Two GPR systems were used in this study: a ground-based system deployed at all sites, and a UAV-based system used at Hohsaas. An overview of the setups for both systems is provided in Fig. S7.

All ground-based GPR surveys were conducted using a Sensors & Software pulseEKKO Pro system. The choice of antenna frequency balanced penetration and robustness against scattering on the one hand, and ease of handling on the other. We used 100 MHz antennas at most sites as a conservative choice, ensuring good-quality bedrock reflections given the expected ice thickness and the scattering from englacial water and inhomogeneities. At Alphubel, the steep terrain favored the more compact 250 MHz antennas which still imaged the bedrock reliably. Antenna positions were tracked with a Leica differential GNSS receiver operating in real-time kinematic (RTK) mode, providing centimeter-level positioning accuracy. Antenna elevations were corrected for the offset between the GNSS and radar antennas and for the antenna height above the ground surface. Depending on surface conditions, the antennas were either carried approximately 50 cm above the surface or, in areas with sufficient snow cover (Sex Rouge, Corvatsch, Alphubel), mounted on a sledge and pulled at approximately 10 cm above the surface. In both configurations, a constant separation of 2 m between transmitter and receiver was maintained. At most sites, the survey included a profile approximately along the glacier centerline connecting the borehole locations.

At Hohsaas, we employed a recently developed UAV-based GPR system , using an M300 RTK UAV (DJI) carrying a GPR unit with a single transmit-receive antenna at a center frequency of 80 MHz. Survey line positions were determined using a D-RTK 2 high-precision GNSS mobile station (DJI), providing centimeter-level RTK positioning (nominal accuracy 1cm+1ppm horizontal, 2cm+1ppm vertical). By flying a regular grid of parallel lines with 4 m spacing oriented approximately along the glacier contours, we obtained a 3D data cube covering nearly the entire glacier area. For the thermal interpolation, we extracted two down-slope profiles from this grid by sampling across the survey lines, each profile intersecting three boreholes.

All ground-based GPR data were processed using a standard workflow in Geolitix ; details are provided in Sect. S2.8.1. The UAV-based GPR data from Hohsaas were processed using a separate workflow described in Sect. S2.8.2. In both cases, the basal reflection was manually picked and ice thickness computed as the difference between surface and bed elevations, establishing the geometric framework needed for interpolation of the temperature data. We estimate the resulting bedrock-depth uncertainty at approximately ±5 m, arising from the manual picking of the basal reflection and the vertical resolution of the radar. We primarily used the radar data to characterize the englacial thermal structure: In GPR data, temperate ice containing free water at grain boundaries produces diffuse scattering and signal attenuation, whereas cold ice is largely transparent to radar . The cold-temperate transition surface (CTS) is therefore identifiable as the boundary between diffuse scattering and low-reflectivity zones, used here to independently validate borehole-derived CTS positions and to extend the thermal interpretation beyond the borehole network.

Annual point mass balance was measured at all Tinytag borehole locations (1TT and 2TT). Snow depth measurements were additionally collected across each glacier in April 2025 using manual probing at 15–60 points per site (Fig. S9). Snow density was measured on site using a coring device, or adopted from measurements on nearby glaciers. At Corvatsch, Sex Rouge and Tortin, these measurements are part of monitoring programs started around 2010. Point measurements of ablation and accumulation were spatially extrapolated to the entire glacier using a distributed mass balance model , which accounts for spatial variability and the inhomogeneous distribution of field observations. The model combines a distributed accumulation model with a temperature-index melt model that includes a potential-solar-radiation term , rather than a full surface energy balance. Melt is computed from air temperature and potential clear-sky solar radiation, while accumulation is derived from precipitation, scaled by a correction factor and distributed using observed snow-accumulation patterns to capture the spatial variability of snow deposition across the glacier. Model parameters are calibrated against the seasonal stake and snow-probing measurements at each site. A full description of the model, including the melt and accumulation formulations and the calibration procedure, is given in . This approach allows inferring surface mass balance distribution on a 10 m spatial grid.

