# Accelerated retreat of northern James Ross Island ice streams (Antarctic Peninsula) in the Early-Middle Holocene induced by buoyancy response to postglacial sea level rise

Matěj Roman, Daniel Nývlt, Bethan Davies, Régis Braucher, Stephen J.A.

Jennings, Michal Břežný, Neil Glasser, Michael Hambrey, Juan Lirio, Ángel

Rodés

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Matěj Roman, Daniel Nývlt, Bethan Davies, Régis Braucher, Stephen J.A. Jennings, et al. Accelerated retreat of northern James Ross Island ice streams (Antarctic Peninsula) in the Early-Middle Holocene induced by buoyancy response to postglacial sea level rise. Earth and Planetary Science Letters, 2024, 641, pp.118803. 10.1016/j.epsl.2024.118803. hal-04628471

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# Accelerated retreat of northern James Ross Island ice streams (Antarctic Peninsula) in the Early-Middle Holocene induced by buoyancy response to postglacial sea level rise

Matěj Roman<sup>a,b,*</sup>, Daniel Nývlt<sup>b,c</sup>, Bethan J. Davies<sup>d</sup>, Régis Braucher<sup>e</sup>, Stephen J.

A. Jennings<sup>b,f</sup>, Michal Břežný<sup>g</sup>, Neil F. Glasser<sup>a</sup>, Michael J. Hambrey<sup>a</sup>, Juan M. Lirio<sup>h</sup>, Ángel Rodés<sup>i,1</sup>, ASTER Team<sup>e,2</sup>

## ARTICLE INFO

Edited by: Dr Y Asmerom

Keywords:

Palaeoglaciology

<sup>10</sup>Be and <sup>26</sup>Al cosmogenic dating

Antarctic deglaciation

Holocene

Ice stream

Grounding line

## ABSTRACT

The knowledge of dynamics and retreat patterns of marine-based ice streams under multiple stressors are of foremost importance for predicting Antarctic Ice Sheet response to climate changes. The Holocene palaeoglaciological record of former ice streams draining the northeast Antarctic Peninsula can elucidate the influences of changes in atmospheric and oceanic circulation and sea-level oscillations on the ice thinning and grounding line retreat. Here, terrestrial cosmogenic nuclide (TCN) dating of erratic boulders across the James Ross Island group sheds light on the pattern and timing of the ice recession along the two main arteries of the palaeo-ice drainage: Croft Trough and Prince Gustav Channel. The approach of using paired <sup>10</sup>Be-<sup>26</sup>Al nuclides enabled an assessment of cosmogenic isotope inheritance and complex burial-exposure history, notably on the high-altitude volcanic mesas. The TCN ages suggest that the Prince Gustav Channel Ice Stream was thinning from at least ~12 ka, with subsequent separation of the Antarctic Peninsula and James Ross Island ice masses by 10–8 ka. The transition from grounded ice to open marine conditions in the Croft Trough occurred rapidly at 8.6–7.2 ka, following the Early Holocene Warm Period, concomitant with eustatic and relative sea-level rise and incursions of warmer circumpolar waters. Grounding line retreat was possibly further accelerated by buoyancy response of thinning ice stream to low-gradient bed topography. The lessons of rapid deglaciation of James Ross Island palaeo-ice streams may provide analogues for recent or future intensification of pressures on Antarctic glaciers.

## 1. Introduction

Past and present cryospheric changes observed in the Antarctic Peninsula are expected to continue during 21st century due to

progressing disequilibrium with the anthropogenically enhanced climate forcing (Noble et al., 2020; DeConto et al., 2021; Hanna et al., 2024). This region is considered a sensitive component of the Earth's climate system, as evidenced by one of the most rapid observed rates of

<footnote><sup>a</sup> Centre for Glaciology, Department of Geography and Earth Sciences, Aberystwyth University, Aberystwyth SY23 3DB, Wales, United Kingdom</footnote>

<footnote><sup>b</sup> Polar-Geo-Lab, Department of Geography, Faculty of Science, Masaryk University, Kotlárníká 2, Brno 602 00, Czechia</footnote>

<footnote><sup>c</sup> Czech Geological Survey, Brno Branch, Leitnerova 22, Brno 658 69, Czechia</footnote>

<footnote><sup>d</sup> School of Geography, Politics and Sociology, Newcastle University, Newcastle-upon-Tyne NE1 7RU, United Kingdom</footnote>

<footnote><sup>e</sup> CEREGE, CNRS-IRD UM34, Aix-Marseille Univ., College de France, INRAE, BP 80, Aix-en-Provence 13545 CEDEX, France</footnote>

<footnote><sup>f</sup> Institute of Biochemistry and Biophysics, Polish Academy of Sciences, Pawińskiego 5A, Warsaw 02-106, Poland</footnote>

<footnote><sup>g</sup> Department of Physical Geography and Geocology, Faculty of Science, University of Ostrava, Chittussiho 10, Ostrava 710 00, Czechia</footnote>

<footnote><sup>h</sup> Instituto Antártico Argentino, 25 de Mayo 1143, General San Martín, Buenos Aires, Argentina</footnote>

<footnote><sup>i</sup> Scottish Universities Environmental Research Centre, Rankine Av, East Kilbride, Glasgow G75 0QF, United Kingdom</footnote>

<footnote>* Corresponding author at: Centre for Glaciology, Department of Geography and Earth Sciences, Aberystwyth University, Llandinam Building, Penglais Campus, Aberystwyth SY23 3DB, Wales, United Kingdom.
*E-mail addresses: mar133@aber.ac.uk, matej.roman@mail.muni.cz (M. Roman).</footnote>

<footnote n="1">¹ Currently at: Department of Geography, Universidade de Santiago de Compostela, Praza da Universidade 1, 15703, Spain</footnote>

<footnote n="2">² ASTER Team: Georges Aumaître, Karim Keddadouche, Fawzi Zaidi</footnote>
atmospheric warming globally in the second half of the 20th century (Vaughan et al., 2003; Turner et al., 2016, 2020). Current impacts of climate changes in the north-eastern Antarctic Peninsula include the disintegration of ice shelves (e.g., Cook and Vaughan, 2010; Baumhoer et al., 2021), the recession of land-based and tidewater glaciers (e.g. Engel et al., 2012; Cook et al., 2016; Seehaus et al., 2018) and variations in atmospheric and oceanic temperatures (e.g., Vaughan et al., 2003; Ambrožová et al., 2020; Turner et al., 2020). The sensitivity and proclivity of regional climate to large oscillations is emphasised by cessa- tion of warming after ~2000 CE (Turner et al., 2016; Oliva et al., 2017) and advances of ice shelf fronts along the eastern Antarctic Peninsula (Christie et al., 2022). Since 2016 CE, the Antarctic Peninsula experi- enced several air temperature extremes (Siegert et al., 2023), which might presage the return of the predicted warming trend (Carrasco et al., 2021; Clem et al., 2022) documented also on local glaciers (Stringer et al., in review).

Regional Holocene glacier dynamics and their interactions with climate, ocean and lithosphere are still poorly constrained despite their significance for contextualizing the ongoing cryospheric changes (e.g., Kaplan et al., 2020; Jones et al., 2022). The James Ross Island group is uniquely placed to provide a terrestrial record of the palaeoglacier

dynamics because it is one of the largest deglaciated areas of the Ant-
arctic Peninsula (Fig. 1; e.g., Smellie et al., 2008; Nývlt et al., 2011;
Glasser et al., 2014). Furthermore, the James Ross Island group is also
one of the most well studied regions in the Antarctic Peninsula regarding
the last deglaciation and Holocene glacier oscillations (Ingólfsson et al.,
1992; Björck et al., 1996; Hjort et al., 1997; Johnson et al., 2009, 2011;
Mulvaney et al., 2012; Davies et al., 2013, 2014; Glasser et al., 2014;
Nývlt et al., 2014; Minzoni et al., 2015; Čejka et al., 2020; Kaplan et al.,
2020; Roman et al., 2024; Jennings et al., unpubl.). Holocene palaeo-
environmental reconstructions from various natural archives (Mulva-
ney et al., 2012; Sterken et al., 2012; Minzoni et al., 2015) facilitate
comparisons with deglaciation records and provide insights into the
forcings that drive glacier behaviour as well as its reciprocal feedbacks
on the evolution of the deglaciated environment. There are unresolved
issues related to the changes in Southern Annular Mode, influencing the
position of Southern Westerly Winds and the upwelling of the Circum-
polar Deep Water (Barbara et al., 2016; Kaplan et al., 2020).

