate the large uncertainties inherent in snow depth in the measured concentration does not intersect the depth profile but actu- past, we show in Fig. 8a, b the effect of two extreme snow depth as- ally runs parallel to it (within the listed uncertainties), one can only ob- sumptions. For this analysis we use the data for sample Trift-07 which tain a minimum erosion depth of >3 m. For in-situ 14C, there is a similar lies outside the LIA extent (Trift-07) and has been exposed constantly depth effect but due to the higher production of in-situ 14Catdepthby since retreat of the much larger Egesen stadial Trift glacier (11.5 ka). muons, the concentration-depth profile decreases more slowly com- The red and blue bands (10Be and 14C) in Fig. 8a show the increase of pared to that of 10Be. The critical depth for in-situ 14C is approximately the cosmogenic nuclide concentration during the time of exposure, 5 m. Therefore, measurable erosion depths for 14C ‘reach deeper’ and where the upper boundary of the band considers no snow cover at all range between 0 cm (Trift-02) and >5 m (Trift-01) (Fig. 4a). This may and the lower boundary takes into account that the location was cov- be considered as an advantage of using in-situ 14C over 10Be in settings ered by 200 cm of snow for six months each year. The solid line within of rapid glacial erosion. Calculated erosion rates gained independently the bands depicts the intermediate snow cover values discussed above from the two nuclides generally agree well with each other (Fig. 9, (50 cm and 100 cm of snow for 6 months), which will be used for fur- Table 4). However, the in-situ 14C erosion rates are slightly higher ther calculation of the erosion depth. In comparison to no snow cover- than the 10Be rates due to the depth effect discussed above. For the fol- age, the intermediate snow scenario (solid line) reduces the final lowing discussion, values determined with 10Be are used and corre- nuclide concentration by about 7%, while 200 cm of snow for 6 months sponding in-situ 14C values are given in parentheses. (lower boundary) reduces it by about 20%. The modelled nuclide con- Erosion rates were not calculated for Trift-07 as it is located outside centration with the intermediate snow depth scenario matches well of the LIA extent. Based on the measured exposure age (10Be: 10710 ± the measured concentration of the herein used example sample Trift- 330 a) the site was likely never covered by a glacier during the Holocene

9 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Fig. 8. 10Be (red) and in-situ 14C (blue) concentration evolution over time. Turquoise vertical bands show phases of glacier coverage. Horizontal lines represent measured nuclide concentrations with uncertainties. (a) Concentration evolution for Trift-07, located outside of the Little Ice Age (LIA) extent and undergoing constant exposure since 11.5 ka. The envelope on the concentration growth curve shows the effect of snow coverage. The upper boundary is for no snow, the lower boundary is for 200 cm of snow during six months/a, solid line in the centre of the band represents an intermediate snow thickness of 50 cm for six months/a during the early and middle Holocene and 100 cm for six months/a in the late Holocene (2.8 ka). The centre band snow value was used for subsequent calculations. (c) Concentration evolution of Trift-02, located at marginal (inside) of the LIA extent. Exposure starts at 11.5 ka and is interrupted in the late Holocene by repeated glacier advances (see text), which leads to the step-like curves. Note that during coverage, 14C concentration decreases due to its short half-life. (e) Concentration evolution of Trift-10, located at the riegel using the intermediate snow scenario (solid lines) from (a). Additional curves (dashed, dotted and dash-dotted) show the effect of glacial erosion using hypothetical values. On the right side (b,d,f) corresponding photographs of the samples on the left are shown. (g) Standing on the riegel looking downvalley. For sample locations see Fig. 3.

10 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Table 4 erosion at the Trift-02 location. 40 m lower down the riegel (Fig. 9), 10 14 Determined glacial erosion depth and rates at Trift from Be and in-situ C the calculated erosion rate for Trift-03 is 0.5–0.6 mm/a (0.6–0.7 mm/ concentrations. a) and the erosion rate from Trift-04 which is located only about 20 m Sample ID 10Be glacial 10Be glacial 14C glacial 14C glacial erosion further down is 0.2–0.3 mm/a (0.2–0.3 mm/a). Approximately 30 m erosion rate erosion depth erosion rate depth downslope, the next sample of the perpendicular transect, Trift-05, mm/a cm mm/a cm gave an erosion rate of 0.5–0.7 mm/a (0.7–0.8 mm/a). A rate of 0.6–

Trift-01 0.7 – 1.1 220 – 321 >1.8 >500 0.8 mm/a (>1.2 mm/a) was determined for Trift-06, the lowest Trift-02 0.0 – 0.0 0 – 0 0.0 – 0.0 0 – 0 sample of the perpendicular and the first of the parallel transect. The Trift-03 0.5 – 0.6 80 – 91 0.6 – 0.7 96 – 111 remaining samples of the parallel transect do not show significant var- Trift-04 0.2 – 0.3 39 – 45 0.2 – 0.3 38 – 42 iation; Trift-08: 0.7–0.8 mm/a (0.7–1.2 mm/a), Trift-09: 0.6–0.7 mm/a Trift-05 0.5 – 0.7 163 – 199 0.7 – 0.8 195 – 251 (0.8–1.1 mm/a), only Trift-10 is, with an erosion rate of 0.3–0.4 mm/a Trift-06 0.6 – 0.8 191 – 245 >1.2 356 – >500 Trift-07 Sample outside Little Ice Age extent (0.3–0.4 mm/a), a little lower. Trift-08 0.7 – 0.8 198 – 253 0.7 – 1.2 215 – 358 Trift-09 0.6 – 0.7 177 – 225 0.8 – 1.1 230 – 331 Trift-10 0.3 – 0.4 104 – 132 0.3 – 0.4 99 – 115 5. Discussion

5.1. Observed trends of determined erosion rates and depths after it was freed from ice around 11.5 ka. For Trift-02, located ~50 m The combination of cosmogenic nuclide data with the MECED model inside of the lateral moraine marking the LIA extent, the measured con- and thereof determined glacial erosion rates (Fig. 9, Table 4)enablesus centration is only slightly lower than the modelled concentration, to analyse the spatial erosion pattern along the bedrock riegel. Two clear which takes into account the several periods of glacier coverage as de- trends are apparent. First, erosion rates increase along the perpendicu- scribed above (Trift-02, Fig. 8a). This suggests no (or very little) glacial lar transect going from the LIA left-lateral moraine down towards the

Fig. 9. 3D view of the Trift riegel towards the west (reproduced with the authorisation of swisstopo (JA100120)). Calculated erosion depths and rates (Table 4) are given for both nuclides. Minimum and maximum values are based on MECED model calculations and determined as shown in Fig. 5. Note the trend of increasing erosion rate from high to low elevations, and the slightly decreasing erosion rate downstream across the riegel. 14C-based glacial erosion rate for Trift-01 lies off the scale (too rapid). Erosion depth and rate not calculated for Trift-07 as location lies outside of LIA extent. Trift glacier extents of different years are shown.