The calibrated mass balance model can also be driven by meteorological forcing outside of the period covered by in-situ measurements. In this case, the model is constrained by observed long-term changes in ice volume . Supported by the detailed measurements of the recent years, informing about the spatial variability of mass balance components, the model thus provides a gridded annual mass balance distribution for each glacier from 1970 to 2025. These grids were used to derive spatially distributed firn cover maps by cumulatively tracking mass balance at each grid cell through time. Firn thickness was updated annually by adding positive mass balance (accumulation) and removing negative mass balance (ablation) from the existing firn column, with ablation removing the most recently deposited layers first. Grid cells were classified as firn-covered if they accumulated snow in two consecutive years. This criterion excluded single years with a larger end-of-season snow extent from the firn extent. Firn layers persisting for more than 20 years were considered ice, a simple assumption consistent with typical densification timescales in the Alps . This approach yields annual maps of firn thickness and extent from 1970 to 2025 for each glacier. Because the reconstruction depends primarily on the accumulation component, its reliability rests on how well snow accumulation and its spatial distribution are captured. For Alphubel, Chessjen, and Hohsaas the accumulation model is constrained by only a single year of snow-probing data, fewer than at the long-monitored sites (Corvatsch, Sex Rouge, Tortin). We therefore assessed the reconstructed firn extent independently against firn/perennial-snow boundaries manually digitized from high-resolution SWISSIMAGE orthophotos ; the comparison shows no systematic bias and comparable agreement across all six glaciers, including the less-constrained polythermal sites (Sect. 3.1). From the annual firn cover maps we computed the time since firn loss for each cell.

Tinytag sensors installed in 2024 recorded nearly a full annual temperature cycle at all six sites (Fig. ). Over the measurement period, surface melt thinned the ice above each sensor, so that effective sensor depths changed between installation (Zinit) and retrieval (Zfinal). For simplicity, all depths stated in the following refer to the installation depth. The reduction in effective depth ranged from approximately 0.4 m to more than 2 m, meaning that sensors progressively sampled shallower ice over the course of the measurement period. We expect this to result in slightly larger observed seasonal amplitudes. Unless stated otherwise, reported temperatures carry an uncertainty corresponding to the calibrated sensor accuracy, ±0.2 °C for the Tinytag and ±0.05 °C for the Geoprecision sensors. Seasonal temperature amplitudes varied by more than two orders of magnitude across sensors and sites, with the largest recorded amplitude of 3.64 °C at Tortin (GT1TT, 4.2 m), the smallest of 0.03 °C at Hohsaas (HS1TT, 15.6 m). As expected, amplitudes decreased systematically with depth and the temperatures at depths of 10 to 15 m lagged roughly six months behind the surface (Fig. a–d), meaning that the lowest temperatures occurred in the summer, the highest in winter.

At Alphubel, Chessjen, Hohsaas, and Sex Rouge, temperatures showed no sign of approaching 0 °C at the instrumented depths (6.8–15.4 m), suggesting that the winter cold wave could not be compensated for during summer. Two sensors are exceptions to this: SR2TT (10 m) and HS1TT (15.6 m), whose temperature maxima (-0.18 and -0.1 °C, respectively) fall within the ±0.2 °C sensor accuracy, meaning temperate conditions at these depths are likely. Shallow sensors at Tortin and Corvatsch were trending toward 0 °C by the end of the observation period and are most likely approaching temperate conditions later in the year. The final readout was conducted before the 2025 summer warming signal had fully propagated through the ice column. GT1TT (4.2 m) and CV1TT (2.0 m) are therefore expected to reach their annual maximum during or shortly after the melt season, while sensors at 10 to 15 m depth likely do so between December and March.