The convenient regional settings and plethora of palae-
oreconstructions enable to postulate the following research questions:
(i) what was the retreat pattern of the marine-based glaciers and ice
streams, with focus on their transition to ice shelves and land-based

<figure><img src="image_4.png" /></figure>

**Fig. 1.** Location of the study area. (A) Position of the Antarctic Peninsula within the Antarctic. (B) The James Ross Island group (denoted by red frame) is located in the Weddell Sea, on the eastern side of the Antarctic Peninsula. The ice extent as reconstructed by the RAISED Consortium is shown in 5-ka time slices (Bentley et al., 2014). (C) The localities where erratic boulders were sampled for TCN dating are denoted by red (solid and dashed) frames, with the numbers referring to Fig. 2 and Fig. 3, respectively. Shown are also the localities of palaeoenvironmental records obtained from Haddington Ice Cap core (Mulvaney et al., 2012) and lacustrine sediments on Beak and Vega Islands (Sterken et al., 2012; Čejka et al., 2020). In addition to REMA (Howat et al., 2019), the base maps are sourced from the geospatial package Quantarctica 3.2 (Matsuoka et al., 2021), including BedMachine v2 Antarctica bed topography (Morlighem et al., 2020), Landsat Image Mosaic of Antarctica (LIMA), and the SCAR Antarctic Digital Database (ADD) v7.0 shapefiles (http://www.add.scar.org/). The bathymetry around James Ross Island synthesises data referenced in Johnson et al. (2011). Inferred past ice flow directions (white arrows) and former ice shelf extent (dashed line) are after Johnson et al. (2011) and Pudsey et al. (2006), respectively.
glaciers; (ii) what was the connection between the atmospheric and oceanic forcing of ice thinning and ungrounding of ice masses, cryo-spheric tipping points and hysteresis; (iii) how can these palaeo-changes inform the present understanding of glacier dynamics in relation to changing climatic and marine conditions. The relatively high density of ages constraining the last deglaciation offer opportunities to explore these complex relationships, which might serve as analogues for other currently more glacierised regions or ice sheets. In this paper, we aim to extend the previously obtained deglaciation chronologies, using terres-trial in situ cosmogenic nuclide (TCN) exposure dating, from the areas adjoining the most dynamical palaeo-ice masses of the James Ross Island group, to understand the factors that led to the demise of these palaeo-ice streams.

# 2. Regional setting

## 2.1. Geology, glaciation, and climate

The James Ross Island group is located in the north-western part of the Weddell Sea, in the vicinity of the northernmost tip of the NE Antarctic Peninsula (Trinity Peninsula), from which it is sundered by the Prince Gustav Channel (Fig. 1). The James Ross Island group is composed of James Ross Island, Vega Island, Snow Hill Island, Seymour Island, and several other smaller landmasses. Glaciers and permanent snowfields account for more than 80 % of the archipelago's area. The bedrock of the James Ross Island group is composed of two main geological units, namely Cretaceous to Neogene sedimentary rocks and subglacial/subaqueous Neogene to Quaternary volcanic and volcanoclastic rocks with associated siliciclastic sediments (e.g., Francis et al., 2006; Nývlt et al., 2011; Smellie et al., 2013; Nehyba and Nývlt, 2015; Mlčoch et al., 2020). The majority of the volcanic rocks formed during glacial periods under ice cover (Smellie et al., 2008), giving rise to specific volcanic landforms, such as volcanic mesas, some of which attain heights of over 400 m a.s.l. (Lachman Crags, Davies Dome, Leal Bluff). However, some volcanic products, such as cinder and tuff cones (e.g. Terrapin Hill or Cerro Domo), were erupted under marine conditions during interglacial ice-poor periods (Smellie et al., 2008; Calabozo et al., 2015; Bertoa del Llano et al., 2021; Smellie, 2021). Granitoid and metasedimentary erratic boulders of Antarctic Peninsula provenance (Leat et al., 1995) are dispersed across the James Ross archipelago area both within the siliciclastic sequences of Neogene age and on the current surface (e.g., Hjort et al., 1997; Nývlt et al., 2011; Davies et al., 2013; Jennings et al., 2021).

During the Last Glacial Maximum, the Antarctic Peninsula Ice Sheet coalesced with the independent Mt. Haddington Ice Cap, situated over southern James Ross Island (Ó Cofaigh et al., 2014; Bentley et al., 2014). After ~18 ka (kiloanni = $10^3$ years), the overlying ice mass became increasingly dynamic with the development of the Prince Gustav Ice Stream, leading to regional ice thinning (Glasser et al., 2014; Nývlt et al., 2014), with the deglaciation of the Ulu Peninsula being underway by ~8 ka. The receding glaciers revealed a deglaciated landscape with large-scale features determined by the underlying bedrock geology, which had been sculpted by overlying ice masses during preceding glacial periods (Jennings et al., 2021). Post-deglaciation modification of the ice-free terrain superimposed paraglacial and periglacial geomorphological features upon the remnant glacial features (Davies et al., 2013; Ruiz-Fernández et al., 2019; Jennings et al., 2021). The recession of the glaciers in northern the James Ross Island group also led to the formation of numerous lakes, often with valuable palaeolimnological records (Björck et al., 1996; Sterken et al., 2012; Nedbalová et al., 2013; Písková et al., 2019; Roman et al., 2019; Čejka et al., 2020; Roman et al., 2024).

The climate of James Ross archipelago is determined by its position at the boundary between continental and maritime Antarctic, and its location in the precipitation shadow of the Antarctic Peninsula (van Wessem et al., 2015). It is characterized by cold, arid southerly barrier

winds and northwesterly foehn winds, contributing to strong surface melt events on glaciers and ice shelves (Clem et al., 2022; Laffin et al., 2022). Precipitation, mostly in the form of snow, is estimated at 300–500 mm a⁻¹ (van Lipzig et al., 2004). The mean annual air temperature in the Ulu Peninsula was –6.6°C in the period of 2004–2021 (Kaplan Pastíriková et al., 2023). The deglaciated areas of the James Ross Island group are underlain by continuous permafrost (Obu et al., 2020).

## 2.2. Study sites

The selection of sites for sampling rocks for TCN dating was based on their proximity to the palaeoglacier flow in the Prince Gustav Channel and Croft Troughs. The large altitudinal gradients of this volcanogenic relief also enable exploring the past ice surface elevation changes. Mesas represent ideal flat surfaces for erratic deposition, provided they were overrun not only by local ice caps, but also by ice flow from the Antarctic Peninsula (e.g. Glasser et al., 2014). The drainage basin of the Croft Trough (Johnson et al., 2011), a fjord extending over the Croft Bay and Herbert Sound, now represents one of the least glacierised ice catchments, in the entire Antarctic Peninsula region. The Croft Trough drainage basin covers ~700 km², of which ~300 km² is presently inundated by the sea and ~250 km² is glacierised (Minzoni et al., 2015). Tidewater glaciers are located only in the southernmost cirque of Croft Bay and in the north-western outlet glacier of Cape Lamb on Vega Island, totalling ~8 % of the coastline length (~107 km overall).

Clearwater Mesa (CWM; GPS [−64.02, −57.71]) and Sykes Cliffs are located in north-central James Ross Island, on the east and west side of Croft Bay, respectively (Fig. 1). Clearwater Mesa is bordered by the Haddington Ice Cap to the south and east, and the tidewater Blancmange Glacier to the north-east. Clearwater Mesa comprises a prominent lavafed delta (Smellie et al., 2013), with an altitudinal range of ~200–275 m a.s.l., and covers an area of ~8 km². This tuya (a flat-topped landform formed by subglacial volcanic eruption; Smellie et al., 2006) has a level surface, with a vertical relief of less than 30 m from valley floors to the tops of the tumuli (Roman et al., 2019). To the east, Clearwater Mesa consists of a lower-lying promontory, stripped of basaltic cap-rock, which connects the mesa with the ice cap.

Sykes Cliffs [−64.005, −57.83] appears to be a rotated slide-block region of volcanic mesa, overlooking Dreadnought Point, akin to the more northerly Lagoons Mesa (Calabazo et al., 2015). Alternatively, the correspondence of the cap-rock level altitude with Clearwater Mesa suggests a common origin at the beginning of the Pleistocene period when palaeo-ice thickness was similar (Smellie et al., 2013) or even in a single event of subglacial volcanism, with mesas subsequently dissected by erosion of the Croft Bay ice stream that formed the Croft Trough in-between.

Cape Lamb [−63.90, −57.61] covers ~27 km² in the south-western part of Vega Island, located to the north-east of James Ross Island across Herbert Sound. The Cape Lamb Member mudstones to sandstones of Late Cretaceous age form the basement of most of the peninsula's deglaciated area, overlain by Neogene to Quaternary cliffs of the James Ross Island Volcanic Group (Smellie et al., 2013).

Long Island [−63.77, −58.21] is located in Prince Gustav Channel, ~5 km from the Antarctic Peninsula. The shattered metamorphic bedrock of Trinity Peninsula Group (Smellie et al., 1996) differs from the James Ross Island sedimentary and igneous rocks, although the erratic-rich drift found here is easily discernible (Kaplan et al., 2020).

