The impact of Holocene deglaciation and glacial dynamics on the landscapes and geomorphology of Potter Peninsula, King George Island (Isla 25 Mayo), NW Antarctic Peninsula

Abstract

The timing and impact of deglaciation and Holocene readvances on the terrestrial continental margins of the Antarctic Peninsula (AP) have been well-studied but are still debated. Potter Peninsula on King George Island (KGI) (Isla 25 de Mayo), South Shetland Islands (SSI), NW Antarctic Peninsula, has a detailed assemblage of glacial landforms and stratigraphic exposures for constraining deglacial landscape development and glacier readvances. We undertook new morphostratigraphic mapping of the deglaciated foreland of the Warszawa Icefield, an outlet of the Bellingshausen (Collins) Ice Cap on Potter Peninsula, using satellite imagery and new lithofacies recognition and interpretations, combined with new chronostratigraphic analysis of stratigraphic sections, lake sediments, and moraine deposits. Results show that the deglaciation on Potter Peninsula began before c. 8.2 ka. Around c. 7.0 ka, the Warszawa Icefield and the marine-facing Fourcade Glacier readvanced across Potter Peninsula and to the outer part of Potter Cove. Evidence of further readvances on Potter Peninsula was absent until the Warszawa Icefield margin was landward of its present position on three occasions: c. 1.7–1.4 ka, after c. 0.7 ka (most likely c. 0.5–0.1 ka), and by 1956 CE. The timing of Holocene deglaciation and glacier fluctuations on Potter Peninsula are broadly coeval with other glacier- and ice-free areas on the SSI and the northern AP and likely driven by interactions between millennial–centennial-scale changes in solar insolation and irradiance, the southern westerlies, and the Southern Annular Mode.

1 Introduction

The average annual temperature of the northern Antarctic Peninsula (AP) and South Shetland Islands (SSI) region (Figure 1A) has increased by almost 3°C (from −2.5°C) since the middle of the last century (Vaughan et al., 2003; Meredith and King, 2005; Turner et al., 2016), exceeding average global warming by ∼1.5°C–1.8°C (Sánchez-Lugo et al., 2018; Kaufman et al., 2020), leading to enhanced ice sheet thinning, glacier retreat, and permafrost thawing (Steig et al., 2009; Serrano et al., 2012; Bockheim et al., 2013; Osmanoğlu et al., 2013; Vaughan et al., 2013).

FIGURE 1

Past ice sheet, ice cap, and outlet glacier behavior provides an important long-term context for recent changes and allows for assessments of whether the current rates of change are beyond the range of natural variability (Bentley et al., 2014). During the global Last Glacial Maximum [g-LGM; 23–19 cal. ka BP, calibrated radiocarbon age in thousands of years before present (Hughes et al., 2013)], the AP ice sheet extended toward the shelf edge, and a secondary, smaller ice cap covered the northwest AP/SSI region, reaching the outer continental shelf ∼50 km north of its present location in the inner fjords of the Maxwell Bay, King George Island (KGI), SSI (Figures 1B, C) (John and Sugden, 1971; Curl, 1980; Ó Cofaigh et al., 2014; Nývlt et al., 2020). Deglaciation on KGI began broadly after 15.4 ± 2.5 ka (Seong et al., 2008), and Maxwell Bay was deglaciated between c. 14.8 and 14.1 cal. ka BP as the LGM ice sheet decoupled from the sea floor (Simms et al., 2011) and became ice-free by 9.1 cal. ka BP (Simms et al., 2011; Watcham et al., 2011). On land, lakes formed in deglaciated ice-free areas above the Holocene marine limit at ∼16 m above present mean sea-level (henceforth, m a.s.l.) on Fildes and Barton Peninsula provide minimum age constraints on deglaciation at, or before, c. 8.1 ka (Watcham et al., 2011; Oliva et al., 2019).

While the deglaciation of the AP and SSI by Termination 1 [the start of the Holocene at 11,700 years b2k—before CE 2000 (Gibbard and Head, 2020)] has been broadly established (Ó Cofaigh et al., 2014), more field data from glacier-free terrestrial areas of the northern AP continental margin are required to assess how natural warming during the transition into the current (Holocene) interglacial drove further local-regional deglaciation and how millennial–centennial-scale changes in climate during the Holocene impacted glacier dynamics (Spada et al., 2013; Bentley et al., 2014). Furthermore, an evaluation of the sensitivity of small ice caps and outlet glaciers of the AP and SSI to smaller-scale amplitude atmospheric and oceanographic forcing during the Holocene is of particular importance for constraining past rates of grounding-line retreat and relative sea level (RSL) change (Hall, 2007; Bentley et al., 2014; Johnson et al., 2022).

Changes in the behavior of the Bellingshausen (Collins) Ice Cap (BIC) on KGI /Isla 25 de Mayo (KGI), the largest island and ice cap in the SSI archipelago, can be constrained using field evidence from its deglaciated peninsulas, such as Potter Peninsula (Figures 1C, 2). Detailed geomorphological mapping and establishing robust chronologies for geomorphological landforms, stratigraphic sections, and lake sediments from glacier- and ice-free areas on these peninsulas can help refine existing models of AP and SSI continental margin deglaciation and allow us to examine the impacts and underlying mechanisms of post-LGM deglaciation and outlet/tidewater glacier readvance(s).

FIGURE 2

In common with other deglaciated peninsulas bordering Maxwell Bay (Oliva et al., 2019; Oliva et al., 2020) and across the northern AP (Ruiz-Fernández et al., 2019), Potter Peninsula is covered by till and other glacial, periglacial, proglacial, and paraglacial morphostratigraphic deposits, containing datable organic remnants (Sugden and John, 1973; Birkenmajer, 1998; Ballantyne, 2002; Ruiz-Fernández et al., 2019). Fourcade (1960), González-Ferrán and Katsui (1970), John and Sugden (1971), Sugden and John (1973), and Birkenmajer (1998) undertook pioneering work on Potter Peninsula, published geomorphological sketch maps, and reconstructed its glacial history, yet no comprehensive and up-to-date geomorphological synthesis exists. Moreover, in recent decades, the retreat of BIC glacier outlets has exposed new coastlines, landforms, and stratigraphic exposures, and a new assessment of the landform assemblages and their chronologies is now required.

We present a new geomorphological synthesis for Potter Peninsula comprising the first comprehensive geomorphological map constructed using satellite imagery and digital mapping techniques and new chronological data from sediment-landform assemblages recording glacier fluctuation on Potter Peninsula (Figures 2, 3). We test the hypothesis that Holocene deglaciation on Potter Peninsula from offshore LGM limits was a continuous process and combine all available evidence to answer the following questions: 1) When did the deglaciation of Potter Peninsula begin? 2) Did readvances from the Warszawa Icefield occur and what was their extent? 3) What were the wider regional–global drivers and impact of deglaciation and glacier readvance(s) on Potter Peninsula and the SSI?

FIGURE 3

To provide samples for dating, we excavated a new, deeper stratigraphic section in a similar location to del Valle et al. (2007), close to Mirounga Point on the outer Peninsula (nPII in Figure 2, Location 1 in Figure 3). We also excavated a river section on Potter Peninsula that cuts through the same moraine sampled by Sugden and John (1973) (PC in Figure 2; Location 2 in Figure 3) and a new “Inland outcrop” ∼700 m from the active glacier front, recently exposed (2013 CE) by the deglaciation of the glacier foreland (Figure 2, Location 3 in Figure 3). Results from these records were compared with ³He cosmogenic nuclide exposure dating of boulders from the Three Brothers Hill moraine system and data from new lake sediment records extracted from Matias Lake (ML in Figure 2), located outside the 1956 CE glacier limit of the Warszawa Ice Field, and Lake L15 (L15 in Figures 2, 3; Supplementary Figure S2), the furthest permanent water body from the active glacier front of the Warszawa Icefield.

2 Geology, climate, and glaciological setting of Potter Peninsula

2.1 Study area and geology

Potter Cove is an inner fjord of Maxwell Bay bordering the Fourcade Glacier, a marine-terminating glacier located on the northeastern side of the fjord (Figure 1C). The Warszawa Icefield to the south of the Fourcade Glacier defines the eastern border of the deglaciated Potter Peninsula (Figure 1C). Both glaciers are outlets of the BIC (Collins Ice Cap) that covers approximately 90% of the surface of KGI, to a maximum thickness of about 400 m (Blindow et al., 2010). Aerial and satellite imagery shows that the Warszawa Icefield and the Fourcade Glacier retreated rapidly between 1956 and 2013 CE, with ∼40% of the present-day deglaciated area of Potter Peninsula becoming ice-free during that time (Figure 3).