11 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 riegel (Fig. 9). Second, erosion rates are highest right at the crest of the Wiederkehr et al., 2015)(Fig. 10). The steep rock wall on the up- riegel and remain so in samples taken parallel to the glacier flow direc- glacier side (adverse slope) of the riegel extends towards the Trifttellti tion (Fig. 9). The first trend shows that the LIA Trift glacier was not ero- in the WSW and the Drosili glacier in the ENE. This structure is well ap- sive near the ice margin (Trift-02). About 40–60 m downslope from parent on the aerial photographs (Fig. 10), suggesting control of the lo- Trift-02 and about 110–130 m below the ice surface, the late Holocene cation of the riegel and the overdeepening by a fault or fracture zone. glaciers eroded 45 cm at a rate of 0.3 mm/a (0.3 mm/a) (Trift-04), and Fault breccias have been observed in similar faults in neighbouring val- locally up to 91 cm (Trift-03) with a rate of 0.5–0.6 mm/a (0.6–0.7 leys (Wehrens et al., 2017). Becker et al. (2014) argue that closely mm/a). At the riegel, which was below approximately 230–270 m of spaced fractures perpendicular to the ice flow direction favour quarry- ice when the glacier was at its LIA size (Fig. 2d), the late Holocene gla- ing and can enhance formation of bedrock ridges (riegels) in areas ciers removed between 1 and 3 m (1–>5 m) of rock. Corresponding ero- where fractures are widely spaced. The fracture or fault zone upstream sion rates are 0.3–1.1 mm/a (0.3–>1.8 mm/a). Additionally, Fig. 9 of the Trift riegel was likely the trigger for formation of the highlights the good agreement of erosion rates determined with 10Be overdeepening there. and in-situ 14C. Notably, in-situ 14C rates tend to be a little higher. As glacial erosion is directly related to glacier velocity, formation of The discrepancy increases where depth of erosion approaches or overdeepenings can be favoured in areas where glacier velocities are exceeds 2 m as determined with 10Be and where erosion rates are highest, such as at glacier confluences or near the ELA (Penck, 1905; >0.6 mm/a (samples Trift-01, Trift-06, Trift-08, Trift-09). The values of Boulton, 1996; MacGregor et al., 2000; Anderson et al., 2006; Cook these samples are considered as minimum erosion rates. and Swift, 2012; Haeberli et al., 2016). According to the historic map Cosmogenic nuclide-based glacial erosion rates have been deter- (Dufour et al., 1864), there was no confluence at the location of Lake mined at three other recently ice-free sites in the Alps: Rhône glacier Trift, during the LIA. Nevertheless, it is likely there was a confluence dur- (Goehring et al., 2011), Goldbergkees (Wirsig et al., 2017)and ing the more extensive glaciations of the Lateglacial, with Trift glacier Tsanfleuron glacier (Steinemann et al., 2020). The first two sites are in reaching down to Gadmertal (Fig. 1) during the Gschnitz and Egesen crystalline bedrock, while the latter site is in limestone. In comparison (Hantke, 1980). The Drosili glacier from the NE and a glacier to the relatively rapid glacial erosion observed at Trift and at from the Trifttellti to the W must have joined the Trift glacier at the lo- Goldbergkees, at the Tsanfleuron limestone bedrock site erosion rates cation of Lake Trift (Fig. 10) during the Lateglacial, creating a confluence ranging from 0 to 0.08 mm/a were measured. These exceptionally low situation just upstream of the riegel. rates are attributed to the presence of a well-developed karst system be- This leads to the question: Is it possible that the Lateglacial and the neath the Tsanfleuron glacier. Loss of all meltwater from beneath the late Holocene glaciers alone carved the 150 m deep rock basin? With glacier down into the karst system greatly reduces sliding and thus ero- the Gschnitz glaciers lasting from ~17–16 ka and the Egesen glaciers sion (Steinemann et al., 2020). Notably, and in comparison to the Trift lasting from 12.9–11.7 ka, a total duration of glacier coverage of roughly study, neither an overdeepening nor a riegel are present at Tsanfleuron. 2.4 ka or slightly more can be assumed. For the overdeepening to have In contrast, complete removal of the cosmogenic nuclide signature due formed in this time period alone, the estimated erosion rate would to deep glacial erosion agrees well with the findings of Wirsig et al. have to have exceeded 50 mm/a, which seems extreme. Including con- (2017). They examined late Holocene erosion depths on a (less pro- tribution of the late Holocene glacier to the carving for another ~3000 nounced) orthogneiss-bedrock riegel at Goldbergkees in the Eastern years, would still require erosion rates of >25 mm/a. Thus, it is highly Alps. Interestingly they found highest erosion depths of >3 m (>5 unlikely that the Trift riegel and overdeepening were sculpted solely mm/a) (also exceeding the limit of cosmogenic nuclides) not along by post-LGM glaciers. the riegel itself but on the adverse slope and just downvalley of the Interesting is the role of the gorge that crosses the Trift riegel and riegel. On top of the riegel the erosion depth ranged between 1 and 2 presently controls the lake level. But how can the presence of a gorge in- m. The pattern of erosion displayed by the Wirsig et al. (2017) data sug- fluence the depth of an overdeepening? In Fig. 11, we summarize the ef- gests that the Goldbergkees riegel is in the process of emerging or grow- fect of the discussed factors and show schematically how the presence ing. Our rates at Trift, on the other hand, are the highest right on the of a gorge can change the trajectory of rock basin and riegel evolution. riegel. It seems the Trift riegel did not persist or even increase in height The depth of an overdeepened basin is controlled by the potential of over multiple glaciations because of limited glacial erosion on its top. the pressurised subglacial water to ascend the adverse slope of subgla- The Trift glacier eroded the riegel at a rather rapid rate (1.1 mm/a cial bedrock obstacles (Evans, 2008; Cook and Swift, 2012). This assures (10Be), >1.8 mm/a (14C)) and yet the riegel still persists. More impor- that the produced subglacial sediment is constantly evacuated and that tantly, erosion in the rock basin upstream of the riegel must have no water accumulates below the glacier (Alley et al., 2003). If the water outpaced erosion on the riegel even though apparently the conditions is not able to escape, ponds may form below the glacier which can lead for negative feedback are met (see Section 5.2), i.e. very steep adverse to a decoupling of the glacier from its bed (floatation) (Fig. 11d). If the slopes (Alley et al., 1999, 2003; Cook and Swift, 2012). In other words, glacier has no direct contact to the bed, it loses the ability to erode it the riegel is too high and the overdeepening is too deep in comparison (Glasser and Bennett, 2004). The same happens if the sediment is not to the present understanding of overdeepenings. removed from the glacier bed by flowing subglacial meltwater; till starts to accumulate. If the till layer reaches a certain thickness it protects the 5.2. Controls on the formation and evolution of the rock basin, riegel and bedrock from erosion (Alley et al., 2003). Subglacial water flow can be gorge at Trift interrupted if the water is no longer pressurised (Fig. 11d, e) leading to supercooling and freezing of the water (Röthlisberger, 1968; Numerous factors control where overdeepenings form and how Werder, 2016). Supercooling of the water can either be caused due to deep they become (Sugden and John, 1976; Evans, 2008, 2013; Cook lowered ice thickness which is decreasing the pressure, or by exceeding and Swift, 2012). Geological structures are clearly an important factor a ratio of ~1.2–1.7 between the adverse bedrock slope and the angle of (Matthes, 1930, 1972). Small irregularities in the bedrock, for example the ice surface (Cook and Swift, 2012; Creyts et al., 2013; Werder, a change in lithology or the presence of a weak, fractured zone can trig- 2016). These negative feedbacks regulate the depth of overdeepenings. ger the formation of an overdeepening. Given that the glacier is temper- As long as the gorge deepens and constantly lowers the threshold for ate, rock mass strength is a first order factor impacting the efficacy of subglacial water drainage (Fig. 11f), a subsequent glacier can deepen glacial erosion (Augustinus, 1995; Dühnforth et al., 2010; Krabbendam the basin even further. We suggest that this mechanism has been oper- and Glasser, 2011). In that light, it is striking that the orientation of ating at the Trift glacier overdeepening. The actual threshold that the the riegel at Trift follows the general WSW-ENE trend of faults and subglacial water has to overcome is the elevation of the upstream lip thrusts in the area (Pfiffner and Ramsay, 2011; Labhart et al., 2015; of the gorge (1650 m a.s.l.), around 50 m above the estimated elevation