# 3. Materials and methods

Granitoid erratic boulders of distinct Antarctic Peninsula Batholith (Leat et al., 1995) provenance were targeted for TCN dating. Only large granite to diorite boulders (with b-axis >1.0 m wherever possible) on stable surfaces were sampled to reduce the probability of post-depositional peri-/paraglacial processes influencing the production
<table><caption>**Table 1**
Sample description and TCN measurements used for calculation of ¹⁰Be and ²⁶Al ages.</caption><thead><tr><th>Sample</th><th>Date sampled</th><th>Altitude (m a.s.l.)</th><th>Length (m)</th><th>Width (m)</th><th>Height (m)</th><th>Sample thickness (cm)</th><th>Topographic shielding</th><th>Pressure (mbar)</th><th>Total scaling</th><th>Quartz mass (g)</th><th>¹⁰Be at/g Qtz (± 1σ)</th><th>²⁶Al at/g Qtz (± 1σ)</th></tr></thead><tbody><tr><td>CWM-38</td><td>21/01/ 2016</td><td>194</td><td>3</td><td>1.5</td><td>0.8</td><td>2</td><td>0.99936</td><td>964.13</td><td>1.50113</td><td>14.559</td><td>42,894 ± 1124</td><td>193,031 ± 18,120</td></tr><tr><td>CWM-59</td><td>27/01/ 2016</td><td>255</td><td>2.5</td><td>1.1</td><td>0.9</td><td>2</td><td>0.99939</td><td>956.41</td><td>1.59793</td><td>19.401</td><td>55,635 ± 1285</td><td>316,016 ± 20,346</td></tr><tr><td>CWM-66</td><td>30/01/ 2016</td><td>230</td><td>1.5</td><td>1</td><td>0.45</td><td>2</td><td>0.99903</td><td>959.57</td><td>1.55719</td><td>10.911</td><td>54,120 ± 2211</td><td>276,177 ± 29,900</td></tr><tr><td>CWM-80</td><td>01/02/ 2016</td><td>170</td><td>2.1</td><td>1.55</td><td>0.8</td><td>2</td><td>0.99919</td><td>967.19</td><td>1.46394</td><td>21.042</td><td>39,157 ± 1629</td><td>231,510 ± 39,600</td></tr><tr><td>VL-37</td><td>12/02/ 2017</td><td>196</td><td>1.8</td><td>1.4</td><td>1</td><td>2</td><td>0.98750</td><td>963.88</td><td>1.48637</td><td>20.243</td><td>43,908 ± 1497</td><td>325,888 ± 25,644</td></tr><tr><td>VL-38</td><td>12/02/ 2017</td><td>183</td><td>2.2</td><td>2</td><td>1</td><td>2</td><td>0.99692</td><td>965.53</td><td>1.48049</td><td>19.977</td><td>48,710 ± 1845</td><td>309,550 ± 21,597</td></tr><tr><td>DM-1</td><td>23/01/ 2019</td><td>284</td><td>1.1</td><td>1</td><td>0.25</td><td>1</td><td>0.99896</td><td>952.76</td><td>1.65764</td><td>19.987</td><td>57,172 ± 3261</td><td>227,726 ± 67,000</td></tr><tr><td>LI-1</td><td>08/02/ 2019</td><td>31</td><td>2.05</td><td>1.7</td><td>1.5</td><td>2</td><td>1.00000</td><td>985.07</td><td>1.26280</td><td>20.186</td><td>40,833 ± 980</td><td>212,617 ± 16,945</td></tr><tr><td>LI-2</td><td>08/02/ 2019</td><td>81</td><td>2.3</td><td>1.7</td><td>1.1</td><td>2</td><td>1.00000</td><td>978.60</td><td>1.33339</td><td>20.059</td><td>36,681 ± 962</td><td>309,351 ± 79,060</td></tr><tr><td>LI-5</td><td>08/02/ 2019</td><td>42</td><td>2.2</td><td>1.8</td><td>1</td><td>2</td><td>1.00000</td><td>983.64</td><td>1.27813</td><td>19.993</td><td>42,550 ± 1057</td><td>249,889 ± 21,123</td></tr><tr><td>DN 07-020</td><td>21/01/ 2007</td><td>157</td><td>1.1</td><td>0.9</td><td>0.65</td><td>2</td><td>0.99798</td><td>968.85</td><td>1.46518</td><td>20.202</td><td>67,840 ± 3154</td><td>-</td></tr><tr><td>DN 07-025</td><td>21/01/ 2007</td><td>147</td><td>1.05</td><td>1</td><td>0.9</td><td>2</td><td>0.99422</td><td>970.12</td><td>1.44443</td><td>18.014</td><td>29,835 ± 1640</td><td>-</td></tr><tr><td>DN 07-026</td><td>21/01/ 2007</td><td>146</td><td>0.6</td><td>0.55</td><td>0.45</td><td>2</td><td>0.99551</td><td>970.25</td><td>1.44478</td><td>8.844</td><td>30,598 ± 2270</td><td>-</td></tr><tr><td>DN 07-027</td><td>21/01/ 2007</td><td>115</td><td>1.2</td><td>1</td><td>0.6</td><td>2</td><td>0.99190</td><td>974.22</td><td>1.39317</td><td>8.813</td><td>38,571 ± 2214</td><td>-</td></tr><tr><td>TH1</td><td>02/02/ 2012</td><td>570</td><td>0.82</td><td>0.43</td><td>0.45</td><td>1-2</td><td>0.99913</td><td>917.52</td><td>2.20606</td><td>15.0495</td><td>60,013 ± 3399</td><td>-</td></tr><tr><td>TH7</td><td>02/02/ 2012</td><td>264</td><td>0.58</td><td>0.48</td><td>0.28</td><td>1-3</td><td>0.99971</td><td>955.28</td><td>1.63849</td><td>23.0197</td><td>47,822 ± 2169</td><td>-</td></tr><tr><td>TH12</td><td>04/03/ 2012</td><td>90</td><td>0.97</td><td>0.66</td><td>0.4</td><td>&lt;2</td><td>0.99547</td><td>977.44</td><td>1.36135</td><td>22.7028</td><td>35,615 ± 2173</td><td>-</td></tr><tr><td>TH13</td><td>05/03/ 2012</td><td>35</td><td>1.3</td><td>0.86</td><td>0.64</td><td>2</td><td>0.99779</td><td>984.55</td><td>1.28550</td><td>22.3971</td><td>42,318 ± 2236</td><td>-</td></tr></tbody></table>
<table><caption>**Table 2**