Lava flows, pyroclastic rocks, and related hypabyssal intrusives are exposed on Potter Peninsula, and the basalts of the peninsula have been potassium-argon dated at 49 ± 1 to 42 ± 1 Ma (Smellie et al., 1984; Kraus and del Valle, 2008a, b). Three Brothers Hill (196 m a.s.l.), an andesitic plug with transverse faults and columnar jointing formed during the final stages of a Palaeogene stratovolcano, is the most prominent topographic feature on Potter Peninsula (Figures 1D, E) (Birkenmajer 1998).

2.2 Climate

The mean annual air temperature on KGI between 1948 and 2011 CE was −2.5°C, with Austral summer temperatures often above zero (Kejna et al., 2013). A warming trend observed since the 1950s has paused, with a slight cooling occurring between 1999 and 2015 CE (Oliva et al., 2017). In contrast to the high interior atmospheric pressure of the West and East Antarctic plateau, the climate of the northern AP and SSI is influenced by the mid-latitude and/or the polar atmospheric cells (Bentley et al., 2009). Atmospheric cells and the Southern Hemisphere Westerly winds (SHW) bring warm and moist air masses formed at lower latitudes over the surrounding oceans and ∼1,000 mm yr⁻¹ of precipitation (Turner et al., 2002; Bentley et al., 2009; Falk et al., 2018), which drives snow accumulation, glaciofluvial erosion, and surface- and groundwater discharge into Potter Cove (Falk and Silva-Busso, 2020). Potter Peninsula becomes largely snow-free during the Austral spring/summer, but ephemeral snow patches persist well into late summer. The wind flow in Potter Cove is predominantly from SW to NW, with a mean annual wind speed of ∼15 m s⁻¹ and weaker katabatic component (Falk et al., 2018).

Between 1996 and 2005 CE, increased cyclonic activity in the Drake Passage, associated with more poleward-focused SHW and an increasingly more positive phase of Southern Annular Mode (SAM), drove sea ice poleward, increasing the advection of warm air across southern Patagonia, the northern AP, and SSI during winter (Kaplan et al., 2016, 2020; Marshall et al., 2017; Oliva et al., 2017; Reynhout et al., 2019). Conversely, between 2006 and 2015 CE, colder winters pushed sea ice north, increasing snowfall across the northern AP and doubling snow cover on some SSI islands between 2009 and 2014 CE (de Pablo et al., 2013; Oliva et al., 2017).

The SAM is the primary mode of interannual and decadal to centennial-scale atmospheric circulation variability in the Southern Ocean and reflects the zonal mean sea level pressure difference between Antarctica (>65°S) and the mid-latitudes (40°S) (Marshall, 2007; Moreno et al., 2018; Kaplan et al., 2020). The SAM index is a measure of the longitudinal mean SHW strength, with positive phases characterized by stronger and poleward-shifted SHW (Marshall, 2007; Charman et al., 2018; Kaplan et al., 2020). Negative phases are characterized by generally colder and, in some places, more humid conditions and weaker (and/or more diffuse), possibly equatorward-shifted, SHW (Marshall, 2007).

Interactions between the SAM and a deepening Amundsen Sea low during positive SAM phases can also bring warmer air from the Atlantic and north-easterly winds with cooler air from the Weddell Sea gyre, leading to colder/wetter conditions on the northern AP and low-lying SSI and increased glacier mass (Liu et al., 2005; Thomas et al., 2008; Goodwin et al., 2016; Oliva et al., 2017; Falk et al., 2018; Kaplan et al., 2020). These processes are intensified further in the Pacific sector that feeds the AP by interactions between SAM and the (sub)decadal-scale El Niño Southern Oscillation (ENSO). Positive (enhanced) phases of ENSO (El Niño) drive warmer interannual global mean temperature variability and negative phases of the SAM (Wang and Cai, 2013; Turner et al., 2016). Decadal scale variability in temperature measured at stations on the AP and SSI has been linked to changes in ENSO (Oliva et al., 2017).

2.3 Glaciological setting

Observations, numerical simulations, and ground-penetrating radar data indicate that the BIC is a partly temperate ice cap, with cold ice at higher elevations and temperate ice at lower elevations (Blindow et al., 2010; Rückamp et al., 2010, 2011). The Warszawa Icefield margin is a polythermal, debris-covered glacier, and ice-cored and thrusted moraines can be found at Stranger Point (Figure 2A). Increased precipitation on the western AP over the last 5 decades (Thomas et al., 2008; Goodwin et al., 2016) has resulted in a slightly positive glacier mass balance on northern AP islands and the SSI overall (Oliva et al., 2017). From 1970 to 1990 CE, the BIC had a stable to slightly negative glacier mass balance and an equilibrium line altitude (ELA) at ∼ 150 m a.s.l. (Curl, 1980; Turner et al., 2002; Bentley et al., 2009; Falk et al., 2018). By 2015 CE, the ELA had risen to ∼260 m a.s.l., and glacier flow rates had accelerated (Falk et al., 2018). Across the SSI, permafrost is discontinuous between 0 and 30 m a.s.l. Colder conditions between ∼30 and 50 m a.s.l. favor the formation of continuous permafrost. Above 50 m a.s.l., permafrost dominates periglacial processes, but permafrost is less common above 90 m a.s.l. (Serrano et al., 2012; Oliva et al., 2020).

The deglaciation of Potter Peninsula left behind numerous glacial landforms. Inland moraines are surrounded by a coastal environment with older (Holocene-age) raised beaches and marine deposits (John and Sugden, 1971; Lindhorst and Schutter, 2014; Heredia Barión et al., 2019). Holocene RSL curves for the SSI are more complex than in other locations on the AP, and several different curves have been constructed (Pallàs et al., 1997; Bentley et al., 2005; Hall, 2010; Roberts et al., 2011; Watcham et al., 2011; Simms et al., 2012, 2021; Johnson et al., 2022). Bentley et al. (2005) (and Johnson et al. (2022)) proposed an early-mid Holocene marine limit of ∼16–18 m a.s.l. for the SSI. Based on the ages of marine-freshwater transitions and the altitude of isolation basins on nearby Fildes Peninsula, Watcham et al. (2011) built on the pioneering work of Mäusbacher et al. (1989) and refined this to a more conservative estimate of ∼16 m a.s.l. On Potter Peninsula, del Valle et al. (2002) constrained the onset of the deglaciation and raised beach formation between 17.1 and 14.4 m a.s.l. at Stranger Point (Pingfo-I, PI in Figure 2) with a minimum age of 5.5–5.4 cal. ka BP from elevated coastal sediments, although Fretwell et al. (2010) and Lindhorst and Schutter (2014) later revised the maximum raised beach altitude on Potter Peninsula to 12.2 m a.s.l.

There are several low-altitude sedimentary sections close to the present-day sea level along the south side of Potter Cove are ∼200 m west of the Argentinian Carlini station, ∼2 km from the front of the modern glacier terminus, and ∼1 km from its 1956 CE position (Figures 2, 3). They are composed of marine intertidal sands with remnants of algae and marine mollusks, bivalves, and the bones of penguins, seals, and cetaceans, deposited in an ice-free environment, and capped by a till that represents the last time a local glacier readvanced over the outer part of Potter Cove/Peninsula (Sugden and John, 1973).

Radiocarbon dating of the deepest buried shells and seaweed in these sections provides a maximum age constraint for deglaciation (Strelin et al., 2014). By analogy with observations from modern beaches on Potter Peninsula, seaweed was probably deposited as large trash, subfossil deposits close to the high tide level (Fretwell et al., 2010; Strelin et al., 2014). Sugden and John (1973) first dated shells from a 2.5 m thick stratigraphic section of the “Potter Cove section” (located close to PC in Figure 2) and proposed the oldest age for the onset of the deglaciation and marine transgression from the outer part of Potter Peninsula of 9,670 ± 230 ¹⁴C years, or 9,570 ± 650 cal. a BP (S&J Birm-48a in Table 2, recalibrated in this study). Del Valle et al. (2007) used these ages to constrain the age of a stratigraphic profile named “Pingfo II” (Location 1 in Figure 3B) 60 m to the west in the same moraine, but the ages obtained by Sugden and John (1973) have not been replicated.

3 Methods

3.1 Geomorphological mapping

Detailed geomorphological mapping is an essential first step in assessing the deglacial history of continental margins around Antarctica (Figures 2, 3). Using a glacial landsystems approach (e.g., Davies et al., 2013), we distinguished landforms and sediment exposures associated with glacier advance and retreat to understand how the landscape on Potter Peninsula evolved during the Holocene (Table 1). Mapping was undertaken in ARC-GIS v.2.5, with final layouts achieved in Adobe Illustrator v. 26.2.1 or CorelDRAW v. 2020.