12 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Fig. 10. Overview of the situation in the Trift glacier region. Aerial photograph overlain on a multidirectional hillshade map (reproduced with the authorisation of swisstopo (JA100120)). Black line: location of the profile (Fig. 2d). Blue lines: Little Ice Age glacier extent (Dufour et al., 1864). Red lines: faults (Pfiffner and Ramsay, 2011), dashed red lines: inferred faults based on observation of the hillshade map and the aerial photograph (cf. Menkveld (1995)). White arrows show location and view direction of photographs in Fig. 2.

of the rock basin floor, and not the height of the bedrock riegel (~1750 m riegel (Bénévent, 1914; Creyts et al., 2013; Haeberli et al., 2016; Beaud a.s.l.). At Trift, the gorge is the critical element in allowing the glacier to et al., 2018), widening and deepening the notch (Fig. 11b). During continue deepening of the rock basin by preventing the negative feed- every period of glacier occupation, the gorge as well as the rock basin backs (Fig. 11f, g). upvalley of the riegel were further deepened (Fig. 11c, f). Fluvial incision The gorge at Trift is therefore not only enigmatic, it is a key feature during may have contributed as well (cf. Valla et al., 2010; for both the development of the apparently ‘too deep’ overdeepening Montgomery and Korup, 2011) to gorge deepening. As soon as the de- and the persistence of the riegel. The presence of mechanically weak veloping gorge was large enough to effectively accept the meltwater ac- bedrock (fault or fracture zone), led to enhanced erosion upstream of cumulating in the rock basin it became the control point for the adverse the present riegel initiating the overdeepening (Fig. 11a). The resistance slope. In this way the riegel and the overdeepening in the Lake Trift rock of the bedrock at the riegel to glacial erosion led not only to formation of basin became completely decoupled (Fig. 11g). The Trift riegel stands the riegel at that location but also to a constriction of the valley (Pippan, out as one of the most impressive riegels in the Alps, at ~150 m high 1965; Dürst Stucki et al., 2012; Haeberli et al., 2016). At some point a with an adverse slope of ~60° (cf. Haeberli et al., 2016) when measured notch into the riegel formed, likely controlled by local NNW-SSE ori- to the top of the riegel. In contrast, when measured from the lake bot- ented faults (Wiederkehr et al., 2015). During glaciations, the presence tom bedrock to the pour point of the gorge a value of 20° is obtained, of the constriction forced pressurised subglacial meltwater over the which does lie within the discussed range of up to ~30° for adverse

13 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

a 2700

2500 glacier 2300

2100

1900 subglacial water

elevation / m a.s.l. 1700 ? fault

1500 weak bedrock zone (granite or gneiss)

b 2700

2500

2300

2100

1900 subglacial water elevation / m a.s.l. 1700 ?

1500 ?

c 2700

2500

2300

2100

1900 subglacial water elevation / m a.s.l. 1700

1500 ? 0 1000 2000 3000 distance / m no gorge fromation gorge formation

d f

at adverse slope gorge allows freezing of meltwater (supercooling) drainage formation of subglacial lake

e g

supercooling drainage allows formation of ‘overdeep’ overdeepening subglacial lake

14 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533 slope inclination (Alley et al., 2003; Cook and Swift, 2012). We empha- Several valley-in-valley morphologies (broader, 150 m-deep size here the largely overlooked control of the presence of a gorge gorges), e.g. Stock and Spitallamm (Fig. 12), are found in granites to (noted by Haeberli et al., 2016) through a riegel for the existence of granodiorites of the upper reaches of Haslital (Hasli Valley). These alter- extra deep overdeepenings. nate with deep overdeepened valleys that are now mainly occupied by dammed lakes. Stock and Spitallamm valley-in-valleys are conspicu- ously wider (100–200 m; cf. Ziegler et al., 2013) than the narrow gorges 5.3. Alpine inner gorges and their relationships to riegels and in the tributary valleys (tens of meters wide; Gälmer, Bächli) to Haslital. overdeepenings We suggest that the valley-in-valley at Stock started out similar to the narrow gorge at Trift and was widened over several periods of glacier Inner gorges in bedrock are a widespread feature in the Alps occupation cycles to its present U-shape. Glacial striations are abundant (Montgomery and Korup, 2011). However, why they form where they along the valley walls in the Stock gorge. In the nearby hanging valleys, form, how they evolve and the temporal framework are still unclear at Gälmer and Bächli (Fig. 12), the riegels are in massive Central Aar (Preusser et al., 2010; Montgomery and Korup, 2011; Dürst Stucki Granite which affords few weak points for a gorge to form. The bedrock – et al., 2012; Haeberli et al., 2016). Is the Trift situation unusual or are lips are cut by narrow, 20 30 m deep notches that formed in SW-NE similarly high riegels cross-cut by deep gorges with surprisingly deep trending faults (Swisstopo, 2019). At the Rhône glacier, the overdeepenings present in other valleys? To take a step towards an- massive, resistant rock (Grimsel Granodiorite) allows development of swering these questions, we gathered information about inner gorges a distinct riegel with a steep adverse slope and a 50 m-deep (known locally as schlucht or lamm) or valley-in-valley morphologies overdeepening (Haeberli et al., 2016). The glacier snout is presently in in the central Swiss Alps near the Trift glacier site (Fig. 12,TableS1 the lake (~38 m deep) which is drained by several-meter-deep notch, and Supplemental material). but no gorge has developed. Some of the deepest and most famous gorges cut through limestone Our brief survey underlines the fact that several processes can lead units: Aareschlucht (160 m deep), Rosenlauischlucht (150 m), to bedrock gorge formation, with rock type and structure exerting Schwartze Lütschine (230 m), Gletscherschlucht (400 m). These gorges strong control (Sugden and John, 1976; Augustinus, 1995). The differ- are generally longer and deeper than gorges in crystalline rocks (gorge ences between gorges cut into limestone and those in crystalline rock depths given in supplemental data). The massifs they cut through are are pronounced, although in both case subglacial meltwater, likely act- several kilometres wide in comparison to the well-defined ‘narrow’ ing over many glacial cycles, plays a key role (Creyts et al., 2013). Gorge riegel as at Trift. Notably, overdeepenings associated with the latter orientation tends to be either aligned with the SW-NE trend of the fi three gorges are not within the limestone bedrock but are located up- orogen (Kühni and P ffner, 2001) or perpendicular to it as at Trift and stream in the crystalline rocks presently in the forefield or under the at the limestone gorges. In crystalline rocks, where faults and shear glaciers (data from Linsbauer et al., 2012; Werder, 2016). Control of zones provide the weak area for overdeepenings to initiate, riegels the location of limestone gorges, or slot canyons, is attributable to a form downstream. Where rock discontinuities are favourable, gorges combination of several factors. First, as recently shown by Steinemann cut through the riegels. In massive, competent and compact rock, like et al. (2020), because of loss of subglacial water into underlying karst for example at Rhône and Gälmer, gorge formation is limited. According systems, glaciers have difficulty to erode massive, thick-bedded lime- to our interpretation of the Trift situation, such sites would tend to host ‘ ’ stones. This means that limestone massifs, for example at the neither too deep overdeepenings (Gälmer 30 m deep, Rhône 50 m) nor Aareschlucht, can remain largely uneroded through several glaciations. pronounced riegels, as no gorge has formed. Second, during glacier coverage vertical scouring by subglacial meltwa- ter incises the limestone, as suggested for the Aare gorge (Dürst Stucki 6. Conclusions et al., 2012). Third, meltwater incision takes advantage of structurally controlled (paleo) karst features (e.g. Veress et al., 2019). The persis- The purpose of this study was to determine the spatial glacial ero- tence of the Aare gorge system, which evidently survived several glaci- sion pattern on the transverse bedrock ridge in front of Lake Trift, in ations, is attested by the presence of numerous ‘abandoned’ gorges that the central Swiss Alps, to better understand the development of the typ- are filled with glacial sediment (Hantke and Scheidegger, 1993). ical longitudinal profiles of glacially shaped valleys and the temporal More comparable to the Trift situation are the gorges and framework. On the cross-valley riegel, ten samples were taken along overdeepenings in crystalline rock. The situation at Fiesch (Fig. 12)isre- two transects, one perpendicular to and one parallel to the former ice- markably similar to that at Trift. Similar to Trift, the riegel is nearly flow direction. Samples were taken on glacially polished surfaces, completely cut by a more than 100 m-deep gorge. The bedrock floor plucked areas were avoided. Cosmogenic 10Be and in-situ 14C results (Werder, 2016) of the overdeepening is estimated to lie 140 m below were combined with field observations, glacier reconstruction history the top of the riegel. As at Trift, the glacier covered the riegel during information and a numerical model to quantify glacial erosion depths the LIA. Another site that bears similarities to Trift is the recently formed and rates. (2013), 100-m-deep Gaulisee (Werder, 2016) which is blocked by a Obtained erosion rates along the perpendicular transect from out- riegel. The main difference to Trift is that the riegel is broader than at side the LIA extent towards the centre of the riegel show a distinct in- Trift (800 m vs. 200 m). In addition, its breadth is aligned with and crease. Samples inside the LIA extent but in marginal locations have not against the orientation (SW-NE) of the foliation of the gneiss erosion rates ranging from zero to around 0.3 mm/a based on 10Be (Swisstopo, 2019). The Gaulisee drainage (gorge) takes advantage of (0.3 mm/a based on in-situ 14C). Determined erosion rates of the sam- these structural weaknesses. ples along the riegel vary between 0.5 mm/a and 1.1 mm/a based on