Calculated TCN ages from the James Ross Island group with their geomorphological context. <sup>10</sup>Be, <sup>26</sup>Al and considered <sup>10</sup>Be-<sup>26</sup>Al ages are in years. Boulders are granites if not stated otherwise.</caption><thead><tr><th>Sample</th><th>Altitude (m a. s.l.)</th><th>Latitude (°)</th><th>Longitude (°)</th><th>Location and description</th><th><sup>10</sup>Be age (± 1σ)</th><th><sup>26</sup>Al age (± 1σ)</th><th><sup>10</sup>Be/<sup>26</sup>Al ratio (± 1σ)</th><th><sup>10</sup>Be-<sup>26</sup>Al age considered (± 1σ)</th></tr></thead><tbody><tr><td>CWM-38</td><td>194</td><td>-64.01660</td><td>-57.68970</td><td>fine-grained blockfield sloping towards the Blancmage Glacier; well-developed desert varnish, faceted</td><td>7053±420</td><td>4716±443</td><td>4.50±0.50</td><td>5951±661</td></tr><tr><td>CWM-59</td><td>255</td><td>-64.02406</td><td>-57.72205</td><td>the central plateau blockfield, near Katerina Lake; desert varnish largely desquamated</td><td>8601±444</td><td>7268±468</td><td>5.68±0.47</td><td>7973±658</td></tr><tr><td>CWM-66</td><td>230</td><td>-64.01139</td><td>-57.70416</td><td>northern plateau blockfield, near Soledad Lake; desert varnish grooved by striae at ~0-180°, rounded edges</td><td>8584±798</td><td>6511±705</td><td>5.10±0.73</td><td>7426±1060</td></tr><tr><td>CWM-80</td><td>170</td><td>-64.03444</td><td>-57.68011</td><td>promontory between mesa and Blancmange Gl., near Andrea Lake; desert varnish on top surface near fissures</td><td>6600±617</td><td>5800±992</td><td>5.91±1.15</td><td>6379±1244</td></tr><tr><td>VL-37</td><td>196</td><td>-63.90129</td><td>-57.61231</td><td>Cape Lamb, Leal Bluff slide-block region, hillslope ~10°; partly desquamated desert varnish</td><td>7291±558</td><td>7957±626</td><td>7.42±0.81</td><td>7625±837</td></tr><tr><td>VL-38</td><td>183</td><td>-63.89620</td><td>-57.60642</td><td>Cape Lamb, foothill of Leal Bluff slide-block region, hillslope ~12°; desert varnish, ploughing/solifluction?</td><td>8122±689</td><td>7659±534</td><td>6.36±0.70</td><td>7848±862</td></tr><tr><td>DM-1</td><td>284</td><td>-64.00690</td><td>-57.83331</td><td>on flat-topped mesa plateau; desert varnish, embedded into water-saturated patterned ground (low height above relief)</td><td>8522±1086</td><td>5043±1484</td><td>3.98±1.28</td><td>7317±2346</td></tr><tr><td>LI-1</td><td>31</td><td>-63.77656</td><td>-58.23042</td><td>in erratic-rich drift, rounded boulder with slender base, but well-embedded, mosaicked desert varnish</td><td>7973±430</td><td>6167±491</td><td>5.21±0.50</td><td>7191±692</td></tr><tr><td>LI-2</td><td>81</td><td>-63.76910</td><td>-58.21119</td><td>in erratic-rich drift, subrounded boulder with sparse desert varnish</td><td>6784±402</td><td>8512±2175</td><td>8.43±2.21</td><td>6842±1795</td></tr><tr><td>LI-5</td><td>42</td><td>-63.78464</td><td>-58.24879</td><td>in erratic-rich drift, subrounded boulder, incipient desert varnish</td><td>8210±458</td><td>7166±606</td><td>5.87±0.59</td><td>7831±793</td></tr><tr><td>DN 07-020</td><td>157</td><td>-63.85963</td><td>-58.10842</td><td>moraine crest with erratic-rich drift, partly desquamated desert varnish</td><td>11,541 ±536</td><td>-</td><td>-</td><td>-</td></tr><tr><td>DN 07-025</td><td>147</td><td>-63.85897</td><td>-58.10768</td><td>~3 m below moraine crest on a slope, frost shattered, desert varnish</td><td>5137±282</td><td>-</td><td>-</td><td>-</td></tr><tr><td>DN 07-026</td><td>146</td><td>-63.85922</td><td>-58.10794</td><td>~2 m below moraine crest on a slope, frost shattered</td><td>5290±392</td><td>-</td><td>-</td><td>-</td></tr><tr><td>DN 07-027</td><td>115</td><td>-63.85795</td><td>-58.10616</td><td>on hillslope of moraine, subrounded, desert varnish</td><td>6980±401</td><td>-</td><td>-</td><td>-</td></tr><tr><td>TH1</td><td>570</td><td>-63.96331</td><td>-57.54139</td><td>Terrapin Hill, near summit; gneissic diorite with red surface stain, on a flat surface with frost-shattered basalt gravel</td><td>6729±381</td><td>-</td><td>-</td><td>-</td></tr><tr><td>TH7</td><td>264</td><td>-63.98100</td><td>-57.55346</td><td>Terrapin Hill, S slope; rounded faceted boulder among broken lapilli tuff bedrock; slope ~3°</td><td>7209±327</td><td>-</td><td>-</td><td>-</td></tr><tr><td>TH12</td><td>90</td><td>-63.98645</td><td>-57.57913</td><td>Boulder Valley moraine ridge; rounded diorite well-embedded in drift of Cretaceous sediment,</td><td>6452±394</td><td>-</td><td>-</td><td>-</td></tr><tr><td>TH13</td><td>35</td><td>-63.98274</td><td>-57.60381</td><td>Boulder Valley above Shoreline Lake; mass movement crest; subrounded to subangular boulder with red veneer</td><td>8119±429</td><td>-</td><td>-</td><td>-</td></tr></tbody></table>
<figure><img src="image_5.png" /><figcaption>**Fig. 2.** Glacial geomorphological maps of the studied localities, with the REMA DEM topography of 8 m (and where available 2 m) resolution, rendered as hillshade (Howat et al., 2019). The maps designation (A)–(E) follows from the Fig. 1. TCN (circles) and $^{14}$C (triangles) dated deglaciation ages are denoted with their identificators.</figcaption></figure>
of TCN in rock. We preferentially aimed for boulders perched upon bedrock, blockfield or erratic-poor/erratic-rich drifts (~ground moraine), but boulders from Stoneley Point and one Terrapin Hill sample were obtained from moraine ridges. These moraines represent the former glacier limits and ages from such landforms should represent the periods of moraine formation, just preceding the deglaciation, and should be understood as minima for ice retreat (Kaplan et al., 2020). In total, 18 samples for surface exposure dating were collected from the upper surface (thickness max. 3 cm) of each boulder using a hammer and chisel.

All samples were crushed and sieved to 0.25–1 mm fraction prior to magnetic separation using a Frantz barrier separator. Quartz was isolated by froth floatation (SUERC samples; Herber, 1969). The extraction method for ¹⁰Be consists of isolation and purification of quartz and elimination of atmospheric ¹⁰Be. For samples prepared at CEREGE, approximately 150 µg of a 3025 ppm ⁹Be solution (~500 µg of a 405.5 ppm ⁹Be solution at SUERC) was weighed and added to the purified quartz, before the dissolution within HF. In ten samples, natural concentrations of ²⁷Al were measured in liquid aliquots by inductively coupled plasma–optical emission spectrometry (ICP–OES, Thermo Icap6500). If needed, ²⁷Al carrier (980 ± 4 ppm ²⁷Al solution from Merck) was added. Beryllium and aluminium were extracted by controlled elution from successive anion and cation resin extraction (DOWEX 1×8, then 50WX8) columns. The final hydroxides Be(OH)₂ and Al(OH)₃ precipitates were dried and heated at 800°C for 1 h to obtain BeO and Al₂O₃. These oxides were mixed respectively with niobium and silver powders prior to AMS measurements at the ASTER AMS Facility at CEREGE in Aix-en-Provence, France. Four samples from Terrapin Hill were measured at SUERC, East Kilbride, UK. ¹⁰Be data were calibrated from an ASTER in-house preparation “STD11”, with a ¹⁰Be/⁹Be ratio of (1.191 ± 0.013) x 10⁻¹¹ (Braucher et al., 2015) using a ¹⁰Be half-life of (1.387 ± 0.0012) x 10⁶ years (Chmeleff et al., 2010; Korschinek et al., 2010), and “NIST SRM4325” for SUERC samples, with a ¹⁰Be/⁹Be ratio of 2.79 × 10⁻¹¹. ²⁶Al/²⁷Al ratios were calibrated with an ASTER in-house preparation “SM-AL-11” of (7.401 ± 0.064) x 10⁻¹² ratio (Merchel and Bremser, 2004) using a ²⁶Al half-life of (7.05 ± 0.17) x 10⁵ years (Samworth et al., 1972). Reported analytical uncertainties include counting statistics, machine stability (~0.5 % for ¹⁰Be; Arnold et al., 2010), and blank correction. ¹⁰Be/⁹Be and ²⁶Al/²⁷Al blank ratios were 5.70 × 10⁻¹⁶ (± 23 %) and 9.00 × 10⁻¹⁶ (± 100 %), respectively.

A modern ¹⁰Be spallation production rate at sea-level and high-latitude of (4.02 ± 0.16) at g⁻¹ yr⁻¹ was used (Borchers et al., 2016). Muon production follows the method of Braucher et al. (2011). The sea-level and high-latitude spallation production rate was then scaled for individual sample altitudes and latitudes using Stone (2000) polynomial whereas muon contributions were only scaled for altitude using Antarctic atmospheric pressure. The topographic shielding factor was calculated from the field observations. The snow shielding was neglected due to there being no constraints on the past snow cover thickness changes, and the sampled boulders standing proud on the surface, mostly well above the current average snow height, which is estimated at ~30–50 cm for the northern James Ross Island (Kňažková et al., 2020).

In order to determine ¹⁰Be exposure ages (Table S1) from the ¹⁰Be concentrations measured in the quartz fractions we used the equation of Braucher et al. (2011). Age uncertainties include 1σ error on counting statistics and instrumental error derived from multiple standard measurements. However, it is important to recall that external uncertainty associated with the production rate (on the order of 5 %) must be added to the analytical ones when comparison with other chronologies is performed (Stone, 2000). All ages have been calculated with a constant production rate implying that no correction was made for geomagnetic field variations, which are negligible in Antarctic latitude. ¹⁰Be exposure ages were considered to represent best the timing of deglaciation, since the production rate and scaling scheme are best known for this isotope (Suganuma et al., 2022), as well as to streamline the comparison with

the published, largely ¹⁰Be-based, deglaciation chronologies. Recognising the issue of complex exposure history, we also present ‘considered’ ¹⁰Be–²⁶Al ages, calculated using the approach explained in Pappu et al. (2011) and Braucher et al. (2022). We base on the fact that both nuclides, originating from the same mineral, have experienced the same exposure history, thus the age and denudation should be similar. For samples with statistically undistinguishable ¹⁰Be and ²⁶Al ages, we assume no inheritance and consider ¹⁰Be age to indicate the deglaciation, whereas for samples with low ²⁶Al/¹⁰Be ratio, the considered ¹⁰Be–²⁶Al age better represents the estimated time since the last exposure.