3.2 Stratigraphic profile, sampling, and lithofacies analysis

Morphostratigraphic descriptions and sampling for radiocarbon dating were undertaken at an outer peninsula stratigraphic profile site excavated in a similar location to the “Pingfo II” profile (del Valle et al., 2007) (Table 2). We refer to the new profile as “new Pingfo II” (nPII in Figure 2 and Location 1 in Figure 3). We also excavated, described, and radiocarbon dated a new river section adjacent to the “Potter Cove section” sampled by Sugden and John (1973). Henceforth, we refer to this as the “new Potter Cove section.” Both sections are located at “PC” in Figure 2 and Location 2 in Figure 3. Following recommendations in Björck et al. (1991c), we also radiocarbon dated terrestrial moss samples embedded in a recently exposed “Inland outcrop” ∼700 m from the active glacier front and inside the 1956 CE limit to constrain glacier readvance(s) on the inner peninsula area (Figure 2 and Location 3 in Figure 3B). Depositional units in the stratigraphic sections were characterized using lithofacies textural (grain size and fabric), composition, sedimentary structures, and stratigraphic relationship criteria (Supplementary Material for additional methods). Summary results were digitized in Figures 5, 6 using SedLog 3.0 (Zervas et al., 2009), and grain size data are shown in Supplementary Table S1.

3.3 Lake sediment analysis

After mapping, classifying, and assessing the suitability of the numerous lakes on Potter Peninsula for sediment coring (Supplementary Figure S2), we cored and present here summary results from two foreland lake basins: Matias Lake (ML in Figures 2, 3B; Lake L5 in Supplementary Figure S2); and Lake L15 (aka GPS Lake).

Matias Lake (Lake L5, Figures 2, 3B; 62.2450°S, 58.6655°W) is a small lake sub-basin up to ∼6 m deep (del Valle et al., 2004) formed in the moraine deposits of the glacier foreland, adjacent and connected to Rudy Lake (Lake L6) in the mid-foreland area (Figure 2; Supplementary Figure S2). Located on the eastern, glacier-facing side of Potter Peninsula at the base of, and shaded by Three Brothers Hill, lake-ice often persists into January or February. In contrast, Rudy Lake and Lake L15 are largely free of lake ice by November–December. The southern basin, Rudy Lake, is a larger but shallower basin, ∼4 m at its deepest point (del Valle et al., 2004). Both lakes (L5 and L6) are located outside of the 1956 CE glacier limit of the Warszawa Ice Field (Figures 1, 2), ∼500 m NW of the “Inland outcrop.” Following a ground penetrating radar (GPR) survey to locate the depocentre in Matias Lake, 13 sediment cores were extracted from a solid lake-ice platform in November 2011 using a 50 mm wide, 1 m long Livingston corer and a 50 mm wide, 0.5 m long Russian corer.

Lake L15 (62.24056°S S, 58.67757°W) is a small shallow bedrock basin (<1–2 m deep) inland from the new Pingfo II section (L15 in Figures 2, 3). It is a seasonally lake-ice-free basin and the furthest permanent water body from the active glacier front on Potter Peninsula (Figures 1, 2; Supplementary Figure S2). Eight cores were taken from a lake-ice platform over the depocentre using a 50 mm wide, 1 m long Livingston corer and a 50 mm wide, 0.5 m long Russian corer.

Livingston cores were split, and all cores were described in the laboratory. Non-destructive ITRAX (Cox Analytical) X-ray fluorescence (XRF) core scanning (XRF-CS) and Bartington Magnetic Susceptibility High-Resolution Surface Scanning Sensor (MS2E) measurements were undertaken at Aberystwyth University and on a Geotek® multi-sensor core logger (MSCL) at Durham University following standard procedures (Gunn and Best, 1998; Davies et al., 2015; Roberts et al., 2022). Contiguous bulk, wet sediment geochemical Energy Dispersive XRF-CS (Energy Dispersive Spectroscopy) data were obtained using a chromium (Cr) X-ray tube (X-radiography image settings: 40 kV, 40 mA, 200 m; XRF-CSCr settings: 30 kV, 40 mA, dwell time of 10 s, at 100 μm (MAT1) and 2 mm (L15-H2, L15-H16) intervals). Machine and sample calibration were undertaken using a synthetic glass standard and XRF fused glass discs of sediment samples from the nearby Yanou Lake and Ardley Lake records on Fildes Peninsula at the start and end of each analytical phase. Raw count per second (cps) data were analyzed using the Cox Analytical Q-spec software.

To account for downcore variations in count rate, density, water, and organic content, XRF-CS data are presented as relative changes in percentages of the total scatter normalized ratio sum (%TSN, equivalent to the %cps sum—Saunders et al., 2018; Roberts et al., 2022) and as natural log ratios and centered log ratios (clr). Log ratios have been shown to produce similar downcore patterns to more traditional and fully quantitative Wavelength Dispersive Spectroscopy, WDS-XRF, and dry subsample analysis (Kylander et al., 2011; Davies et al., 2015; Roberts et al., 2017; Dunlea et al., 2020). Data were filtered to remove the small number of downcore spectra with kcps less than the mean minus two-SD, caused mainly by small gaps in the core, and MSE values greater than mean plus four-SD, which represents a poor fit of the measured to theoretical energy spectra. “Measurable” elements were determined by autocorrelation, with significant zones, and their boundary positions, defined by constrained cluster analysis (CONISS) with broken stick analysis applied to square-root-transformed %TSN values of key elements using the R packages Vegan and Rioja (Juggins, 2012; Oksanen, 2014).

Hyperspectral image (HSI) scanning analysis, which measures reflected optical properties between wavelengths from 400 to 1,000 nm (Butz et al., 2015), was undertaken at the University of Bern using the Specim Ltd. single core scanner (PFD-CL-65-V10E line scan camera, 400–1,000 nm) following the protocol of Butz et al. (2015). The spatial resolution (pixel size) was set at 69 × 69 μm, and the spectral resolution is 2.8 nm sampled at an interval of 0.78 nm. Raw data were normalized with a BaSO4 reference, and spectral endmembers were calculated using the software ENVI 5.03. Quantitative estimates of pigments were obtained using the Relative Absorption Band Depth (RABD) method, which uses ratios and normalized reflectance data from distinct wavelengths. The spectral index RABD660;670 (RABD at 660–670 nm) was calculated from the continuum between 590 and 730 nm (Butz et al., 2015) with I-band wavelengths between 660 and 670 nm using equations RABD660;670 = (6*R590+7*R730)/13/Rmin(660;670) and RABD660;670 [I-Band] = (6*R590+7*R730)/13/Rmin(660;670)/Rmean (Rein and Sirocko, 2002). The spectral index RABD660;670 is proportional to the concentration of sedimentary chlorophyll-a, chlorophyll-b, and related colored diagenetic pheopigments (referred to as total chlorophylls TChl), which are generally interpreted as aquatic primary production indicators (Leavitt and Hodgson, 2001; Butz et al., 2017; Schneider et al., 2018; Makri et al., 2020).

3.4 Chronology and chronostratigraphy

3.4.1 Radiocarbon dating

Twenty-one accelerator mass spectrometry (AMS) radiocarbon (¹⁴C) ages were obtained from seaweed, marine mollusc shells, penguin bones, bones from undetermined species, remnants of terrestrial mosses embedded in marine coastal proximal and terrestrial moraine sediments, and freshwater subaquatic moss in the uppermost unit (0–6 cm depth) of the Lake L15-H16 (Figure 3E; Supplementary Figure S6; Table 2). In order to reduce the risk of contamination with modern material, samples for radiocarbon dating from the stratigraphic sections were taken from freshly cleaned outcrops and stored in zip-lock plastic bags at 4°C. Samples were prepared at the Alfred Wegener Institute Helmholtz Centre for Polar and Marine Research, Durham University, and the British Antarctic Survey. AMS and ¹³C/¹²C isotope ratio measurements undertaken at ETH Zürich, the NERC Radiocarbon Laboratory, East Kilbride, and by Beta Analytical, Miami, were used to calculate conventional radiocarbon ages following established procedures (¹⁴C years; Table 2; Supplementary Material for details).

Calibration of marine sample radiocarbon ages (marine shells, penguin bones, bones from undetermined species, and seaweed) was undertaken in Oxcal v. 4.4 using the Marine20 calibration curve (Gordon and Harkness, 1992; Bronk Ramsey, 2009; Heaton et al., 2020, 2022). We used a newly recalculated local marine reservoir age offset (ΔR) of 666 ± 76 ¹⁴C years that represents the weighted mean ΔR of the four closest and most applicable radiocarbon-dated marine samples collected before 1950 CE from the northern AP and Signy Island in the online Marine20 database (http://calib.org/marine/) (Table 2; Supplementary Material for details). Terrestrial and freshwater subaquatic moss samples were calibrated using the Southern Hemisphere SHCal20 (Hogg et al., 2020) calibration curve in Oxcal v. 4.4 (Bronk Ramsey, 2009). Post-bomb (>1950 CE) ages were corrected according to ¹³C/¹²C isotopic ratios from measured pMC with the “present day” pMC value defined as 107.5% (2010 CE) and calibrated using the SHCal13 SH Zone 1-2 Bomb curve in CALIBomb (Reimer and Reimer, 2004; Hua et al., 2013). Median and mean calibrated ages and their 2σ (95.4% confidence level) maximum to minimum error ranges are presented as thousands of years before the present (cal. ka BP). Calibrated ages have been rounded to the nearest 10 years in Table 2 and to the nearest hundred years in the text to reflect realistic total (i.e., all internal and external) uncertainties. Post-bomb ages are presented as year CE and have been rounded to the nearest year.