Fig. 11. Sketches showing the divergent development of an overdeepening and riegel depending on the presence or absence of a gorge. (a) Initial condition with a temperate glacier on a crystalline bedrock bed. Note that the bedrock has a weak zone, e.g. an area with a high density of fractures or faults or a tectonic breccia. (b) Glacier erodes its bedrock bed. In the weak zone erosion is higher and a rock basin starts to form. (c) Glacial erosion continues and the basin deepens with time. Because the subglacial meltwater is pressurised it still can overcome the adverse slope of the depression. At this point there are two possible lines of evolution; one with no gorge formation (d & e) and one with gorge formation (f & g). (d) Glacial erosion continues but if the overdeepening reaches a certain depth and the adverse slope becomes too steep the subglacial water cannot be evacuated (supercooling). The subglacial meltwater starts accumulating in the basin and protects the bedrock from further erosion. (e) Glacial erosion continues but not at the position of the subglacial lake, thus the overdeepening keeps its shape. (f) The pressurised subglacial meltwater cuts into the bedrock forming first a notch, then a gorge. The threshold, which the meltwater has to overcome, is constantly lowered. The glacier is able to deepen the basin further. (g) The glacier continues eroding because the gorge still allows drainage of the subglacial meltwater, this leads to the surprisingly deep overdeepening with the high riegel.

15 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Fig. 12. Overview of gorges in the area around Trift. Background map is a simplified geological map (Source: Swiss Federal Office of Topography) overlain on a hillshade map created with a 25 m resolution DEM (Source: Swiss Federal Office of Topography). Glacier outlines of 2010 and 1850 are based on data from Fischer et al. (2014), Maisch et al. (2000) and Paul (2006). Location of suggested overdeepenings beneath present day glaciers are based on data from Linsbauer et al. (2012) and Werder (2016). Gorges were mapped based on high resolution Lidar data SwissALTI3D (reproduced with the authorisation of swisstopo (JA100120)). The gorges are classified based on the presence of an overdeepening upstream into; gorges with overdeepening (red), gorges without overdeepening (yellow), and gorges with possible overdeepening (orange). Classification was done with help of the SwissAlti3D, the thickness of unconsolidated deposits map (Swisstopo, 2015b), the bedrock elevation model map (Swisstopo, 2015a), supplemented with borehole information from the “Geoportal des Kantons Bern” (Felsreliefkarte). Abbreviations of glaciers: A: Grosser Aletsch glacier, UA: Unteraar glacier, G: Gauli glacier, T: Trift glacier, R: Rhône glacier. Additional information on the gorges and overdeepenings shown is found in the supplemental material.

10Be (0.6–>1.8 mm/a based on in-situ 14C). The maximal erosion rate incising deeper into the bedrock. This inhibited supercooling of the sub- value indicates glacial erosion to a depth of at least 3 m from 10Be data glacial meltwater which would lead to the negative feedbacks causing (>5 m from in-situ 14C), thus almost complete erosion of the cosmo- the glacier to stop deepening the rock basin. Therefore, the presence genic signature. Despite the fact that glacial erosion along the crest of of the gorge seems to be the key factor in the shaping and further deep- the riegel was extreme and only minimum erosion rates could be deter- ening of the deep rock basin upvalley of the Trift riegel. Although there mined, the findings of this study highlight that a small glacier (9.3 km), are numerous deep gorges in this part of the Alps, a similar configura- was still able to erode over 3 m of bedrock at the riegel during late Ho- tion, exceptionally deep overdeepening, high riegel and deep gorge, is locene advances. mainly found at sites where, as at Trift, crystalline lithologies of varying We suggest that initiation of formation of both the Trift erodibilities alternate along a dominant structural trend that is perpen- overdeepening and the Trift riegel was controlled by the bedrock struc- dicular to the glacier flow direction. tures. The riegel follows the WSW-ENE orientation of a fault zone, Supplementary data to this article can be found online at https://doi. which parallels the trend of the main thrusts and faults in the area. org/10.1016/j.geomorph.2020.107533. The distinctive ~150 m deep overdeepening formed through a complex synergy of glacier hydrology, bedrock structures and the local narrowing of the valley. The weak zone triggered the overdeepening, Declaration of competing interest its intensity was probably amplified by glacier confluence and induced increased ice-flow velocity. The narrowing of the valley and repeated The authors declare that they have no known competing financial glacier occupation led to the gorge incision by pressurised subglacial interests or personal relationships that could have appeared to influ- water. The gorge steadily lowered the threshold for the meltwater, by ence the work reported in this paper.