The published ¹⁰Be surface exposure ages from cited sources (89 in total; Table S2) were recalculated using the above-mentioned approach. CosmoCalc was used to produce the co-isotope (‘banana’) plot (Vermeesch, 2007). The normal Kernel Density Estimate and Probability Density Function of the age distributions was performed in Density-Plotter v8.5 (Vermeesch, 2012). Recalibration of the published ¹⁴C ages (14 in total) was carried out in OxCal v. 4.4 (Bronk Ramsey, 2009) using the SHCal20 atmospheric and Marine20 curves (Heaton et al., 2020; Hogg et al., 2020). Radiocarbon data are reported as conventional ¹⁴C years BP with 1σ error, and as calibrated years with 2σ (Table S3).

## 4. Results

The cosmogenic data from the James Ross Island group granitoid erratic boulders are shown in Table 1, and the calculation of the estimated time since their last exposure from beneath ice cover in Table 2. On Clearwater Mesa, four granite erratic boulders were sampled (Fig. 2), which yielded ¹⁰Be ages in the range 8.6–6.6 ka (Figs. 3 and 4). Two older samples, CWM-59 and CWM-66, located on the high-lying central blockfield of the mesa, give ¹⁰Be ages 8.60 ± 0.44 and 8.58 ± 0.80 ka. The samples CWM-38 and CWM-80 return ages 7.05 ± 0.42 and 6.60 ± 0.62 ka, and these are situated in a more ice-proximal and lower-altitude settings. The erratic CWM-66 bears marks of glacial striations, that were infrequently observed also on basaltic boulders and in one case on the basaltic bedrock outcrop on Clearwater Mesa (Fig. 5). On the other side of the Croft Bay (along the transverse altitudinal profile i-i', Fig. 3), the ¹⁰Be age of the sample DM-1 from Sykes Cliffs is 8.52 ± 1.09 ka, although DM-1 displays the lowest ¹⁰Be/²⁶Al ratio, and hence discrepancy between the ¹⁰Be and ²⁶Al (but also the highest analytical errors). DM-1 has the lowest prominence of sampled boulders, with a height of only 0.25 m emanating above the terrain. Terrapin Hill samples yield ¹⁰Be ages between 8.1 and 6.5 ka. The lowest erratic TH13, situated closest to the present coastline, was exposed earliest by 8.12 ± 0.43 ka. Samples from the Terrapin Hill southern slope and summit, TH7 and TH1, give lower ages of 7.21 ± 0.33 and 6.73 ± 0.38 ka, respectively. The diorite boulder TH12 on a moraine ridge returns an exposure age of 6.45 ± 0.39 ka.

Samples VL-38 and VL-37 from Cape Lamb yield ¹⁰Be ages 8.12 ± 0.69 and 7.29 ± 0.56 ka. These samples come from erratics on the sloping foothill of Cape Lamb volcanic delta, yet the close correspondence of ages, especially considered ¹⁰Be–²⁶Al ages of ~7.8 and ~7.6 ka, respectively, suggests no post-depositional movement affecting the dates. ¹⁰Be ages from the Long Island erratics are within range of 8.21 ± 0.46 to 6.78 ± 0.40 ka, with the apparent direction of ice recession from NE to SW. Stoneley Point erratics have a wide range of ages from 10.57 ± 0.49 to 4.62 ± 0.25 ka. The ages of the two youngest samples are statistically indistinguishable.

Kernel density estimate based on mean cosmogenic ages places the peak age of the deglaciation at 7.0–7.2 ka both in the Croft Trough and Prince Gustav Channel areas, although the samples from Long Island and Stoneley Point display larger scatter (Fig. 6). The relationship between ¹⁰Be and ²⁶Al ages is shown in Fig. 7. The co-isotope plot displays ages of all samples at or below the line consistent with no erosion and constant exposure, at 1σ. At 2σ error, CWM-38 and LI-1 are entirely within the zone of complex exposure or intermittent shielding, even though DM-1 exhibit the lowest ¹⁰Be/²⁶Al ratio (but also the highest analytical
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** (A) Map of Croft Trough and surrounding landmasses, with published and new deglaciation ages (in ka). Our cosmogenic ages ($^{10}$Be in black and $^{26}$Al in red font) are shown in white boxes. The transverse altitudinal transects are denoted by red lines with small roman numerals. The inset frame depicts the Long Island. TCN ages are denoted by circles, $^{14}$C ages by triangles, OSL ages by diamonds. (B) Four transverse profiles through the Croft Trough show the deglaciation age relationships on the western and eastern (true left and right) flanks of the palaeo-ice stream, ordered from north to south (i.e. trough mouth to headwall) as i-i' to iv-iv'. Note that the older TCN samples from Nývlt et al. (2014), not representing the last deglaciation, are not shown.</figcaption></figure>

uncertainty).

# 5. Discussion

## 5.1. Interpretation of TCN surface exposure ages and glacial geomorphology

The new cosmogenic nuclide exposure data constrain the Holocene deglaciation around Croft Bay and Prince Gustav Channel, subsequent to the establishment of a newly independent Mount Haddington Ice Cap following the drawdown of the mainland Antarctic Peninsula Ice Sheet. During and just after the LGM, the Antarctic Peninsula Ice Sheet and Mt. Haddington Ice Cap were drained mainly through several ice streams and glaciers occupying the deepest areas, i.e. Prince Gustav Channel Ice Stream (Evans et al., 2005) and Croft Trough Glacier (Johnson et al., 2011; Minzoni et al., 2015). Ice thinning on the landmasses of the James Ross Island group has been underway since ~18 ka in the higher localities, such as Lachman Crags (Glasser et al., 2014), or by ~13 ka on Cape Lachman, near the confluence of the Prince Gustav and Croft Trough ice streams (Ingólsson et al., 1992; Johnson et al., 2011; Nývlt et al., 2014). The new TCN data agree well with, and supplement, other TCN and $^{14}$C deglaciation chronologies from this area of northern James Ross Island (Ingólsson et al., 1992; Hjort et al., 1997; Johnson et al., 2011; Glasser et al., 2014; Nývlt et al., 2014; Minzoni et al., 2015; Kaplan et al., 2020; Fig. 8) and wider Antarctic Peninsula region (e.g. Ó Cofaigh et al., 2014; Heredia Barión et al., 2023). We estimate $^{10}$Be

exposure ages to best indicate the timing of deglaciation; nonetheless, the coupled $^{10}$Be-$^{26}$Al results demonstrate the limitations posed by complex exposure history, and thus treat $^{10}$Be ages as maxima.

Marine geological data indicate that a grounded outlet glacier draining the Mount Haddington Ice Cap, occupied Herbert Sound during the Last Glacial Maximum. This 'Croft Trough Glacier' thinned to the point of flotation in Herbert Sound at 10 ± 2.4 ka, induced by atmospheric warming concomitant with the sea level rise (Minzoni et al., 2015). Subsequently this floating ice shelf disintegrated/receded in the Early Holocene between 10 to 7.2 ka when open marine conditions commenced (Minzoni et al., 2015). As the grounded ice receded from Cape Lachman, deglaciation of the Naze ensued next, with surface exposure ages of 10–7.4 ka (Johnson et al., 2011) and slightly lower $^{14}$C ages of 7.7–7.0 cal. Ka BP (Hjort et al., 1997). The Naze is, however, already located farther from the influence of the Croft Trough palaeo-ice stream, and the ice rather flowed to the Humps Trough and subsequently northeastward to the Vega Basin (Johnson et al., 2011). In the areas immediately adjacent to the Croft Trough palaeo-ice stream, thinning of the glacier surface to below 200 m a.s.l. occurred at ~8.6–7 ka, resulting in near-synchronous exposure of sample sites at this altitude (Figs. 3, 9). The deglaciation started at Clearwater Mesa (~250 m a.s.l.) at ~8.6 ka, and the Sykes Cliffs (~280 m a.s.l.) at ~8.5 ka ago. The higher-lying localities of Cape Lamb (~190 m a.s.l.) on Vega Island became deglaciated at 8.1–7.3 ka, which corresponds to the onset of ice retreat by ~7.7 ka on the opposite shore of Herbert Sound, in Santa Martha Cove and Alpha Glacier foreland (Glasser et al., 2014; Kaplan et al., 2020). The
<figure><img src="image_7.png" /><figcaption>Fig. 4. Photographs of sampled erratic boulders and their $^{10}$Be and $^{26}$Al exposure ages with 1σ denoted (in white and red, respectively). (A–D) Samples from Clearwater Mesa, James Ross Island; (E, F) Samples from Cape Lamb on Vega Island; (G-I) Sampled boulders from the Long Island; (J-M) Samples from Stoneley Point, James Ross Island; (N-Q) Sampled boulders from Terrapin Hill, James Ross Island. We assume zero erosion on the boulder surface after glacier deposition, as the granite boulders are resistant rocks with often developed desert varnish that was targeted for sampling, even if partial desquamation was present.</figcaption></figure>