Bayesian age-depth models for the new Pingfo II sedimentary sequence and the lake sediment records were constructed using the Bacon R package (Blaauw and Christen, 2011) (Figure 6B; Supplementary Figure S5). Summed probability density analysis of the new Potter Peninsula radiocarbon data was undertaken using the Rcarbon package v. 1.4.2 (Bevan, 2021) and compared to non-parametric phase modeling (i.e., a probabilistic version of the Oxcal SUM command) using the Bchron v. 4.7.6 R package (Haslett and Parnell, 2008; Parnell, 2021). New radiocarbon data from Potter Peninsula were then incorporated into a new compilation of existing maximum age constraints on glacier readvance from across KGI, and non-parametric phase modeling was undertaken (Figures 6C, D).

3.4.2 Lead-210 (²¹⁰Pb), caesium-137 (¹³⁷Cs), and americium-241 (²⁴¹Am) analysis

²¹⁰Pb 137Cs and ²⁴¹Am dating of the uppermost 10 cm of the lake sediment records was undertaken using ∼4 g of homogenized dried sediment, added into tubes to a predefined level and sealed gas tight. After at least 21 days of storage to obtain radioactive equilibrium between ²²⁶Ra and ²²²Rn, activities of radionuclides were measured by well-type gamma spectrometry (Ge-detector, GWC 2522-7500 SL, Canberra Industries Inc., United States) and processed with GENIE 2000 3.0 (Canberra Industries Inc., United States). Data analysis and dating model calculations followed standard procedures defined by Appleby and Oldfield (1978). Lead-210 age estimates were derived using the constant rate of supply (CRS) method (Appleby and Oldfield, 1978) and incorporated into Bayesian age-depth models (Figure 6B; Supplementary Figure S5) (Supplementary Material for details).

3.4.3 Cosmogenic helium-3 (³He) nuclide surface exposure dating

Stable cosmogenic ³He accumulates in pyroxene and olivine phenocrysts in the upper few cm of basaltic boulders on the Three Brothers Hill moraine system, recording the total time that the rock surface has been exposed to cosmic rays. Using a hammer and chisel to remove the upper few centimeters of exposed surfaces, we collected five samples from boulders perched on moraines between 35 and 100 m a.s.l. (Supplementary Figure S1). The sampled boulders were >50 cm in diameter and the most suited for cosmogenic dating analysis as they showed no signs of sediment cover or significant erosion. Their size and shape meant it was very unlikely they had been overturned, and post-depositional movement was minimized by sampling boulders on locally flat and stable surfaces, away from steep slopes and cliffs (Supplementary Figure S1). Differential GPS (dGPS) measurements were undertaken using a Trimble Pathfinder ProXH to determine the precise location and altitude of boulders in relation to the DALL 66019M002 (S62°14′16.335″, W58°39′52.364″, ellipsoidal height 39.376 m) triangulation station located on the Argentine Carlini base, a few hundred meters away from the sampled erratics (Table 3). dGPS precision is better than 10 cm in all axes, but ellipsoid correction errors are larger.

Sampling and processing methods are described in Tables 3–5 and in detail in the Supplementary Material. Exposure ages were calculated using the CRONUScalc calculator (Version 2.0; Marrero et al., 2016) with the time-dependent Lal (1991)/Stone (2000) scaling model (Lm) for altitude at Antarctic pressure conditions and the primary calibration dataset for ³He in pyroxene, which yields a long term sea-level high latitude (SLHL) scaled production rate of 122 ± 13 at g⁻¹ a⁻¹ (Borchers et al., 2016). We report internal and external uncertainties in Table 4. External age uncertainties include production rate uncertainties, and exposure ages determined with other scaling models (e.g., Lifton et al., 2014) vary by up to ∼6%. Following Balco et al. (2008), external uncertainties are used for comparison with calibrated AMS radiocarbon ages and error ranges.

All statistical analysis was undertaken, and figures were constructed using R v. 4.1.0/RStudio v. 1.4.1717 and packages Tidyverse, ggplot2, Vegan, Rioja, Ggally v. 2.1.2, RBacon, Rcarbon, Bchron, and Sigmaplot v. 14.0, C2 (Juggins, 2007), with final figure layouts achieved in Adobe Illustrator v. 26.2.1 or CorelDRAW v. 2020.

4 Results

We identified four distinct altitudinal and glacially impacted environments on Potter Peninsula: 0–15 m a.s.l. (Holocene-age and active beaches); 15–50 m a.s.l. (vegetated (perma)frost-affected bedrock plateaus); 50–100 m a.s.l. (fresh-looking, recent glacial deposits); >100 m a.s.l. (Three Brothers Hill) (Figure 2).

The main features of these landform assemblages are summarized and interpreted in Table 1 and illustrated in Figures 3–5. In the following sections, we have organized results obtained from the stratigraphic sections, lake records, and chronological analysis according to the landform assemblages described above.

FIGURE 4

FIGURE 5

4.1 Landform assemblage LA-1: Active glacial sediment landforms

4.1.1 Inland Outcrop on the recently deglaciated inner Warszawa Icefield foreland

The contact between the Warszawa Icefield and its proglacial area is shallow and contains many active glacial sedimentary landforms (Figure 4A; Table 1; Section 1 for further details). In 2013, we discovered a new exposure ∼700 m away from the present-day glacier limit (62.2473°S, 58.6466°W). This “Inland Outcrop” is a partly degraded, non-vegetated, 2.3 m thick massive matrix-supported sandy diamicton. Two radiocarbon ages from basal moss fragments incorporated into the base of a diamicton have mean ± 2σ ages of 1.7 ± 0.1 and 1.8 ± 0.1 cal. ka BP, with a density phase model 95% probability range of 2.2–1.4 cal. ka BP (Figures 5, 6; Table 2, nos. 26, 27). This section represents glacier bulldozing of moss that grew on previously glacier-free areas, and the youngest of the moss radiocarbon ages provides a maximum age constraint on glacier readvance across the mid-upper glacier foreland.

FIGURE 6

4.2 Landform assemblage LA-2: Inactive glacial assemblage

4.2.1 Stratigraphic river sections from northern mid-outer Potter Peninsula

The new Pingfo II (nPII) section (62.2389°S, 58.6729°W; Figure 2A and Location 1 in Figure 3) and “Potter Cove’’ (PC) river sections (62.23895°S, 58.67113°W; Figure 2 and Location 2 in Figure 3) are part of a small lateral moraine ridge running obliquely to the southern coastline of Potter Cove.

The nPII section is composed of five units of shell-bearing marine sand, and terrestrial sand, sand gravel and gravel (Lithofacies 1–5 in Figure 6A) overlain by a sandy diamicton unit composed of matrix-supported pebble to cobble-sized clasts (Lithofacies 6; Figure 6A) as follows: Lithofacies 1 is a laminated silt/very fine sand with remnants of shells of bivalve Laternula elliptica and algae; Lithofacies 2 are composed of beds of massive, matrix to clast supported gravels, ∼10–15 cm thick, with subangular to angular clasts up to 5 cm embedded in a sandy matrix; Lithofacies 3 is composed of horizontally laminated very fine to silty sand, with brownish-to-orange algae and shelly remains with sporadic subangular clasts up to 7 cm; Lithofacies 4 is composed of horizontally laminated fine-to-medium sand, with silty-clay and sporadic sub-angular clasts up to 10 cm, shells of Laternula elliptica, penguin bones, and a distinctively compacted brownish (algae) deposit at 1.60 m covering fine pebbles; Lithofacies 5 is a stratified matrix and clast-supported gravel intercalated with sand and laminated silt, subrounded and subangular clasts up to 10 cm and larger “dropstones” up to 30 cm diameter and has a contorted upper boundary; Lithofacies 6, which is a sandy-clay, matrix-supported, clast rich, massive to weakly laminated diamicton, with striated angular to sub-angular clasts up to 50 cm. The upper surface of the nPII and PC sections is a flat terrace at 6 m a.s.l. littered with larger clasts on the surface facing Potter Cove and forms part of the inactive vegetated glacial assemblage (Figures 3, 6A; Table 1).