16 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Acknowledgements Brückl, E., Brückl, J., Chwatal, W., Ullrich, C., 2010. Deep alpine valleys: examples of geo- physical explorations in Austria. Swiss J. Geosci. 103 (3), 329–344. https://doi.org/ 10.1007/s00015-010-0045-x. We appreciate the careful, critical reviews by Ian S. Evans and Buechi, M.W., Graf, H.R., Haldimann, P., Lowick, S.E., Anselmetti, F.S., 2018. Multiple Qua- Melaine Le Roy; their input led to marked improvement in the manu- ternary erosion and infill cycles in overdeepened basins of the northern Alpine fore- – script. We sincerely thank Ewelina Bros, Reto Grischott, and Ueli land. Swiss J. Geosci. 111 (1), 133 167. https://doi.org/10.1007/s00015-017-0289-9. Burschil, T., Tanner, D.C., Reitner, J.M., Buness, H., Gabriel, G., 2019. Unravelling the shape Steinemann for their support and fruitful discussions in the field, Sarah and stratigraphy of a glacially-overdeepened valley with reflection seismic: the Lienz Kamleitner for comments on an earlier version of this manuscript, and Basin (Austria). Swiss J. Geosci. https://doi.org/10.1007/s00015-019-00339-0. fi Stefan Strasky (Swisstopo) for map data. For assistance with data anal- Christl, M., Vockenhuber, C., Kubik, P.W., Wacker, L., Lachner, J., Al mov, V., Synal, H.A., 2013. The ETH Zurich AMS facilities: performance parameters and reference mate- ysis we thank Sascha Maxreiner. We would like to thank Maarten rials. Nucl. Instrum. Methods Phys. Res., Sect. B 294, 29–38. https://doi.org/10.1016/ Lupker for technical help with the in-situ 14C extraction line. We appre- j.nimb.2012.03.004. ciate insightful suggestions from Urs H. Fischer, Wilfried Haeberli and Cook, S.J., Swift, D.A., 2012. Subglacial basins: their origin and importance in glacial sys- – fi tems and landscapes. Earth Sci. Rev. 115, 332 372. https://doi.org/10.1016/j. Adrian P ffner. Special thanks go to the members of the Laboratory of earscirev.2012.09.009. Ion Beam Physics at the ETH Zürich, for support in the laboratory and Creyts, T.T., Clarke, G.K.C., Church, M., 2013. Evolution of subglacial overdeepenings in re- excellent AMS measurements. MeteoSwiss, the Swiss Federal Office of sponse to sediment redistribution and glaciohydraulic supercooling. J. Geophys. Res. Earth Surf. 118 (2), 423–446. https://doi.org/10.1002/jgrf.20033. Meteorology and Climatology provided weather data (snow depth). Dalban Canassy, P., 2012. On the Stability of Steep Glacier Tongues: A Combined Seismo- We especially acknowledge funding from the Swiss National Science logical and Ice Dynamical Study Performed on Triftgletscher. (PhD Thesis). ETH Zü- Foundation (SNSF projects 175794 and 156187). rich, Zürich. Dalban Canassy, P., Bauder, A., Dost, M., Fäh, R., Funk, M., Margreth, S., Müller, B., Sugiyama, S., 2011. Hazard assessment investigations due to recent changes in Data availability Triftgletscher, Bernese Alps, Switzerland. Nat. Hazards Earth Syst. Sci. 11 (8), 2149–2162. https://doi.org/10.5194/nhess-11-2149-2011. All data generated or analysed during this study are included in the Dehnert, A., Lowick, S.E., Preusser, F., Anselmetti, F.S., Drescher-Schneider, R., Graf, H.R., Heller, F., Horstmeyer, H., Kemna, H.A., Nowaczyk, N.R., Züger, A., Furrer, H., 2012. published paper. Evolution of an overdeepened trough in the northern Alpine Foreland at Niederweningen, Switzerland. Quat. Sci. Rev. 34, 127–145. https://doi.org/10.1016/j. References quascirev.2011.12.015. Dufour, G.-H., Müllhaupt, H., Koegel, H., 1864. Interlachen, Sarnen, Stanz, Topographische Karte der Schweiz 13, 1864. Eidgenössisches Topographisches Bureau, Genf. Alley, R.B., Mayewski, P.A., Sowers, T., Stuiver, M., Taylor, K.C., Clark, P.U., 1997. Holocene climatic instability: a prominent, widespread event 8200 yr ago. Geology 25 (6), Dühnforth, M., Anderson, R.S., Ward, D., Stock, G.M., 2010. Bedrock fracture control of gla- – 483–486. https://doi.org/10.1130/0091-7613(1997)025<0483:Hciapw>2.3.Co;2. cial erosion processes and rates. Geology 38 (5), 423 426. https://doi.org/10.1130/ G30576.1. Alley, R.B., Strasser, J.C., Lawson, D.E., Evenson, E.B., Larson, G.J., 1999. Glaciological and Dunai, T.J., 2010. Cosmic rays. In: Dunai, T.J. (Ed.), Cosmogenic Nuclides: Principles, Con- geological implications of basal-ice accretion in overdeepenings. Special Paper of cepts and Applications in the Earth Surface Sciences. Cambridge University Press, the Geological Society of America 337, 1–9. https://doi.org/10.1130/0-8137-2337- Cambridge, pp. 1–24 https://doi.org/10.1017/CBO9780511804519.002. X.1. Dürst Stucki, M., Schlunegger, F., 2013. Identification of erosional mechanisms during past Alley, R.B., Lawson, D.E., Larson, G.J., Evenson, E.B., Baker, G.S., 2003. Stabilizing feedbacks glaciations based on a bedrock surface model of the central European Alps. Earth in glacier-bed erosion. Nature 424 (6950), 758–760. https://doi.org/10.1038/ Planet. Sci. Lett. 384, 57–70. https://doi.org/10.1016/j.epsl.2013.10.009. nature01839. Dürst Stucki, M., Schlunegger, F., Christener, F., Otto, J.C., Götz, J., 2012. Deepening of inner fi Anderson, R.S., Molnar, P., Kessler, M.A., 2006. Features of glacial valley pro les simply ex- gorges through subglacial meltwater — an example from the UNESCO Entlebuch area, plained. J. Geophys. Res. 111 (F1), F01004. https://doi.org/10.1029/2005JF000344. Switzerland. Geomorphology 139-140, 506–517. https://doi.org/10.1016/j. Anselmetti, F., Drescher-Schneider, R., Furrer, H., Graf, H.R., Lowick, S.E., Preusser, F., Riedi, geomorph.2011.11.016. M.A., 2010. A ~180,000 years sedimentation history of a perialpine overdeepened gla- Evans, I.S., 2008. Glacial erosional processes and forms: mountain glaciation and glacier cial trough (Wehntal, N-Switzerland). Swiss J. Geosci. 103, 345–361. https://doi.org/ geography. Ch. 11. In: Burt, T.P., Chorley, R.J., Brunsden, D., Cox, N.J., Goudie, A.S. 10.1007/s00015-010-0041-1. (Eds.), The History of the Study of Landforms or the Development of Geomorphology, Augustinus, P.C., 1995. Glacial valley cross-profile development: the influence of in situ v. 4: Quaternary and Recent Processes and Forms (1890–1960s) and the Mid-Century rock stress and rock mass strength, with examples from the Southern Alps, New Revolutions. The Geological Society, London, pp. 413–494. Zealand. Geomorphology 14 (2), 87–97. https://doi.org/10.1016/0169-555X(95) Evans, I.S., 2013. Glacial landforms, erosional features: major scale forms. In: Elias, S.A. 00050-X. (Ed.), The Encyclopedia of Quaternary Science. vol. 1. Elsevier, Amsterdam, Balco, G., Stone, J.O., Lifton, N.A., Dunai, T.J., 2008. A complete and easily accessible means of pp. 847–864. calculating surface exposure ages or erosion rates from 10Be and 26Al measurements. Fabel, D., Harbor, J., 1999. The use of in-situ produced cosmogenic radionuclides in glaci- Quat. Geochronol. 3 (3), 174–195. https://doi.org/10.1016/j.quageo.2007.12.001. ology and glacial geomorphology. Ann. Glaciol. 28 (1), 103–110. https://doi.org/ Beaud, F., Flowers, G.E., Venditti, J.G., 2016. Efficacy of bedrock erosion by subglacial water 10.3189/172756499781821968. flow. Earth Surface Dynamics 4 (1), 125–145. https://doi.org/10.5194/esurf-4-125- Fabel, D., Harbor, J., Dahms, D., James, A., Elmore, D., Horn, L., Daley, K., Steele, C., 2004. 2016. Spatial patterns of glacial erosion at a valley scale derived from terrestrial cosmogenic Beaud, F., Venditti, J.G., Flowers, G.E., Koppes, M., 2018. Excavation of subglacial bedrock Be-10 and Al-26 concentrations in rock. Ann. Assoc. Am. Geogr. 94 (2), 241–255. channels by seasonal meltwater flow. Earth Surf. Process. Landf. 43 (9), 1960–1972. https://doi.org/10.1111/j.1467-8306.2004.09402001.x. https://doi.org/10.1002/esp.4367. Fischer, U.H., Haeberli, W., 2012. Glacial Overdeepening. Results of a Workshop Held in – – Becker, R.A., Tikoff, B., Riley, P.R., Iverson, N.R., 2014. Preexisting fractures and the forma- Zürich, Switzerland, 20 21 April 2012 Arbeitsbericht NAB 12-48. pp. 1 60. tion of an iconic American landscape Tuolumne Meadows, Yosemite National Park, Fischer, M., Huss, M., Barboux, C., Hoelzle, M., 2014. The new Swiss glacier inventory USA. GSA Today 24 (11), 4–10. https://doi.org/10.1130/GSATG203A.1. SGI2010: relevance of using high-resolution source data in areas dominated by very small glaciers. Arct. Antarct. Alp. Res. 46 (4), 933–945. https://doi.org/10.1657/ Beel, C.R., Goehring, B.M., Lifton, N.A., 2015. How many and from where? Assessing the 1938-4246-46.4.933. sensitivity of exposure durations calculated from paired bedrock 14C/10Be measure- ments in glacial troughs. Quat. Geochronol. 29, 1–5. Frey, H., Haeberli, W., Linsbauer, A., Huggel, C., Paul, F., 2010. A multi-level strategy for an- ticipating future glacier lake formation and associated hazard potentials. Nat. Hazards Bénévent, E., 1914. Sur les encoches du verrou glaciare. Comptes rendus hebdomadaires Earth Syst. Sci. 10 (2), 339–352. https://doi.org/10.5167/uzh-33174. des séances de l'académie des sciences 158, 742–744. Geopraevent, 2019. Gletscherueberwachung Trift. https://www.geopraevent.ch/project/ Bini, A., Buoncristiani, J., Couterrand, S., Ellwanger, D., Felber, M., Florineth, D., Graf, H., gletscherueberwachung-trift/ (Zürich). Keller, O., Kelly, M., Schlüchter, C., 2009. Die Schweiz während des letzteiszeitlichen GLAMOS, 2018. The Swiss glaciers 1880-2016/17, Glaciological Reports No 1-138 Year- Maximums (LGM), 1:500 000 (Bundesamt für Landestopografie). books of the Cryospheric Commission of the Swiss Academy of Sciences (SCNAT). Blomdin, R., Harbor, J., 2017. Glacial erosional processes and landforms. In: Richardson, N. https://doi.org/10.18752/glrep_series. C.D., Goodchild, M.F., Kobayashi, A., Liu, W., Marston, R.A. (Eds.), International Ency- Glasser, N.F., Bennett, M.R., 2004. Glacial erosional landforms: origins and significance for clopedia of Geography: People, the Earth, Environment and Technology. John Wiley palaeoglaciology. Prog. Phys. Geogr. 28 (1), 43–75. https://doi.org/10.1191/ &Sonshttps://doi.org/10.1002/9781118786352.wbieg0719. 0309133304pp401ra. Boulton, G.S., 1996. Theory of glacial erosion, transport and deposition as a consequence Goehring, B.M., Schaefer, J.M., Schlüchter, C., Lifton, N.A., Finkel, R.C., Jull, A.J.T., Akçar, N., of subglacial sediment deformation. J. Glaciol. 42 (140), 43–62. https://doi.org/ Alley, R.B., 2011. The Rhone Glacier was smaller than today for most of the Holocene. 10.3189/S0022143000030525. Geology 39 (7), 679–682. https://doi.org/10.1130/G32145.1. Briner, J., Swanson, T., 1998. Using inherited cosmogenic 36Cl to constrain glacial erosion Grischott, R., Anselmetti, F., Funk, M., 2010. Seismic Survey Lake Trift, Tech. Rep.: EAWAG rates of the Cordilleran . Geology 26 (1), 3–6. https://doi.org/10.1130/0091- and VAW. 7613(1998)0262.3.CO2. Haeberli, W., Linsbauer, A., Cochachin, A., Salazar, C., Fischer, U.H., 2016. On the morpho- Brocklehurst, S.H., Whipple, K.X., 2004. Hypsometry of glaciated landscapes. Earth Surf. logical characteristics of overdeepenings in high-mountain glacier beds. Earth Surf. Process. Landf. 29, 907–926. https://doi.org/10.1002/esp.1083. Process. Landf. 41 (13), 1980–1990. https://doi.org/10.1002/esp.3966.