<figure><img src="image_8.png" /><figcaption>Fig. 5. Examples of striations observed in north-eastern Clearwater Mesa. (A) Dated boulder CWM-66 has developed striae of 0–180° orientation, with probable ice flow direction S to N (presuming the striations were formed following deposition). (B) Basaltic bedrock with two transposed sets of striae at azimuths of ~75–255° (dominant) and 130–310° (secondary), with inferred ice flow direction WSW to ENE and SE to NW, respectively.</figcaption></figure>

low-lying Boulder Valley beneath Terrapin Hill began deglaciating by ~8.1 ka. Only the (at present ice-proximal) Clearwater Mesa promontory, verging on the Blancmange Glacier, and Terrapin Hill slopes were deglaciated later around 6.6–6.5 ka, respectively. The deposition of erratic boulders at Dreadnought Point, downslope of the Sykes Cliffs, at ~6.8 ka (Kaplan et al., 2020) also postdates other adjacent ages, but these should be treated as minima. Terrestrial evidence of deglaciation from the Croft Bay area suggests that ice thinning occurred rapidly, with an identifiable altitudinal median gradient of ~0.5 m a⁻¹ (with 95 %

confidence interval of 0.15–8.14 m a⁻¹, Fig S1, based on Croft Trough TCN ages, excluding those from distal Cape Lachman and minimal ages of Kaplan et al., 2020, using a Monte Carlo regression estimate of ‘ice-TEA’ model; https://icetea.org; Jones et al., 2019). The timing also exactly matches the inferred full deglaciation of the Herbert Sound by ~7.2 ka BP at the marine sediment coring Site 2 NBP0502 (Minzoni et al., 2015).

The Prince Gustav Channel Ice Stream began to thin between 18 and 12.2 ka (Glasser et al., 2014) and grounding-line reached the Röhss Bay
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Kernel density estimate (thick black line) and Probability density function (grey line) of the $^{10}$Be exposure ages from the Croft Trough and Prince Gustav Channel ice-stream catchment areas with the mean age (μ ± 1σ). Additionally, weighted mean of the sample ages after discarding statistical outliers that were identified with a Chi² test (Fig. S1; Ward and Wilson, 1978) is shown. Individual exposure ages are denoted by a Gaussian curve with width proportional to their 1σ. Lines are colour-coded by their location as in Fig. 7; moreover, yellow and light green lines represent samples from Terrapin Hill and Stoneley Point, respectively.</figcaption></figure>

area at ~10.8 cal. ka BP (Evans et al., 2005; Pudsey et al., 2006). Long Island was deglaciated between 8.2 and 6.8 ka according to our $^{10}$Be ages from granite erratic boulders. These ages correspond with the four younger erratic cobble samples of Kaplan et al. (2020) from the NE part of Long Island, that range from 7.6 ± 0.3 to 6.9 ± 0.2 ka. No trend in altitude or direction of ice retreat was confirmed for Long Island. The two older cobble samples of Kaplan et al. (2020) at ~9.7 ka equal the age of deglaciation across the Prince Gustav Channel, on Stoneley Point moraine (Glasser et al., 2014). The age of sample DN07-020 is ~3 ka higher than the age of immediately adjacent boulder JRI62 (Glasser et al., 2014; recalculated), but coincides with the age of the San Carlos Hill samples ~4 km to the northeast. Our other three younger samples (6.2–4.6 ka) from Stoneley Point were found on a slope beneath the moraine crest, which suggests a readvance with the later local deglaciation, or possibly post-depositional disturbances. Overall, these TCN ages on both sides of the trough, as well as the more northerly located palaeolimnological record on Beak Island (Sterken et al., 2012), pinpoint the timing of the fastest ice mass lowering and deglaciation of the Prince Gustav Channel palaeo-ice stream to the period 10–6.8 ka.

The striations discovered on Clearwater Mesa (Fig. 5) suggest the mesa was at least in its lower-lying northern part overrun by wet-based or polythermal erosive ice mass, such as on the higher-altitude Patalamon Mesa (Johnson et al., 2009). However, the erosive action must have been limited in scope, as the primary volcanogenic relief with tumuli is well preserved across the mesa (Roman et al., 2019). The glaciation at

Patalamon Mesa was only sufficient to erode the bedrock in depressions between the tumuli, indicating the maximum ice thickness there was close to 45–200 m, when the basal melting point is achieved under typical local boundary conditions (Johnson et al., 2009). Clearwater Mesa is located at a lower altitude than Patalamon Mesa, suggesting shorter duration of glaciation and longer ice-free time windows for frost shattering and periglacial action to develop the autochthonous blockfields found almost ubiquitously across the mesa (at larger scale, the tumuli with its relief in tens of metres are well preserved, but the surface is periglacially shattered only surficially, down to ~1 m depth). Moreover, the location of several samples' concentrations below the line of simple exposure history (especially CWM-38; Fig. 7 A) indicates that these erratics may have been exposed and buried multiple times during the Quaternary (Bentley et al., 2006; Strub et al., 2015). Possible causes involve the reworking of the Antarctic Peninsula erratic material from older glacigenic and glaciomarine Neogene diamictons (Davies et al., 2013; Nývlt et al., 2014), found in outcrops interbedded between the underlying Cretaceous sediments and overlying volcanic strata (Hambrey and Smellie, 2006 Nelson et al., 2009; Nývlt et al., 2011). Alternatively, the assumption that all erratics were deposited in the last deglaciation is not valid (e.g. Johnson et al., 2011; Glasser et al., 2014) and the Antarctic Peninsula Ice Sheet has been transporting and depositing the granite erratics already in the previous glacial-interglacial cycles (Davies et al., 2013). Such interpretation is supported by the evidence from formerly glaciated plateau landscapes, such as the Cairngorms, that were shaped through multiple Pleistocene glacial periods (Hall et al., 2024). Their conclusion then recommends leaning on the youngest exposure ages as those least influenced by nuclide inheritance.

The co-isotope 'banana' plot (Fig. 7 A) indicates that some of the samples on volcanic mesas, including the DM-1 on Sykes Cliffs and to a lesser degree CWM-38 and CWM-66 on Clearwater Mesa, may have been buried for over 1 Ma, or even 2 Ma at 2σ, thus approaching the age of formation of the mesas – Sykes Cliffs, and Clearwater Mesa probably as well, were formed in the early Pleistocene, according to $^{40}$Ar/$^{39}$Ar dating of cap-rock lavas (Smellie et al., 2008; Roman et al., 2019). However, complex burial history may not be well discernible with $^{10}$Be and $^{26}$Al on time scales of <10⁴ ka due to their long half-lives and $^{26}$Al measurement errors. Therefore, the cosmogenic ages with low $^{26}$Al/$^{10}$Be ratio should be treated with care and only as the very upper limit of the timing of deglaciation due to the possibly complex exposure history and nuclide inheritance.

## 5.2. Holocene pattern of palaeoglacier behaviour and palaeoclimatic teleconnections

The wider palaeoenvironmental implications of the behaviour of the Prince Gustav Channel and Croft Trough palaeo-ice streams pertain to the evolution of Late-Glacial to Holocene climates in the Antarctic, Patagonia and Southern Ocean. The Croft Trough Glacier transition from marine-based grounded ice to ice shelf occurred during the rapid eustatic sea level rise (Sterken et al., 2012; Glasser et al., 2014; Lambeck et al., 2014), following the Early Holocene warm period (Bentley et al., 2009; Mulvaney et al., 2012; Jones et al., 2022). The local relative sea level (RSL) was also rising as the Antarctic glacio-isostatic adjustment was outpaced by the melting Northern Hemisphere ice sheets in the final stages of last deglaciation (Lambeck et al., 2014; Roberts et al., 2011). Regional RSL reconstructions show an Early Holocene maximum, or 'high-stand', of ~15 m above present sea level occurred at ~8 ka, following by gradual fall to the present day as the isostatic rebound responded to local deglaciation (Roberts et al., 2011; Theilen et al., 2023).