Twenty-one samples analyzed for AMS radiocarbon dating included remains of different marine organisms, seaweed, with Laternula elliptica shells and bones (Figure 6A; Table 2, nos. 1–21). Their mean and median calibrated radiocarbon ages cluster between 7.7 and 7.2 cal. ka BP, with a density phase model 95% probability range of 8.2–7.0 cal. ka BP. Shells in the lowermost sand unit are small fragments and unlikely to be in situ but also imply that the intertidal sand was accumulating in the range of c. 8.2–6.9 cal. ka BP (2σ calibrated age range). All new ages obtained from the nPII section are distinct (at 95% probability) from older radiocarbon ages of similar marine shells measured by Sugden and John (1973) with a density phase model 95% probability range of 10.1–9.4 cal. ka BP (Figure 6C; Table 2). Bayesian age-depth modeling shows that deposition of Lithofacies 1–4 was broadly in stratigraphic order, with a homogenous sedimentation rate of ∼10 mm a-1 (Figure 6B).

Four new AMS radiocarbon ages from marine shells, and a ∼2–10 mm thick filamentous seaweed (brown algae) layer collected from between the marine sand unit and the overlying diamicton in the nearby PC river section (Table 2, nos. 22–25; Figure 6C), were dated independently and had mean calibrated ages of c. 7.5 ± 0.2 cal. ka BP, respectively (2σ range 7.7–7.0 cal. ka BP) (Table 2). Their density phase model 95% probability age range is in close agreement with new ages from the nPII section (Figures 6B, C).

4.2.2 Lake sediment records from the central mid-outer area of Potter Peninsula

Lake L5 (Matias Lake): we extracted 13 short cores from the depocentre in Matias Lake and along a surface transect toward Rudy Lake. Sediment recovery depth ranged 20–60 cm before encountering an impenetrable diamicton layer. Cores MAT1 (L5-H1) (27 cm) and MAT2 (L5-H2) (29 cm) in Figure 7 were extracted from the deepest part of Matias Lake (5.8 m; 62° 14′ 42.054″S, 58° 39′ 53.82″W; Figure 2; Supplementary Figure S2). Two lithological zones were present across all cores. The basal zone 1 diamicton was predominantly matrix-supported, largely impenetrable, and only retained in some cores. The overlying sediments in zone 2 are composed of fine (mm scale) grey silty clay and Fe-rich orange clay laminations, probably reflecting periodic (possibly even seasonal) overturning and/or oxidation processes within the lake and glaciolacustrine deposition following the retreat of the Warszawa Icefield landward from Matias Lake and Rudy Lake (Figure 7B).

FIGURE 7

Obtaining basal radiocarbon ages from basal and bulk minerogenic sediments in the Matias Lake cores proved challenging due to a general lack of organic carbon (Figure 7; Supplementary Figures S5, S6; Table 2). The ²¹⁰Pb CRS age model shows that the uppermost 10 cm have been deposited since c. 1850 CE, and the well-defined ¹³⁷Cs peak at 5 cm depth is coherent with the ²¹⁰Pb age model (Figure 7B; Supplementary Table S2). A 0.5 cm thick tephra deposit between 8 and 9 cm depth (1908 ± 10 CE) is characterized by elevated Ca and Sr and most likely represents an early C20th eruption from Deception Island, 120 km southwest of Potter Peninsula.

Lake L15 (GPS Lake): we extracted eight cores from a grid of 25 (∼1 m spaced) holes drilled through the ∼75–100 cm thick lake ice above the deepest accessible point (∼2 m) at Lake L15. Core L15-H2 (62° 14′ 26.052″S, 58° 40′ 39.36″W) in Figure 7A contained 30 cm of sediment, L15-H4 11 cm, L15-H9 39 cm, L15-H16 55 cm (62° 14′ 26.016″S, 58° 40′ 39.252″W; Figure 7A), L15-H17 45 cm, L15-H19 34 cm, L15-H20 61 cm, and L15-H22 49 cm. The same three lithological zones exist in all records. Zone 1 is a fine silt/clay matrix-supported diamicton, zone 2 is a fine silt/clay laminated sediment, and zone 3 is a freshwater subaquatic moss. Hyperspectral and XRF-CS profiles show that zones 1 and 2 have very low primary productivity (Figure 7A), with the transition to elevated RABD660;670 values into zone 3 reflecting increased active photosynthesis by subaquatic moss and higher subaquatic primary production. zone 3A shows a maximum concentration of TChl, which remains high in zone 3B, but with relatively lower TChl concentrations. This relative decrease might also be related to higher lithogenic sedimentation rates and/or greater porosity.

Radiocarbon ages of freshwater moss and macrophytic material from the uppermost lithological Unit 3 in Lake L15 range from post-1950 CE to c. 0.6 cal. ka BP and are statistically distinct from ages obtained from terrestrial moss layers embedded in the “Inland Outcrop” (Figures 2A, 3B, 7; Supplementary Figure S6; Table 2). The Lake L15 ²¹⁰Pb Constant Rate of Supply (CRS) age model has a low sedimentation rate in the top 10 cm, and the sample at ∼4–5 cm is at least 150 years old (Supplementary Table S3). Although the prerequisite for the CRS model was not fulfilled because the flux of ²¹⁰Pb has changed through time and is not constant, the ²¹⁰Pb data are consistent with the radiocarbon ages that show sediment between 6 and 6.5 cm dates to 620 ± 80 cal. a BP, and 3–3.5 cm and 0–0.5 cm depth were deposited in the “post-bomb” (post-1950 CE) era, most likely between −40 and −44 cal yr BP (1990–1994 CE) (Figure 7; Table 2; Supplementary Table S5). The ¹³⁷Cs data are inconsistent with the ²¹⁰Pb CRS age model (Supplementary Table S3), and the steep increase in ¹³⁷Cs in the upper 2 cm may relate to a ¹³⁷Cs “soil reservoir,” which is leaching ¹³⁷Cs from the catchment into the lake from snow or lake-ice melting.

4.3 Landform assemblages LA-2, LA-3, and LA-4

4.3.1 Stranger point and three brothers hill moraine systems and associated proglacial/periglacial landforms and deposits

Well-preserved morainic crests and glacial erratics are rare in the Three Brothers Hill moraine system mainly because the moraines have been heavily degraded into debris flows and are now covered in talus rock accumulations (Table 1; Figure 8; Supplementary Figure S1).

FIGURE 8

We sampled erratics firmly embedded on the moraine for He-3 cosmogenic nuclide exposure dating to minimize the effect of post-depositional movement. In summary, ages are scattered between 9.4 ± 2.6 ka and 41.2 ± 5.6 ka (Figure 4K; Supplementary Figure S1; Table 4) and have external uncertainties up to 28% (e.g., sample 01-Potter, Table 4). The two youngest ages from boulders 01-Potter (9.4 ± 2.6 ka) and 05-Potter (18.6 ± 2.9 ka) had striated surfaces indicative of glacial transport and in situ preservation (Table 4). The large age scatter in the ³He cosmogenic nuclide exposure ages and the large individual age errors are due to the low concentration of cosmogenic helium and/or the poorly known ³He/⁴He ratio of the non-cosmogenic component. The possibility of previous exposure and complex and poorly understood nuclide inheritance processes cannot be excluded either. Therefore, we consider the ³He cosmogenic nuclide exposure results to be promising but tentative.

5 Discussion

We examine the new evidence for these stages in detail in the following sections (Sections 5.1–Sections 5.5) in the context of regional environmental change across the SSI and northern AP during the Holocene and suggest possible mechanisms for deglaciation and glacier readvances (Section 5.6).

5.1 Early Holocene terrestrial deglaciation

Ice retreat on land after deglaciation from offshore LGM limits across the SSI occurred c. 11–8 cal. ka BP, constrained by ages obtained from glacially polished bedrock surfaces on the Barton and Weaver peninsulas, cosmogenic nuclide exposure ages, and the onset of lake sedimentation on Barton Peninsula, Fildes Peninsulas (KGI), and Byers Peninsula (Livingston Island) (Björck et al., 1996; Seong et al., 2008; Watcham et al., 2011; Toro et al., 2013; Oliva et al., 2016a, b; Nývlt et al., 2020).

Our tentative minimum ³He cosmogenic nuclide exposure deglaciation age of 9.4 ± 2.6 ka from Three Brothers Hill on Potter Peninsula (Figure 4K) is broadly consistent with the high-altitude deglaciation history of KGI (Seong et al., 2008), and only one of the radiocarbon ages was obtained by Sugden and John (1973) (S&J Birm-48a in Table 2: 9,670 ± 230 ¹⁴C years; 9,570 ± 650 cal. a BP), which is clearly an outlier (Figures 6B–D, 9B, D).