At Chessjen, Hohsaas, and Sex Rouge, mean temperatures increase with depth. Temperatures at greater depth are 1.2–1.5 °C warmer than those nearer the surface at Chessjen, 0.6–0.9 °C warmer at Hohsaas, and approximately 0.1 °C warmer at Sex Rouge. This distinction is less obvious at Alphubel, where mean temperatures appear to vary more with borehole position than with depth, ranging from -2.0 to -2.2 °C at AH2TT (the upper borehole) and from -2.3 to -2.7 °C at AH1TT further downslope. At Alphubel, Chessjen, Hohsaas, and Sex Rouge, where the deepest sensors exhibit amplitudes below 0.3 °C, we estimated the depth of zero annual amplitude (ZAA; see Sect. S2.7) to range from 18 m at Hohsaas to 23 m at Sex Rouge (see annotation in panels a–d).

Short-term temperature jumps are visible in both the Tinytag and Geoprecision time series data at Alphubel and Chessjen. At CJ2TT, the shallow sensor recorded an abrupt warming of 1.64 °C on 18–19 July (Fig. b), and at Alphubel, the second-highest borehole (AH1G) showed a temperature jump of approximately 0.15 °C on 17–18 August (Fig. S14a). Since these events occurred during the melt season and have transient character, we interpret them as the effect of englacial meltwater, likely routed via crevasses and englacial channels. The resulting temperature jumps reflect heat input to the surrounding ice, through direct exchange with liquid water, latent heat release upon refreezing, or both – processes central to the thermal structures discussed in Sect. .

The availability of previously published borehole data at Sex Rouge and Corvatsch provides a rare opportunity to investigate how englacial temperatures at these sites have evolved over decadal timescales (Fig. ). The earlier measurements represent multiple, single-day snapshots, while our data comprise a full annual time series. To compare, we extracted the matching days of the year from our time series for each prior measurement date. Matching the day of year controls for the seasonal phase of the signal but not for inter-annual weather variability, which still influences temperatures within the seasonally affected zone sampled here. We therefore compare the historical single-day measurements against the full modern annual range at each depth (min-max envelope in Fig. ), and interpret differences within the seasonal zone in terms of the broader thermal structure and its evolution rather than as precise climatic-trend magnitudes.

At Sex Rouge, data from 2013–2015 documented a polythermal structure with a cold surface layer of approximately 20 m overlying a temperate core. We also found a persistent cold surface layer with a strong temperature gradient trending to 0 °C with depth (Fig. a and b), suggesting that this polythermal structure likely still persists today. Relative to the 2013–2015 measurements, temperatures at 10 m rose by 0.65 °C at SR2TT and by 0.23 °C at SR1TT, while at 15.2 m temperatures fell by 0.11 °C at SR1TT, remaining within sensor accuracy (±0.2 °C). However, with approximately 20 m of surface ice lost through ablation since 2013 , our sensors sample ice that was located considerably deeper during the earlier measurement period. The hollow red diamonds in Fig.  show our current temperatures projected to their estimated depths at the time of the prior measurements, illustrating this reference-frame shift. This vertical shift complicates our interpretations. The observed signal reflects the effects of downward cold wave penetration and surface removal of cold ice through ablation, rather than atmospheric forcing alone. These competing drivers are discussed further in Sect. .

At Corvatsch, documented cold-based conditions with temperatures of approximately -4 °C at 13 m depth in 1999/2000. The glacier has since thinned to a point where the entire ice column, now only 9.3 m thick (as of 2026), is subject to seasonal temperature fluctuations (Fig. f). Temperatures at 4.3 m depth exhibit a clear seasonal cycle but remain sub-zero throughout the year, indicating that summer warming is insufficient to bring the ice to the melting point. Our measured temperatures at 4.3 m depth were on average -0.36 °C colder than the measurements at comparable depths from . The lower sensor reached the bed at 9.3 m but failed during equilibration, recording -1.3 °C with temperatures still declining at the time of failure, suggesting the bed remains cold today. Among the six study sites, Corvatsch occupies a unique topographic position on a narrow ice ridge. The ridge geometry limits snow accumulation and firn formation while exposing both flanks to lateral conductive heat loss, providing an additional cooling pathway beyond surface forcing alone . The persistently cold conditions observed here are therefore at least partly a consequence of this inherently cold topographic setting, alongside the long firn-free history of the site, discussed further in Sect. .