17 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Hallet, B., Hunter, L., Bogen, J., 1996. Rates of erosion and sediment evacuation by glaciers: Menkveld, J.W., 1995. Der geologische Bau des Helvetikums der Innerschweiz. A review of field data and their implications. Glob. Planet. Chang. 12 (1), 213–235. (Phd Thesis). University of Bern, Bern. https://doi.org/10.1016/0921-8181(95)00021-6. Miller, G.H., Briner, J.P., Lifton, N.A., Finkel, R.C., 2006. Limited ice-sheet erosion and com- Hantke, R., 1980. Letzte Warmzeiten, Würm-Eiszeit, Eisabbau und Nacheiszeit der Alpen- plex exposure histories derived from in situ cosmogenic 10Be, 26Al, and 14ConBaffin Nordseite vom Rhein- zum Rhone-System. Ott, Thun. Island, Arctic Canada. Quat. Geochronol. 1 (1), 74–85. https://doi.org/10.1016/j. Hantke, R., Scheidegger, A.E., 1993. Zur Genese der Aareschlucht (Berner Oberland, quageo.2006.06.011. Schweiz). Geographica Helvetica 48 (3), 120–124. https://doi.org/10.5194/gh-48- Montgomery, D.R., Korup, O., 2011. Preservation of inner gorges through repeated Alpine 120-1993. glaciations. Nat. Geosci. 4, 62. https://doi.org/10.1038/ngeo1030. Heisinger, B., Lal, D., Jull, A.J.T., Kubik, P., Ivy-Ochs, S., Knie, K., Nolte, E., 2002a. Production Nishiizumi, K., Imamura, M., Caffee, M.W., Finkel, R.C., McAninch, J., 2007. Absolute cali- of selected cosmogenic radionuclides by muons: 2. Capture of negative muons. Earth bration of 10Be AMS standards. Nucl. Instrum. Methods Phys. Res., Sect. B 258 (2), Planet. Sci. Lett. 200, 357–369. https://doi.org/10.1016/S0012-821X(02)00641-6. 403–413. Heisinger, B., Lal, D., Jull, A.J.T., Kubik, P., Ivy-Ochs, S., Neumaier, S., Knie, K., Lazarev, V., Nussbaumer, S.U., Steinhilber, F., Trachsel, M., Breitenmoser, P., Beer, J., Blass, A., Grosjean, Nolte, E., 2002b. Production of selected cosmogenic radionuclides by muons: 1. Fast M., Hafner, A., Holzhauser, H., Wanner, H., Zumbühl, H.J., 2011. Alpine climate during – muons. Earth Planet. Sci. Lett. 200, 345 355. https://doi.org/10.1016/S0012-821X the Holocene: a comparison between records of glaciers, lake sediments and solar ac- (02)00640-4. tivity. J. Quat. Sci. 26, 703–713. https://doi.org/10.1002/jqs.1495. Hippe, K., 2017. Constraining processes of landscape change with combined in situ cos- Patton, H., Swift, D., Clark, C., Livingstone, J., Cook, S., Hubbard, A., 2015. Automated mapping 14 10 – mogenic C- Be analysis. Quat. Sci. Rev. 173, 1 19. https://doi.org/10.1016/j. of glacial overdeepenings beneath contemporary ice sheets: approaches and potential quascirev.2017.07.020. applications. Geomorphology, 232 https://doi.org/10.1016/j.geomorph.2015.01.003. Hippe, K., Lifton, N., 2014. Calculating isotope ratios and nuclide concentrations for in situ Paul, F., 2006. The New Swiss Glacier Inventory 2000: Application of Remote Sensing and cosmogenic 14C analyses. Radiocarbon 56 (3), 1167–1174. https://doi.org/10.2458/ GIS. (Phd Thesis). Geographisches Institut der Universität Zürich, Zürich. 56.17917. Paul, F., Machguth, H., Hoelzle, M., Salzmann, N., Haeberli, W., 2008. Alpine-wide distrib- Holzhauser, H., 2007. Holocene glacier fluctuations in the Swiss Alps. Environnements et uted modelling: a tool for assessing future glacier change? In: cultures à l’Âge du Bronze en Europe occidentale 29–43. Orlove, B.S., Wiegandt, E., Luckman, B.H. (Eds.), Darkening Peaks: Glacier Retreat, Sci- Hooke, R.L., 1991. Positive feedbacks associated with erosion of glacial cirques and ence, and Society. University of California Press, Berkeley, pp. 111–125 overdeepenings. Geol. Soc. Am. Bull. 103, 1104–1108. https://doi.org/10.1130/0016- 7606(1991)103<1104:PFAWEO>2.3.CO;2. Penck, A., 1905. Glacial features in the surface of the Alps. The Journal of Geology 13 (1), 1–19. https://doi.org/10.1086/621202. Hooyer, T.S., Cohen, D., Iverson, N.R., 2012. Control of glacial quarrying by bedrock joints. fi Geomorphology 153-154, 91–101. https://doi.org/10.1016/j.geomorph.2012.02.012. P ffner, O.A., Ramsay, J.G., 2011. Structural Map of the Helvetic Zone of the Swiss Alps, In- Ivy-Ochs, S., 2015. Glacier variations in the European Alps at the end of the last glaciation. cluding Vorarlberg (Austria) and Haute Savoie (France), 1:100,000, Erläuterungen. Cuadernos de Investigación Geográfica 41 (2), 295–315. Bundesamt für Landestopographie, Wabern. fi Ivy-Ochs, S., Kerschner, H., Reuther, A., Maisch, M., Sailer, R., Schaefer, J., Kubik, P., Synal, P ffner, O.A., Lehner, P., Heitzman, P., Felber, M., Frei, W., Pugin, A., Marchant, R., Besson, fl fi A., Schlüchter, C., 2006. The timing of glacier advances in the northern European O., Stamp i, G., 1997. Incision and back lling of Alpine valleys: Pliocene, fi Alps based on surface exposure dating with cosmogenic 10Be, 26Al, 36Cl and 21Ne. and Holocene processes. In: P ffner, O.A., Lehner, P., Heitzman, P., Mueller, S., Steck, Geol. Soc. Am. Spec. Pap. 415, 43–60. https://doi.org/10.1130/2006.2415(04). A. (Eds.), Deep Structure of the Swiss Alps-Results From NRP 20. Birkhäuser, Basel, – Jansen, J.D., Codilean, A.T., Stroeven, A.P., Fabel, D., Hättestrand, C., Kleman, J., Harbor, J., pp. 265 288. Heyman, J., Kubik, P., Xu, S., 2014. Inner gorges cut by subglacial meltwater during Phillips, F.M., Argento, D.C., Balco, G., Caffee, M.W., Clem, J., Dunai, T.J., Finkel, R., Goehring, Fennoscandian ice sheet decay. Nat. Commun. 5, 3815. https://doi.org/10.1038/ B., Gosse, J.C., Hudson, A.M., Jull, A.J.T., Kelly, M.A., Kurz, M., Lal, D., Lifton, N., Marrero, ncomms4815. S.M., Nishiizumi, K., Reedy, R.C., Schaefer, J., Stone, J.O.H., Swanson, T., Zreda, M.G., King, L., 1987. Gletscherschwankungen und Moränen: Studien zur postglazialen 2016. The CRONUS-Earth Project: a synthesis. Quat. Geochronol. 