FIGURE 9

The interpretation of lithofacies assemblages in the nPII and PC stratigraphic sections underpins our revised deglaciation history for Potter Peninsula. Both sections have a lithofacies association typical of polythermal tidewater glaciers (Stewart et al., 1991; Bennett et al., 1999). Lithofacies 1, 3, and 4 were deposited in a distal glaciomarine environment isolated from wave action, with the occasional release of ice-rafted debris. Lithofacies 2 is a dense debris flow entering the glaciomarine environment. Escape structures made by Laternula elliptica cut across debris flow beds deposited between the glaciomarine sediments suggest that the high sedimentation rate caused by the proximity of a glacier likely forced the bivalves to the surface (Figure 6A) (del Valle et al., 2007; Strelin et al., 2014).

Our oldest sample from a depth of 3.30 m in the nPII and PC sections has a calibrated age of 7,910 ± 260 cal. a BP and a 2σ age range of 8,170–7,660 cal. a BP (Table 2, no. 15), in agreement with more recent studies (Strelin et al., 2014). Similar ages from a further 25 samples from Lithofacies 1–4 convincingly show that a glacier-free environment existed at a low altitude in the mouth of Potter Cove before 8.2 cal. ka BP (Figures 6A–C; Table 2). This new minimum deglaciation age is younger than, but consistent with, previous ages from these deposits (cf. Sugden and John, 1973) and our new minimum ³He cosmogenic nuclide exposure age from Three Brothers Hill of 9.4 ± 2.6 ka.

Our paleoenvironmental interpretations for the nPII and PC sections are in broad agreement with previous studies. However, we could not replicate older chronologies even though the nPII is 2.5 m deeper than the nearby sections dated by Sugden and John (1973) and del Valle et al. (2007). Our Bayesian age-depth and probability density phase modeling (Figure 6B) reinforces that the older Sugden and John (1973) ages from Lithofacies units 1–5 represent the reworking of an older phase of glacier-free conditions in the outer peninsula or they are statistical outliers, possibly related to improvements in AMS radiocarbon dating since the 1970s.

Our revised early Holocene terrestrial deglaciation history for Potter Peninsula is consistent with evidence for enhanced glacier retreat from Maxwell Bay and sustained deglaciation on the nearby Barton and Fildes peninsulas at low altitude by c. 8.0 cal. ka BP (Figure 9B) (Seong et al., 2008; Milliken et al., 2009; Oliva et al., 2019), and glacier retreat on Livingston Island, c. 8.3–7.5 cal. ka BP (Oliva et al., 2016b). These events followed a period of sustained early Holocene warming across the northern AP region (Figures 9J,K) (Ingólfsson et al., 2003; Strelin et al., 2006; Bentley et al., 2009; Mulvaney et al., 2012).

5.2 A c. 7 ka glacier readvance

Radiocarbon ages from the marine shell-bearing horizons of Unit 5 in the nPII section provide a maximum age constraint of c. 7.0 cal. ka BP for a readvance of the former tidewater Fourcade Glacier to the outer mouth of Potter Cove. Glaciodeltaic facies on the surficial morainal bank terrace imply subaqueous discharge close to the calving zone, with contorted deformation structures near the top of Unit 5 most likely resulting from glaciotectonic push and shear (cf. Hambrey and McKelvey, 2000) indicative of subglacial deformation. Lithofacies 6 represents a period of inter-tidal sand accumulation, and the massive matrix-supported, clast-rich diamicton deposited on top of both sections is interpreted as a subglacial till formed by a readvance at or shortly after c. 7.0 cal. ka BP (Figures 6A, D,9D).

Changes in RSL recorded in raised beaches and isolation basins of the SSI provide the most reliable evidence of deglaciation, ice loading, and readvance (e.g., Johnson et al., 2022). As isolation basins are less prone to reworking and remobilization than raised beaches, we attempted to corroborate evidence for a < c. 7 ka readvance in the nPII profile from isolation basins on Potter Peninsula (Figure 2C) below the 8 ka SSI Holocene marine limit at ∼16 m a.s.l. (Fretwell et al., 2010; Watcham et al., 2011) (Figure 9C) by looking for sediment records with marine to terrestrial transitions, but most have unsuitable gravel (or boulder-strewn) lake beds.

RSL curves for the AP and SSI exhibit a regionally coherent shift from rising to falling RSL focused on c. 8.0–7.5 ka, reflecting ice retreat across the northern AP and SSI and isostatic uplift outpacing sea level rise (Roberts et al., 2011; Watcham et al., 2011; Johnson et al., 2022), which possibly prevented further marine-upwelling induced retreat of the Potter Cove glacier. The new c. 7 ka maximum age constraint on glacier readvance for Potter Cove/Peninsula from the nPII section is coeval with a phase of renewed, but comparatively minor, uplift at c. 7.2–7.0 cal. ka BP associated with a standstill or readvance (and/or thickening) of the BIC first determined from isolation basin studies on Fildes Peninsula (Watcham et al., 2011).

A c. 7 ka readvance is also in broad agreement with surface exposure ages of c. 7.0 cal. ka BP from moraine boulders, which represent a glacier advance on Hurd Peninsula, Livingston Island (Hall, 2007), and land-based glacier readvances on James Ross Island, northeast AP, between 7.3 and 7.1 cal. ka BP (Figure 9I) (Strelin et al., 2006; Kaplan et al., 2020). Terrestrial and marine proxies in records from the northern AP indicate extensive glacial melt and open waters without sea ice in Maxwell Bay, and generally warmer than Holocene average conditions, similar to the present-day shortly after c. 7 ka BP (Figures 9A, F, G) (Milliken et al., 2009; Watcham et al., 2011; Mulvaney et al., 2012; Peck et al., 2015; Roberts et al., 2017). Therefore, a precise climate connection between these records and glacier readvance at c. 7 ka on Potter Peninsula remains unclear and is discussed further in Section 5.6.

5.3 Mid-late Holocene deglaciation

Radiocarbon ages from basal lake sediments suggest that glaciers across the SSI and northern AP had receded to within or close to their present-day limits by c. 6 cal. ka BP (Mäusbacher et al., 1989; Watcham et al., 2011; Roberts et al., 2017; Giralt et al., 2020; Gomez et al., 2020). Marine sediment records from Maxwell Bay show decreasing sea surface temperatures (SST) and increasing sea ice cover between c. 5.9 and 2.6 cal. ka BP (Milliken et al., 2009), but we found no evidence for glacier readvance on Potter Peninsula in this period (Figures 6C, 9D).

Hall (2010) and Watcham et al. (2011) suggested that RSL fall was broadly continuous after 6 cal. ka BP (Figure 9C). Although del Valle et al. (2002) considered abandoned elevated coastal sediments at 17.1–14.4 m a.s.l. at Stranger Point with radiocarbon ages of 5.5–5.4 cal. ka BP to provide a minimum age for beach formation and the onset of the deglaciation on Potter Peninsula, it seems more likely that these deposits relate to continued uplift following early Holocene deglaciation and retreat after the c. 7 ka readvance. The latter interpretation is consistent with a maximum age of 4.3 cal. ka BP for glacier sediments that overlie the highest continuous raised beach at Stranger Point at ∼12 m a.s.l. and RSL curves for the SSI (Fretwell et al., 2010; Watcham et al., 2011; Lindhorst and Schutter, 2014; Johnson et al., 2022).

Jeong (2006) proposed that the diameter of sorted vegetated polygons on KGI is directly related to formation age, with the largest polygon (3.23 m) dated at 5.4 ± 0.1 cal. ka BP. Polygons ∼2.5 m in diameter were measured in this study and are tentatively assigned ages >2.6 ka using this method. Deglaciation ages derived using this method have obvious limitations but provide the only available age estimate for the periglacial landform assemblage (LA-4) on Potter Peninsula. Nevertheless, this age is consistent with ages of 4.5 to 2.8 cal. ka BP for a period of sustained mid-to-late Holocene warming, termed the Mid-Holocene Hypsithermal (MHH) in Bentley et al. (2009). Now referred to as the Holocene Hypsithermal (HH in Figures 9K, 10D), it has been identified in lake sediments from across the northern AP and (sub)Antarctic islands when glaciers were mostly situated close to, or landward of, their present-day limits (Björck et al., 1991a; Björck, 1991b; Ingólfsson et al., 2003; Sun et al., 2005; Bentley et al., 2009; Strother et al., 2015; Foster et al., 2016; Roberts et al., 2017).

FIGURE 10

5.4 Neoglaciation in the last 2,000 years

New geomorphological mapping and chronological evidence presented in this study show the Warszawa Icefield/Fourcade Glacier advanced beyond present-day limits on at least three occasions during the late Holocene: after c. 1.7–1.4 ka, following the return to “neoglacial” conditions (cf. Ingólfsson et al., 2003; Bentley et al., 2009), after c. 0.7 ka and at 1956 CE.