Because deep boreholes reaching the glacier bed were drilled only at Alphubel, Chessjen, and Hohsaas (Sect. ), the following analysis of the full vertical thermal structure is limited to these three sites. The shallow (≤15 m) boreholes at Corvatsch, Sex Rouge, and Tortin sample only the seasonally influenced near-surface layer. At all three deep-drilled sites, englacial temperature profiles reveal polythermal conditions (Figs. –), with a recurring spatial pattern: a temperate core in the upper reaches transitioning to partially and fully cold-based conditions near the terminus.

In all three glaciers, we determined the CTS to be at relatively shallow depths of 17–20 m in the upper reaches, while ice near the termini was found to be predominantly cold (Figs. a, a, and a, b). At Alphubel, temperate ice was confirmed between 25 and 33 m depth at AH1G and between 20 and 25 m at AH3G, with the transition to cold conditions occurring approximately 110 m upstream of the terminus (Fig. a). Notably, in both AH1G and AH3G, temperatures approach 0 °C at intermediate depths but drop below 0 °C again toward the bed, suggesting the temperate zone forms a sort of core zone. As neither AH1G nor AH3G have sensors shallower than 17 m, the cold ice overlying the temperate zone is not directly sampled but is assumed to correspond to the seasonally influenced surface layer (Sect. ). At Chessjen, the CTS was identified at the two higher-elevation boreholes (CJ1G and CJ3TT) at depths of 17 to 38 m, with the CTS descending toward the bed approximately 180 m upstream of the terminus, beyond which cold conditions dominate throughout (Fig. a). At Hohsaas, the CTS was identified at a depth of 20 m at the higher-elevation borehole (HS1G), from where it rises toward the surface and remains shallow (0–5 m) down-glacier until approximately 150 m from the terminus, where it begins to descend toward the bed. At the northern profile L1, the ice becomes progressively thinner and colder, suggesting a more laterally extensive cold zone in this part of the glacier (Fig. a).

Despite a broadly similar thermal structure across the three sites, basal conditions differ. At Alphubel, bed temperatures remained cold throughout, reaching as low as -2.1 °C at the terminus. At Chessjen, CJ2G reached the bed at 17 m depth, confirming temperate conditions in early August 2025, but cooled to -0.43 °C by mid-December (Fig. S15b). We interpret this as a seasonally evolving basal thermal regime: meltwater sustains temperate conditions through summer, while basal refreezing occurs once melt ceases in autumn. At Hohsaas, HS2G confirmed temperate basal conditions close to the terminus despite cold conditions in the overlying ice column, though basal coverage between HS2G and HS1G remains limited.

To complement the point measurements from the boreholes with a spatially continuous view of the thermal structure, we next consider the radar profiles. The GPR profiles at all three sites are broadly consistent with the borehole-derived thermal conditions, though the degree of agreement varies. At Alphubel (Fig. c), the profile shows a substantially cleaner signal than at the other two sites, with little scattering even in the area where boreholes confirmed temperate ice. This difference is at least partly attributable to survey timing: unlike the late-summer surveys at Chessjen and Hohsaas (August–September), the Alphubel GPR survey was conducted at the end of winter in mid-May (Fig. S2). At this time of year and at 3800 m a.s.l., liquid water is expected to be at its seasonal minimum regardless of thermal state, potentially suppressing scattering that would otherwise indicate temperate ice. Still, the clean signal below the seasonally influenced layer suggests limited englacial water content at depth, supporting the cold bed interpretation derived from the borehole temperatures. At Chessjen (Fig. c), the scattering pattern corresponds well with the interpolated temperature field. Radar scatter is high throughout the temperate zone up-glacier and diminishes toward the terminus below CJ1G. Notably, scattering extends to the glacier surface in the upper part of the profile, suggesting the ice column is temperate throughout and that the CTS detected here reflects the seasonal penetration depth of the winter cold wave rather than a stable year-round thermal boundary. At Hohsaas (Fig. c), a qualitatively similar pattern is observed, with increased scattering around the upper borehole HS1G and a largely clean signal in the cold terminus. At both Chessjen and Hohsaas, enhanced near-bed scattering suggests that temperate basal conditions are likely more extensive than the borehole network resolves.