31, 119–154. Gletscher- und Vegetationsgeschichte des Sustenpassgebietes. Selbstverlag Giessener https://doi.org/10.1016/j.quageo.2015.09.006. Geographische Schriften, Giessen. Pippan, T., 1965. Morphological studies concerning glaciated areas in Norway's high Kinzl, H., 1932. Die grössten nacheiszeitlichen Gletschervorstösse in den Schweizer Alpen mountains with special reference to alpine landforms. Die Erde (Zeitschrift der Ge- und in der Mont Blanc-Gruppe (Berlin). sellschaft für Erdkunde, Berlin) 96 (2), 105–121. Kohl, C.P., Nishiizumi, K., 1992. Chemical isolation of quartz for measurement of in-situ - Preusser, F., Reitner, J., Schlüchter, C., 2010. Distribution, geometry, age and origin of produced cosmogenic nuclides. Geochim. Cosmochim. Acta 56 (9), 3583–3587. overdeepened valleys and basins in the Alps and their foreland. Swiss J. Geosci. Krabbendam, M., Glasser, N.F., 2011. Glacial erosion and bedrock properties in NW 103, 407–426. https://doi.org/10.1007/s00015-010-0044-y. Scotland: abrasion and plucking, hardness and joint spacing. Geomorphology Reitner, J., Gruber, W., Römer, A., Morawetz, R., 2010. Alpine overdeepenings and paleo- 130 (3–4), 374–383. https://doi.org/10.1016/j.geomorph.2011.04.022. ice flow changes: an integrated geophysical-sedimentological case study from Tyrol Kühni, A., Pfiffner, O., 2001. The relief of the Swiss Alps and adjacent areas and its relation (Austria). Swiss J. Geosci. 103, 385–405. https://doi.org/10.1007/s00015-010-0046-9. to lithology and structure: topographic analysis from a 250-m DEM. Geomorphology Röthlisberger, H., 1968. Erosive processes which are likely to accentuate or reduce the 41, 285–307. https://doi.org/10.1016/S0169-555X(01)00060-5. bottom relief of valley glaciers. International Association of ScientificHydrology Labhart, T.P., Gisler, C., Renner, F., Schwizer, B., Schaltegger, U., 2015. Blatt 1211 Meiental, 87–97. mit Südostteil von Blatt 1191 Engelberg. Bundesamt für Landestopografie, Wabern. Ruff, M., Wacker, L., Gäggeler, H.W., Suter, M., Synal, H.A., Szidat, S., 2007. A gas ion source Lal, D., 1991. Cosmic ray labeling of erosion surfaces: in situ nuclide production rates and for radiocarbon measurements at 200 kV. Radiocarbon 49, 307–314. https://doi.org/ erosion models. Earth Planet. Sci. Lett. 104, 424–439. https://doi.org/10.1016/0012- 10.1017/S0033822200042235. 821X(91)90220-C. Schimmelpfennig, I., Schaefer, J.M., Akçar, N., Koffman, T., Ivy-Ochs, S., Schwartz, R., Finkel, Le Roy, M., Nicolussi, K., Deline, P., Astrade, L., Edouard, J.L., Miramont, C., Arnaud, F., 2015. R.C., Zimmerman, S., Schlüchter, C., 2014. A chronology of Holocene and Little Ice Age Calendar-dated glacier variations in the western European Alps during the glacier culminations of the Steingletscher, Central Alps, Switzerland, based on high- Neoglacial: the Mer de Glace record, Mont Blanc massif. Quat. Sci. Rev. 108, 1–22. sensitivity beryllium-10 moraine dating. Earth Planet. Sci. Lett. 393, 220–230. https://doi.org/10.1016/j.quascirev.2014.10.033. https://doi.org/10.1016/j.epsl.2014.02.046. Leith, K., Moore, J.R., Amann, F., Loew, S., 2014. Subglacial extensional fracture develop- Schlüchter, C., 1988. The deglaciation of the Swiss-Alps: a paleoclimatic event with chro- ment and implications for Alpine Valley evolution. J. Geophys. Res. Earth Surf. 119 nological problems. Bulletin de l'Association Francaise pour l'Etude du Quaternaire 2 – (1), 62 81. https://doi.org/10.1002/2012JF002691. (3), 141–145. Linsbauer, A., Paul, F., Haeberli, W., 2012. Modeling glacier thickness distribution and bed Solomina, O.N., Bradley, R.S., Hodgson, D.A., Ivy-Ochs, S., Jomelli, V., Mackintosh, A.N., topography over entire mountain ranges with GlabTop: application of a fast and ro- Nesje, A., Owen, L.A., Wanner, H., Wiles, G.C., Young, N.E., 2015. Holocene glacier fluc- bust approach. J. Geophys. Res. 117, F03007. https://doi.org/10.1029/2011JF002313. tuations. Quat. Sci. Rev. 111, 9–34. https://doi.org/10.1016/j.quascirev.2014.11.018. Lupker, M., Hippe, K., Wacker, L., Steinemann, O., Tikhomirov, D., Maden, C., Haghipour, Steinemann, O., Ivy-Ochs, S., Grazioli, S., Luetscher, M., Fischer, U.H., Vockenhuber, C., N., Synal, H.-A., 2019. In-situ cosmogenic 14C analysis at ETH Zürich: Characterization Synal, H.A., 2020. Quantifying glacial erosion on a limestone bed and the relevance and performance of a new extraction system. Nucl. Instrum. Methods Phys. Res., Sect. for landscape development in the Alps. Earth Surf. Process. Landf. https://doi.org/ B457,30–36. https://doi.org/10.1016/j.nimb.2019.07.028. 10.1002/esp.4812. MacGregor, K.R., Anderson, R.S., Anderson, S.P., Waddington, E.D., 2000. Numerical simu- Stone, J.O., 2000. Air pressure and cosmogenic isotope production. J. Geophys. Res. Solid lations of glacial-valley longitudinal profile evolution. Geology 28 (11), 1031–1034. Earth 105 (B10), 23753–23759. https://doi.org/10.1029/2000JB900181. https://doi.org/10.1130/0091-7613(2000)0282.3.CO2. Sugden, D.E., John, B.S., 1976. Glaciers and Landscape: A Geomorphological Approach. Ar- MacGregor, K.R., Anderson, R.S., Waddington, E.D., 2009. Numerical modeling of glacial erosion and headwall processes in alpine valleys. Geomorphology 103, 189–204. nold, London. https://doi.org/10.1016/j.geomorph.2008.04.022. Swisstopo, 2005. Geologische Karte der Schweiz 1:500 000. Bundesamt für fi Maisch, M., Wipf, A., Denneler, B., Battaglia, J., Benz, C., 2000. Die Gletscher der Schweizer Landestopogra e, Wabern. Alpen. Gletscherhochstand 1850, Aktuelle Vergletscherung, Gletscherschwund- Swisstopo, 2015a. Höhenmodell der Felsoberfläche. Bundesamt für Landestopographie, Szenarien (Schlussbericht NFP 31). Vdf Hochschulverlag ETH Zürich, Zürich https:// Wabern. doi.org/10.18750/inventory.sgi1850.r1992. Swisstopo, 2015b. Mächtigkeit des Lockergesteins. Bundesamt für Landestopographie, Matthes, F.E., 1930. Geologic History of the Yosemite Valley. 160. U.S. Geological Survey Wabern. Professional Paper, Reston, VA https://doi.org/10.3133/pp160. Swisstopo, 2019. GeoCover Guttannen 1230, kompilation. Bundesamt für Matthes, F.E., 1972. Geologic history of the Yosemite Valley. In: Embleton, C. (Ed.), Gla- Landestopografie swisstopo, Wabern. ciers and Glacial Erosion. Springer, Palgrave, London, pp. 92–118. Synal, H.-A., Stocker, M., Suter, M., 2007. MICADAS: a new compact radiocarbon AMS sys- McGee, W.J., 1894. Glacial Canons. The Journal of Geology 2 (4), 350–364. https://doi.org/ tem. Nucl. Instrum. Methods Phys. Res., Sect. B 259 (1), 7–13. https://doi.org/ 10.1086/606974. 10.1016/j.nimb.2007.01.138.