5.4.1 <1.7–1.4 ka readvances

The “Inland Outcrop,” a massive matrix-supported sandy diamicton sequence, ∼700 m from the present glacier front, was revealed in 2013 following at least ∼60 years of retreat of the Warszawa Icefield across the inner Potter Peninsula (Figure 2A). This diamicton was probably derived from the foliation of entrained subglacial debris at the glacier front, forming unstable ridges of till with sedimentary characteristics of re-sedimentation from a primary subglacial till by melt-out. During melt seasons, its foliation likely collapsed, and any orientation has now been lost (cf. Hambrey et al., 2015).

Warmer and/or more humid conditions that existed before c. 2.2 cal. ka BP, implied by the presence of moss within the “Inland Outcrop” (Figures 6C, 9D), are consistent with similar conditions implied by the first of the two closely-spaced phases of freshwater subaquatic moss deposition in lake records from Fildes Peninsula at c. 2.2–1.9 cal. ka BP and 1.5–1.3 cal. ka BP (Roberts et al., 2017). Remnants of moss layers scattered randomly at its base have median calibrated ages of 1.8 and 1.7 cal. ka BP and a 95% summed probability range of 2.2–1.4 cal. ka BP (Table 2), meaning they were most likely emplaced when the Warszawa Icefield had readvanced further landward by c. 1.7–1.4 cal. ka BP (Figures 6C, 9D).

We link the “Inland Outcrop” to the ice-cored Stranger Point moraine system based on their relative positions to the present glacier margin and the degree of degradation and extent of mass-wasting processes (e.g., landslides and weathering) within the diamicton (Figures 2, 8A–D, F; Table 1). As these sequences are in front of a former theoretical limit of the Warszawa Icefield and form part of the inactive glacial landform assemblage (LA-2) (Figure 2; Table 1), we propose that both sequences were formed by the same glacier advance and subsequent melt-out processes that produced thrusting and deformation of subglacial till at the glacier margin of the Warszawa Icefield (Figures 2, 8C; Table 2).

A readvance sometime after 2.2 cal. ka BP and by c. 1.7–1.4 cal. ka BP in the central foreland of Potter Peninsula is consistent with ages of c. 1.9 cal. ka BP obtained from raised beaches up to 7.3 m a.s.l. overlying volcanic bedrock and till on the south side of Potter Peninsula that imply increased late Holocene ice loading (Lindhorst and Schutter, 2014). The timing is also consistent with offshore evidence of a glacier readvance dated to c. 2.6–1.6 cal. ka BP from the outermost marine moraine complex, M1, in Potter Cove (Figure 1C; Wölfl et al., 2016) and marine diamictons deposited by ice rafting in Maxwell Bay between c. 1.4 and 1.1 cal. ka BP (Ingólfsson et al., 2003; Yoon et al., 2004). The M1 moraine represents the furthest preserved maximum position of the Fourcade Glacier in the late Holocene (Wölfl et al., 2016).

Although late Holocene ice loading and readvances on Potter Peninsula are too small to be modeled by regional-scale ice sheet and glacio-isostatic adjustment models (e.g., Whitehouse et al., 2012a, b), the most recent RSL curve for the SSI (Johnson et al., 2022) (Figure 9C) hypothesizes that the continuous decline from the Holocene marine limit between 8 and 6 ka was interrupted by increased isostatic uplift between 1.5 and 0.5 ka BP. Rebound at this time relates to the “most significant” readvance event of the mid-late Holocene (Hall, 2010) and is consistent with published raised beach data and evidence for a post c. 2 ka “neoglacial” readvance on Potter Peninsula.

Colder “neoglacial” conditions persisted on the SSI and northern AP from c. 2 ka (2070 ± 50 cal a BP in Čejka et al. (2020)) onward (Ingólfsson et al., 2003; Bentley et al., 2009; Mulvaney et al., 2012; Roberts et al., 2017; Palacios et al., 2020; Čejka et al., 2020). However, readvance of the BIC on Fildes Peninsula remained restricted close to, or within, present-day limits (Hall, 2003; Hall, 2007; Hass et al., 2010), whereas on the Barton Peninsula (KGI) and the Byers Peninsula (Livingston Island), radiocarbon dating of basal lake sediments and exposure ages from bedrock outcrops show that no substantial glacier readvances have occurred in the last 2.4 cal. ka BP (Seong et al., 2008; Oliva et al., 2016a; Oliva et al., 2016b; Oliva et al., 2019). There is also some evidence for short-lived warmer phases in terrestrial records from the northern AP and the sub-Antarctic islands between c. 1.8 and 0.7 cal. ka BP (Bentley et al., 2007, 2009; Hall, 2007; Hall, 2010; Foster et al., 2016; Guglielmin et al., 2016; van der Bilt et al., 2017) and generally warmer SST on the western AP between c. 2 and 0.8 cal. ka BP (Shevenell et al., 2011; Etourneau et al., 2013).

5.4.2 <0.7 ka readvances

Our new geomorphological mapping revealed vegetated protalus ramparts between 30 and 60 m a.s.l. on Potter Peninsula, approximately 200 m below the current ELA, implying that cooler and drier environmental conditions existed when they were formed. By examining their cross-cutting relationships with raised beaches on Livingston Island, Serrano and López-Martı́nez (2000) and Kääb (2007) linked their formation to that of low altitude protalus landforms and rock glaciers which advanced between c. 0.8 and 0.3 ka. Applying a similar principle, the Stranger Point moraine system crosscuts raised beaches at 6 m a.s.l., implying a maximum age of 0.7 cal. ka BP for a readvance at this location (Heredia Barión et al., 2019). This age estimate is consistent with coastal gravel deposits at 6 m a.s.l. that were deposited on top of a till deposit at the west and north coast of Potter Peninsula sometime after 0.65 cal. ka BP (Sugden and John, 1973; Curl, 1980; Clapperton and Sugden, 1988; Lindhorst and Schutter, 2014; Heredia Barión et al., 2019).

Supporting evidence for a readvance after c. 0.7 ka and subsequent retreat also exists in the lake sediment records on Potter Peninsula. 1) The first evidence is the modeled age for the contact between the basal diamicton unit and the uppermost laminated sediments in the Matias Lake record of c. 0.3 cal. ka BP (Supplementary Figure S5). Del Valle et al. (2004) considered similar laminated sediments in adjacent Rudy Lake to be annual couplets (varves) with sedimentation rates of 1.2–1.7 mm yr⁻¹. This rate is similar to that obtained by ²¹⁰Pb and ¹³⁷Cs dating of the uppermost unit of laminated sediments in Matias Lake. The presence of basal gravels with large clasts grading into the uppermost unit of fine “varves” in our new cores from Matias Lake implies the lake was seasonally free of lake ice but with the glacier front nearby. Lead-210 dating of the uppermost laminated sediments of Matias Lake shows that glaciers had retreated from the Matias-Rudy Lake area by the mid-1850s. 2) The second is a minimum age of 0.5 ka from Rudy Lake, based on its relationships with adjacent moraines and varve counting of lake sediments above a diamicton deposit (Birkenmajer, 1998; del Valle et al., 2004). 3) The third is the shift to higher productivity and reduced erosion of sediment between zones 2 and 3 in Lake L15, evident in the higher TChl content and lower magnetic susceptibility (MS), Ti and Ca values in zone 3. This is indicative of ice retreat from the outer Potter Peninsula and the establishment of persistent seasonally lake-ice-free conditions sometime after 0.6 cal. ka BP (Figure 7A).

Data from marine sediment cores and the submerged moraine complexes M2 and M3 in Potter Cove show that the Fourcade Glacier was located forward of its present position in the inner cove between 0.5 and 0.1 cal. ka BP (Figure 1C) (Hass et al., 2010; Majewski et al., 2012; Wölfl et al., 2016; Munoz and Wellner, 2018). On Fildes Peninsula, ice-cored moraines composed of till, ice-proximal meltwater fan deposits and ice-marginal pond sediments were forming adjacent to present-day glacial margins by thrusting processes, incorporating remnants of mosses dated to c. 0.83–0.65 cal. ka BP (Hall, 2007). These mosses represent warmer/milder conditions before a <0.7 ka readvance, and their ages are coeval with the timing of the Mediaeval Climate Anomaly (MCA; Figure 9K) c. 1–0.7 cal ka BP, 950–1250 CE (Mann et al., 2009; Kaufman et al., 2020).