The absence of moraines in the trough indicates a rapid and uninterrupted grounding line recession between 10 ± 2.4 and ~7.2 ka, with likelihood skewed towards the latter of these ages (Minzoni et al., 2015). Except for the somewhat older ages from Cape Lachman (Ingólfssson
<figure><img src="image_10.png" /><figcaption>Fig. 7. (A) Co-isotope graph of $^{10}$Be and $^{26}$Al concentrations/ratios plots all samples at or below the line consistent with no erosion and constant exposure, within the $1\sigma$ error. At $2\sigma$, only CWM-38 and LI-1 are entirely within the zone of complex exposure or intermittent shielding, although DM-1 displays the lowest $^{10}$Be/$^{26}$Al ratio (but also the highest uncertainty). (B) Scatter plot of $^{10}$Be vs $^{26}$Al ages, with their respective $1\sigma$ uncertainties.</figcaption></figure>

et al., 1992; Johnson et al., 2011; Nývlt et al., 2014) and a disputable $^{14}$C age from a moraine in east Cape Lamb (Zale and Karlén, 1989), most ages in the proximity of the Croft Trough (Hjort et al., 1997; Johnson et al., 2011; Glasser et al., 2014; Kaplan et al., 2020; this study) posit the most pronounced ice-surface lowering to the period 8.6–7.2 ka along the entire longitudinal profile, thus corroborating the hypothesis of rapid thinning and transition to floating ice shelf (Minzoni et al., 2015). The Croft Trough currently attains depth of 300–400 m throughout its length, so grounded ice would have to thin to approximately 50 m above (present) sea level to begin floating. For the central basin of the Croft Trough to become ungrounded by ~7.2 ka (Minzoni et al., 2015), the ice cover would have needed to lower by ~230 m, from mesa tops to a plausible ice shelf freeboard (height of the ice-shelf surface above sea level), within a few centuries between ~8.6 and ~7.2 ka. The floating ice shelf then must have disintegrated imminently, for the open marine conditions commenced at ~7.2 ka, inferred from abundant open water diatoms and pelagic sedimentation suggesting minimal glacial influence (Minzoni et al., 2015). Such a rate of ice thinning is an order of magnitude higher than that inferred for Prince Gustav Channel Ice Stream, which lowered >230 m during the interval 18 to 12.2 ka (Glasser et al., 2014), but broadly comparable to other Antarctic surface exposure (Small et al., 2019; Johnson et al., 2020; Susanuma et al., 2022; Balco et al., 2023) and ice core records (Grieman et al., 2024). The new ice-core-derived record of Grieman et al. (2024) from the more southerly Skytrain ice rise in the Weddell Sea Embayment also indicates such a sequence and timing of events, with the dynamic thinning occurring first due to ungrounding around 8.2–7 ka, followed by the ungrounded ice shelf retreat at 7.7–7.3 ka.

The mechanism of grounding line retreat, coincidental with the ice thinning and concomitant sea level rise, might have been analogous to rapid recession described at low-gradient continental shelf areas near Larsen Inlet (west of Prince Gustav Channel; Dowdeswell et al., 2020) or mid-Norwegian shelf (Batchelor et al., 2023). Investigations of buoyancy-driven grounding line retreat at these localities suggest that rates of several hundred metres per day can temporarily be reached (Batchelor et al., 2023), although substantially lower retreat rates would still suffice to explain a multicentennial transition of grounded glacier to ice shelf in the Croft Trough. Once the recession began, however, the low-gradient ice-stream bed may have promoted an irreversible and unstoppable grounding line retreat, given the lack of recessional

moraines and other landforms suggesting temporary standstills (Minzoni et al., 2015). Rapid grounding line retreat in the Early Holocene was inferred also from other parts of Antarctica (Jones et al., 2015; Johnson et al., 2020, 2022; Nichols et al., 2023; Venturelli et al., 2023; Grieman et al., 2024), making it a suitable analogue for studying the present-day signs of marine ice sheet instability (Favier et al., 2014; Rignot et al., 2014). The retrograde (landwards-sloping) topography of the Prince Gustav Channel is predisposed to such marine ice sheet instability, even to a greater degree than for a relatively flat-bedded Croft Trough, although the rate of grounding line retreat of this ice stream is less easily elucidated despite the abundance of ages constraining the ice recession.

The deglaciation of Prince Gustav Channel, as deduced from Long Island and NW coast of the Ulu Peninsula, probably began earlier than in Croft Trough, at ~10 ka (Figs. 3, 6; Glasser et al., 2014; Kaplan et al., 2020). The ice lowering/retreat triggered the separation of the Antarctic Peninsula Ice Sheet and local James Ross Island ice earliest around Cape Lachman (Fig. 9; Ingólfsson et al., 1992), where these ice-flow units confluence. The older ages ascertained from Cape Lachman could be due to the earlier splitting of the Croft Trough and Prince Gustav Channel ice streams (Johnson et al., 2011; Nývlt et al., 2014), and generally by their more distal location from the ice divides or a topographic control induced by the nearby mouth of the Croft Trough (Minzoni et al., 2015). A $^{14}$C age of ~10.8 ka BP (considered as maximum) from glaciolacustrine unit at Cape Lachman is thus indicates the presence of a terminoglacial lake, dammed by a glacier in the Croft Trough just beginning its retreat (Fig. 9 A; Ingólfsson et al., 1992), whereas the northern-facing side of Cape Lachman may have been ice-free following the earlier retreat of the Prince Gustav Channel ice stream (Johnson et al., 2011; Nývlt et al., 2014). However, the possible persistence of an ice shelf in the Prince Gustav Channel after the ungrounding (Davies et al., 2012) has implications for the deglaciation of the Ulu Peninsula interior, especially the area of the Abernethy Flats and Monolith Lake, due to the buttressing and back-stress exerted on the tidewater glaciers (e.g., Hodgson et al., 2018; Johnson et al., 2020). After the loss of Prince Gustav Channel Ice Shelf (Pudsey et al., 2006), the tidewater glaciers retreated into the interior of the Ulu Peninsula, exposing the large amphitheatre-like deglaciated area around Brandy Bay (Davies et al., 2014; Roman et al., 2024; Jennings et al., *unpublished*).

Concerning the palaeoclimatic teleconnection patterns, most of the
<figure><img src="image_11.png" /><figcaption>**Fig. 8.** Oblique views to the south (A) and to the north (B) from above James Ross Island towards the Croft Trough and adjacent landmasses on REMA DEM and bathymetry. The $^{10}$Be maximum deglaciation ages from this study and other published ages (in ka) as well as inferred palaeo-ice streams flow direction are shown (cf. Fig. 3, where also moraines readvance ages of Kaplan et al., 2020, are displayed). The sample notation follows colour coding in Fig. 3.</figcaption></figure>
<figure><img src="image_12.png" /><figcaption>**Fig. 9.** Tentative reconstruction of glacier extent/grounding line migration at time slices for ~10, 8, 6 and 0 ka in the James Ross Island group and adjacent Antarctic Peninsula, constrained by the available deglaciation ages. A colour scale denotes a qualitative degree of confidence in glacier margins position. Note that contrary to the reconstruction of the Raised Consortium (Bentley et al., 2014; Ó Cofaigh et al., 2014), the 10-ka time slice displays a more progressed state of deglaciation in Prince Gustav Channel, both from the south and east.</figcaption></figure>

deglaciation occurred immediately after the Early Holocene Warm Period (Fig. 10; Mulvaney et al., 2012), when the band of Southern Westerly Winds shifted poleward and the Southern Annular Mode was persistently positive (Moreno et al., 2018; Kaplan et al., 2020; Roberts et al., 2022; Heredia Barión et al., 2023). Solar forcing, namely the increased spring/summer insolation on the southern hemisphere and