18 O. Steinemann, S. Ivy-Ochs, K. Hippe et al. Geomorphology 375 (2021) 107533

Valla, P.G., Van Der Beek, P.A., Carcaillet, J., 2010. Dating bedrock gorge incision in the 146, Geologischer Atlas der Schweiz 1:25 000. Bundesamt für Landestopografie, French Western Alps (Ecrins-Pelvoux massif) using cosmogenic 10Be. Terra Nova Wabern. 22 (1), 18–25. https://doi.org/10.1111/j.1365-3121.2009.00911.x. Wirsig, C., 2015. Constraining the Timing of Deglaciation of the High Alps and Rates of Veress, M., Telbisz, T., Tóth, G., Lóczy, D., Ruban, D.A., Gutak, J.M., 2019. Glaciokarsts. Subglacial Erosion With Cosmogenic Nuclides. (PhD Thesis). ETH Zurich, Zürich. Springer, Cham https://doi.org/10.1007/978-3-319-97292-3. Wirsig, C., Ivy-Ochs, S., Akçar, N., Lupker, M., Hippe, K., Wacker, L., Vockenhuber, C., Wacker, L., Fahrni, S., Hajdas, I., Molnar, M., Synal, H.-A., Szidat, S., Zhang, Y., 2013. A ver- Schlüchter, C., 2016. Combined cosmogenic 10Be, in situ 14Cand36Cl concentrations satile gas interface for routine radiocarbon analysis with a gas ion source. Nucl. constrain Holocene history and erosion depth of Grueben glacier (CH). Swiss Instrum. Methods Phys. Res., Sect. B 294, 315–319. https://doi.org/10.1016/j. J. Geosci. 109 (3), 379–388. https://doi.org/10.1007/s00015-016-0227-2. nimb.2012.02.009. Wirsig, C., Ivy-Ochs, S., Reitner, J.M., Christl, M., Vockenhuber, C., Bichler, M., Reindl, M., Wanner, H., Solomina, O., Grosjean, M., Ritz, S.P., Jetel, M., 2011. Structure and origin of 2017. Subglacial abrasion rates at Goldbergkees, Hohe Tauern, Austria, determined Holocene cold events. Quat. Sci. Rev. 30, 3109–3123. https://doi.org/10.1016/j. from cosmogenic 10Be and 36Cl concentrations. Earth Surf. Process. Landf. 42 (7), quascirev.2011.07.010. 1119–1131. https://doi.org/10.1002/esp.4093. Wehrens, P., Baumberger, R., Berger, A., Herwegh, M., 2017. How is strain localized in a Young, N.E., Briner, J.P., Maurer, J., Schaefer, J.M., 2016. 10Be measurements in bedrock meta-granitoid, mid-crustal basement section? Spatial distribution of deformation constrain erosion beneath the Greenland Ice Sheet margin. Geophys. Res. Lett. 43 in the central Aar massif (Switzerland). J. Struct. Geol. 94. https://doi.org/10.1016/j. (22), 11,708–711,719. https://doi.org/10.1002/2016GL070258. jsg.2016.11.004. Ziegler, M., Loew, S., Moore, J.R., 2013. Distribution and inferred age of exfoliation joints in Werder, M.A., 2016. The hydrology of subglacial overdeepenings: a new supercooling the Aar Granite of the central Swiss Alps and relationship to Quaternary landscape evo- threshold formula. Geophys. Res. Lett. 43 (5), 2045–2052. https://doi.org/10.1002/ lution. Geomorphology 201, 344–362. https://doi.org/10.1016/j.geomorph.2013.07.010. 2015gl067542. Wiederkehr, M., Labhart, T.P., Schwizer, B., Gisler, C., Renner, F., Schweiz. Bundesamt für, L., Schweiz. Geologische, L., 2015. Meiental mit Südostteil von Blatt Engelberg, No.

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