More widely, using radiocarbon and optically stimulated luminescence dating of raised beaches, Simms et al. (2012; 2021) showed that the last glacier readvances occurred across the SSI and the AP between c. 0.45 and 0.25 ka BP and that increased ice loading is consistent with increasing rates of post-MCA RSL decline on KGI in the last 500 years (Bentley et al., 2005; Hall, 2010; Watcham et al., 2011; Johnson et al., 2022) (Figures 9C, K). After this readvance, a short-term and rapid acceleration in rebound occurred between 1700 and 1840 CE (0.2–0.1 cal. ka BP), with a modeled peak uplift rate of 12.5 mm a-1 (Simms et al., 2012). Simms et al. (2012) also showed that RSL curves for the SSI showed accelerated fall rates in the last 500 years, indicative of prior increased ice loading, whereas those from western and north-eastern AP exhibit a gradually decreasing uplift rate, implying continuous deglaciation, despite extensive geomorphological evidence of mid-late Holocene glacier readvance (Bentley et al., 2005; Hall, 2010; Roberts et al., 2011; Watcham et al., 2011; Hodgson et al., 2013; Kaplan et al., 2020).

Evidence for a <0.7 ka readvance on Potter Peninsula is also broadly coeval with: 1) more persistently colder conditions on or around the SSI and the AP (Figure 9E) (Yoon et al., 2004; Liu et al., 2005; Hall, 2007; Hass et al., 2010; Monien et al., 2011; Angiel and Dąbski, 2012; Majewski et al., 2012; Hu et al., 2013; Roberts et al., 2017; Simms et al., 2021); 2) generally increased sea ice in Maxwell Bay and colder conditions in Bransfield Basin between c. 1.4 and 0.1 cal. ka BP (Heroy et al., 2008; Milliken et al., 2009; Mulvaney et al., 2012; Abram et al., 2013); 3) consistently colder reconstructed SST in proxy records from the western AP between c. 0.5 and 0.2 cal. ka BP (Figure 9H) (Shevenell et al., 2011; Etourneau et al., 2013); 4) evidence of a post “MCA” readvance elsewhere on the AP and some sub-Antarctic Islands (Bentley et al., 2007; Hall, 2007; Bentley et al., 2009; Hall, 2010; Roberts et al., 2010; Davies et al., 2014; Guglielmin et al., 2016; van der Bilt et al., 2017; Verfaillie et al., 2021), at a broadly similar time to the Northern Hemisphere “Little Ice Age” (0.55–0.25 cal. ka BP, 1400–1700 CE) (Mann et al., 2009; Kaufman et al., 2020).

5.5 Recent (post-1956 CE) deglaciation and active geomorphological processes

Aerial photographs show that the Warszawa Icefield was also in an advanced position close to the “Inland Outcrop” in 1956 CE (Figure 3). In 1956 CE, the front of the Fourcade (tidewater) glacier was located 1,200 m away from the mouth of Potter Cove (Figures 1, 2) and has since experienced a net retreat of more than 1,000 m along the flow line. Aerial and satellite imageries show that the Warszawa Icefield and the Fourcade Glacier then retreated rapidly to their present-day position between 1956 and 2013 CE (Figure 3B), and since 2016 CE, satellite imagery has shown that it has terminated on land (Jerosch et al., 2018).

After deglaciation, periglacial/frost action and mass movement became widespread on Potter Peninsula (Figures 3D, 4J). Rapid mass movements and rockfalls have transferred the sedimentary products of the periglacial processes down-slope, creating talus accumulation at the foot of bedrock plateaus and around Three Brothers Hill (Figure 2; Table 1). Cryoturbated sorted circles and polygons form on gentle slopes above 50 m a.s.l. and where surficial sediments are saturated with water due to the presence of a permafrost table that inhibits drainage. Repeated freeze-thaw processes, rather than permafrost (Matsuoka, 2001), most likely led to the formation of cryoturbated stone stripes on steeper slopes and gelifluction lobes at the base of the bedrock plateaus.

The inactive Stranger Point moraine system has been actively degraded by mudflows, solifluction, and thermokarst processes (Figures 2, 4G, 8A,C) (cf. López-Martínez et al., 2012; Oliva and Ruiz-Fernández, 2017). Meanwhile, thermokarst processes are prevalent in active ice-cored moraines close to the glacier margin at Stranger Point and across Potter Peninsula (Figures 4F, 8A, C) (John and Sugden, 1971; Sugden and John, 1973; Hochschild, 1995; Lukas, 2011). The large, suspended sediment load produced by intense glaciofluvial processes at the glacier front and the incision of relict moraines (Figures 4A, H) forms active glaciofluvial fans and deltas that prograde toward and into Potter Cove (Figure 2). Meltwater and suspended sediment change the physical and chemical properties of Potter Cove (Henkel et al., 2013; Monien et al., 2017; Henkel et al., 2018), impacting the local biological communities and, ultimately, the biogeochemical cycle of the Southern Ocean (Quartino et al., 2013; Pasotti et al., 2015a, b; Sahade et al., 2015; Monien et al., 2017; Falk et al., 2018; Henkel et al., 2018).

The uppermost 6.5 cm of Lake L15 and its surrounding environment at the base of Three Brothers Hill is dominated by freshwater subaquatic, littoral, and terrestrial mosses and cyanobacterial mats with post-1950 CE ages (Supplementary Figure S2; see Supplementary Material for further details). Elsewhere on KGI, increased primary productivity and moss colonization of the lake-water interface is commonplace in seasonally lake-ice-free and stable lacustrine environments undisturbed by glaciofluvial inputs and indicative of warmer and more favorable climatic conditions (Björck et al., 1991c; Roberts et al., 2017).

5.6 Mechanisms of change

After ice retreat from offshore LGM limits (Termination I), early Holocene deglaciation and subsequent glacier fluctuations on Potter Peninsula primarily reflect interactions between changes in climate and RSL (Bentley et al., 2009). It seems likely that a c. 7 ka glacier readvance was driven by a combination of changes in glacier stability associated with increased precipitation, regional glacio-isostatic adjustment, and/or local relative sea-level variations (Figures 10E–G) rather than major thickening of the glacier front, fjord geometry, and fjord depth (Mercer, 1961; Van der Veen, 1995).

Subsequent changes in Holocene glacier mass balance across the SSI were controlled by hemispheric to global scale climate processes and long-term (millennial-scale) trends in solar insolation at 62°S (i.e., increasing to c. 5 ka maximum, declining after c. 5 ka) (Bentley et al., 2009). Centennial–millennial-scale fluctuations in solar irradiance and interannual to decadal-scale changes in the SHW and SAM modulate long-term trends in insolation, driving centennial–millennial changes in temperature, precipitation, wind, sea ice formation, and glacier advance and retreat on the northern AP, the SSI, and southern South America (Figures 9, 10) (Bentley et al., 2009; Lamy et al., 2010; Varma et al., 2012; Charman et al., 2018; Moreno et al., 2018; Baggenstos et al., 2019; Reynhout et al., 2019; Kaplan et al., 2020; Verfaillie et al., 2021). The interannual relationship between insolation, temperature ENSO, and the SAM is also thought to operate across decadal-to-millennial timescales during the Holocene (Perren et al., 2020) (Figures 10B–D), impacting the AP (e.g., Dickens et al., 2019).

Deglaciation of the SSI before c. 8.2 ka coincided with a sustained phase of positive SAM-like conditions, increased global irradiance, increasing spring/summer insolation at 62°S, warmth and sustained glacier retreat, and more open water on and around the northern AP (Figures 9F–K, 10A and references therein). A c. 7 ka readvance on Potter Peninsula was initiated during a short-lived period of lower than Holocene average global solar irradiance (Figure 10A) that would have lowered the ELA across the SSI and negative (colder/more humid) SAM-like conditions (Figure 10D), likely leading to increases in year-round precipitation as snowfall (Kaplan et al., 2020).

In contrast, after c. 5 ka, a steadily declining trend in insolation at 62°S coupled with more persistently colder/wetter and more negative SAM-like conditions, stronger SHW, and an enhanced ENSO (especially during the last 4 ka) likely drove late Holocene readvances after c. 1.7–1.4 cal. ka BP, <0.7 cal. ka BP and by 1956 CE (Figures 10A, D, G) (Charman et al., 2018; Saunders et al., 2018; Yang et al., 2019; Perren et al., 2020).

More recently, the longer-term pattern of interannual variability in the SAM has been altered by the ozone hole over Antarctica, which, combined with increases in greenhouse gases and global temperature, is forcing a more positive trend in the SAM and further enhancing the ENSO (Marshall, 2007). The regionally synchronous retreat of most land-terminating glaciers across the northern AP and SSI since the mid-20th Century provides strong evidence that glacier retreat on KGI was initiated by increasing air temperatures (Braun and Gossmann, 2002; Bentley et al., 2009; Kaplan et al., 2020). Recent (post-1956 CE) deglaciation has exposed new onshore areas on Potter Peninsula to intense sediment reworking and redistribution, producing characteristic paraglacial, proglacial, and periglacial landforms on the glacier foreland of Potter Peninsula.

6 Conclusion

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