elevated solar irradiance, together with the overall postglacial reor-
ganisation of atmospheric and oceanic circulation, may have contrib-
uted to these centennial to millennial excursions of the westerlies,
bringing more frequently warm air masses from lower latitudes (Heredia
Barión et al., 2023). Such atmospheric forcing exerted influence on
Weddel Gyre circulation and probably promoted Upper Circumpolar
<figure><img src="image_13.png" /><figcaption>**Fig. 10.** Synthesis of palaeoclimatic records pertinent to Croft Glacier retreat: A) deglaciation ages from the Croft Trough area (denoted by circles for TCN ages and triangles for $^{14}$C ages with 1σ error bars, and coded by colour as in Fig. 3) are plotted against altitude (in metres above present sea level). Moraine ages of Kaplan et al. (2020) and sample TH12 represent periods of glacier expansions and can be interpreted as minima of deglaciation age before ice readvance. A weighted least square regression (constructed using 'iceTEA' model; Jones et al., 2019) of ice lowering rate based on TCN ages from Croft Trough (excluding distal ages from Cape Lachman and abovementioned minimal ages) is displayed in black box and approximated by a black arrow; B) palaeoenvironmental inferences from the Croft Bay marine sediment core NBP0502-2 and multi-beam swath bathymetry data (Minzoni et al., 2015); C) RSL curve based on Beak Island isolation basins (Roberts et al., 2011) with marine limits inferred from macrofossils in Brandy Bay (squares) and The Naze (diamonds) of Hjort et al. (1997); D) ice volume equivalent global sea-level curve of Lambeck et al. (2014), with indicated events of rapidly rising eustatic sea levels. MWP denotes a meltwater pulse; E) loss-on-ignition (LOI 550°C) as a proxy for organic content in Beak Island Lake 1 sedimentary record, with higher values interpreted as warmer water and air temperatures (Sterken et al., 2012); F) reconstructed air temperature anomaly over James Ross Island from Mt. Haddington ice core based on deuterium variations (Mulvaney et al., 2012), with pan-Antarctic millennial-scale climatic excursions denoted; G) $\delta^{18}$O and Na concentration data from Skytrain Ice Rise ice core [−79.74, −78.54] (Grieman et al., 2024) as proxies for ice surface elevation and marine influence, respectively, with thick lines showing 20-yr bin averages. The raw $\delta^{18}$O and sodium data are truncated at higher values for clarity. Periods of the inferred dynamic ice thinning and grounding line retreat are outlined by coloured boxes.</figcaption></figure>
Deep Water advection onto the inner eastern Antarctic Peninsula con-
tinental shelf (Spence et al., 2014; Barbara et al., 2016), accelerating
regional ice shelf and grounding line retreat, and inducing a shorter sea
ice season. A shared regional signal of oceanic and atmospheric forcing
is also supported by a close concurrence with the timing of ice thinning
and ungrounding from the Skytrain ice record (Fig. 10, G), located
further south in the Weddell Sea Embayment (Grieman et al., 2024),
although local modulations of ice sheet dynamics may have acted
differently compared to our narrower, topographically constrained ice
streams.

RSL reached its apex in the Early Holocene, as eustatic sea level rise outpaced isostatic readjustment (Hjort et al., 1997; Roberts et al., 2011; Sterken et al., 2012; Glasser et al., 2014; Suganuma et al., 2022). The combined oceanic and atmospheric forcing caused the retreat and lowering of the James Ross Island ice masses, after which Croft Trough did not host an ice shelf nor grounded ice throughout the remainder of the Holocene (Minzoni et al., 2015). This contrasts with the more southerly Larsen A and south Prince Gustav Channel ice shelves, which re-formed during the Neoglacial (Brachfeld et al., 2003; Balco et al., 2013). The ice masses flowing into the Croft Trough likely crossed the tipping point where the ice shelf could not re-form due to hysteresis (Schoof et al., 2007), unlike the more southern regions of the eastern Antarctic Peninsula (Balco et al., 2013). The ‘Mid-Holocene Hyp- sithermal’, considered to peak at 5–2 ka (Mulvaney et al., 2012; Sterken et al., 2012; Kaplan et al., 2020), is another period of marked ice retreat, especially in the interior of James Ross Island (Björck et al., 1996; Roman et al., 2024; Jennings et al., unpubl.). After the onset of Neo- glacial at ~2.1 ka (Mulvaney et al., 2012; Sterken et al., 2012; Cejka et al., 2020; Theilen et al., 2023), the James Ross Island land-terminating glaciers expanded only locally and formed the ice-cored moraines of the present day (Carrivick et al., 2012; Kaplan et al., 2020).

These palaeoreconstructions inform our understanding of the present
glacier dynamics in the context of changing climatic and marine con-
ditions in the Antarctic Peninsula region. The rapid grounding-line
retreat rates, hitherto unobserved via direct measurements (Dow-
deswell et al., 2020; Batchelor et al., 2023), may have occurred under
comparable, intense atmospheric and oceanic forcing as in the Croft
Trough. These examples might serve as analogues for recent and antic-
ipated intensification of extreme pressures on the Antarctic cryosphere,
be it in the form of i) strong föhn warming and surface melt of Larsen C
Ice Shelf via atmospheric rivers (Clem et al., 2022), ii) poleward
Southern Westerlies wind field shift causing advective warming of
subsurface coastal waters, in turn affecting retreat rates at glacier
grounding lines (Spence et al., 2014), or iii) more frequent upwelling of
(relatively warm) Circumpolar Deep Water to the continental shelf
around Antarctica as a response to higher meltwater input, with impli-
cations for the global oceanic overturning circulation (Li et al., 2023).

6. Conclusions

The TCN dating of erratic boulders in the James Ross Island archi-
pelago established the spatio-temporal pattern of the ice recession along
the two main palaeo-ice streams draining James Ross Island and Ant-
arctic Peninsula: Croft Trough and Prince Gustav Channel. The new
cosmogenic nuclide exposure data augment the existing deglaciation
chronologies and constrain the Holocene deglaciation in one of the most
opportune settings of the Antarctic Peninsula, thanks to the presence of
expansive deglaciated areas with distinct geology, and large altitudinal
gradients of the volcanogenic relief, enabling insights into the ice thin-
ning history. The approach of using paired $^{10}$Be-$^{26}$Al nuclides facilitated
an assessment of cosmogenic isotope inheritance and complex burial-
exposure history. Several erratics from the higher-altitude locations of
volcanic mesas may have been deposited preceding the Last Glacial
Maximum, thus rendering the exposure ages as maximum, rather than
minimum or mean, at least for the samples with low $^{26}$Al/$^{10}$Be ratio.

Nevertheless, the low scatter of TCN ages along the long axis of the Croft
Trough increases confidence that these mark deglaciation and suggests
the timing of transition from grounded ice to open marine conditions
occurred rapidly at ~8.6–7.2 ka. The Prince Gustav Channel Ice Stream
was rapidly thinning from at least ~12 ka, with subsequent separation of
the Antarctic Peninsula and James Ross Island ice masses by 10–8 ka.
With the Early Holocene Warm Period tailing off, the eustatic sea-level
rise (outpacing the glacio-isostatic rebound) and incursions of warmer
circumpolar waters were the primary factors of grounding line retreat
and transition to ice shelves, possibly accelerated in the case of Croft
Trough by hydrostatic response of ungrounding, near-fully-buoyant ice
stream to low-gradient topography. The examples of rapid deglaciation
from James Ross Island archipelago may serve as analogues for recent
and future intensification of extreme pressures on the Antarctic
cryosphere.

CRediT authorship contribution statement

**Matej Roman:** Writing – original draft, Visualization, Methodology, Investigation, Funding acquisition, Formal analysis, Conceptualization.
**Daniel Nývlt:** Writing – review & editing, Supervision, Investigation, Funding acquisition, Conceptualization. **Bethan J. Davies:** Writing – review & editing, Supervision, Investigation, Funding acquisition, Conceptualization. **Régis Braucher:** Writing – review & editing, Supervision, Methodology, Investigation. **Stephen J.A. Jennings:** Writing – review & editing, Investigation, Funding acquisition. **Michal Brežný:** Writing – review & editing, Investigation. **Neil F. Glasser:** Writing – review & editing, Supervision, Funding acquisition. **Michael J. Hambrey:** Writing – review & editing, Investigation. **Juan M. Lirio:** Writing – review & editing, Investigation. **Ángel Rodés:** Writing – review & editing, Methodology.

Declaration of competing interest

The authors declare that they have no known competing financial
interests or personal relationships that could have appeared to influence
the work reported in this paper.

Data availability

Data will be made available on request.

Acknowledgement

This work was supported by the Operational Programme Research, Development, and Education – Project Postdoc@MUNI (No. CZ.02.2.69/0.0/0.0/16_027/0008360) at Masaryk University, NERC grant to N. Glasser: “Glacial history of the NE Antarctic Peninsula Region over centennial to millennial timescales” (reference NE/F012942/1), and Czech Science Foundation (GAČR) grant No. 22–206210. The authors appreciate the use of the Czech Antarctic Research Programme infrastructure, supported by the MEYS projects VAN 2021/1, VAN 2020/1, LM2015078, CZ.02.2.69/0.0/0.0/16_027/0008360, and CZ.02.1.01/0.0/0.0/16_013/0001708, as well as the J. G. Mendel Station, and acknowledge its staff for all the support they provided. The ASTER AMS national facility (CEREGE, Aix-en-Provence) is supported by the INSU/CNRS, the ANR through the “Projets thématiques d’excellence” program for the “Equipements d’excellence” ASTER-CEREGE action and IRD. The authors would like to acknowledge the logistical support of the Dirección Nacional del Antártico and Instituto Antártico Argentino during the two field campaigns on Clearwater Mesa and Cape Lamb. The first author’s contribution was supported by JCMM Brno PhD Talent and MUNI GeoExcellence scholarships. We would like to thank two anonymous reviewers for their constructive comments, which greatly improved the original manuscript.
## Supplementary materials

Supplementary material associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2024.118803.

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