While polythermal glacier structures can arise through a variety of mechanisms , the consistent pattern observed across all three sites, with temperate ice confined to higher elevations and cold conditions prevailing toward the terminus, points to a shared underlying control on englacial heat distribution, a point that we further examine in the following.

Temperate ice within glaciers can be generated or sustained by several heat sources, including geothermal heat flux, strain heating associated with ice flow, latent heat release from refreezing of percolating meltwater, and advection of temperate ice from upstream . The glaciers investigated here are all small (<0.5 km2), meaning that internal heat production from strain heating is expected to be limited , consistent with the low surface displacement rates observed at all sites (Table S17). Since the temperate ice identified here lies in the uppermost reaches of each glacier, where no upstream ice input exists, advection from upstream can also be ruled out as a means of supplying temperate ice. Geothermal heat flux in the region is modest (<80 mWm-2; ) and, as a basal heat source, cannot account for the temperate conditions observed throughout substantial portions of the ice column at these sites.

Instead, we believe that the observed thermal structure is primarily controlled by firn cover distribution and its temporal evolution. Firn acts as a meltwater reservoir, retaining water that percolates to depth, refreezes, and releases latent heat, warming the underlying ice and sustaining temperate conditions . The densification of temperate firn into glacier ice provides an additional, if secondary, supply of temperate ice to the system. Where firn is absent and meltwater runs off, both of these heat sources are lost and the ice can cool. This process chain is supported by modeling studies showing that meltwater refreezing and entrapment can dominate the internal heat budget of small polythermal glaciers, making the thermal structure highly sensitive to changes in surface melt and firn availability .

To examine the role of firn cover as a control on englacial thermal structure, we reconstructed its evolution across the six studied glaciers from the 1970s until today using modeled annual mass balance distribution constrained with GLAMOS mass balance data . Across the six sites, firn-free durations at the borehole locations span from less than a decade at Alphubel to more than three decades at Sex Rouge, Tortin and Corvatsch, providing a range of exposure histories against which to assess the thermal response to firn loss (Fig. ).

Firn-free exposure broadly reflects the thermal state across the six sites, with longer firn-free histories associated with more advanced cooling. The clearest examples are Sex Rouge and Corvatsch, both firn-free for more than three decades, where cold surface conditions prevail throughout the measured depth range. At Sex Rouge, the apparent near-surface warming relative to previous measurements (Sect. ) is most plausibly explained by surface ablation removing cold ice from above rather than a genuine change in thermal forcing. At Tortin, where borehole locations have similarly been firn-free for more than 30 years, shallow measurements alone are insufficient to characterize the full thermal structure.

At Chessjen and Hohsaas, the spatial pattern of time since firn loss broadly reflects the glacier's thermal structure, with the CTS deepest in areas that lost firn earliest, suggesting the up-glacier retreat of the firn line has left a spatial imprint on the thermal regime. At both sites, firn-free duration increases down-glacier, ranging from approximately 15–20 years at the uppermost borehole positions to 25–30 years toward the terminus. That temperate ice nonetheless persists beneath the upper reaches, despite one to two decades without firn, reflects the slow pace of conductive cooling combined with continued heat input. Conductive cooling of a temperate layer tens of meters thick proceeds on a timescale of decades ; over the firn-free period, the cold front driven by firn loss would be expected to penetrate roughly 20–25 m, broadly consistent with, though slightly deeper than, the 17–20 m CTS depths observed. That temperate ice is preserved to somewhat shallower depths than conduction alone would predict points to additional heat input. Two mechanisms contribute. First, the downward migration of the cold front is slowed by the latent heat released as residual pore water in the temperate ice refreezes ahead of it, so that temperate conditions persist longer than pure conduction would imply. Second, meltwater delivery to the firn-free temperate zone has likely not ceased: at both sites, patches of firn persist in the highest reaches of the glacier, up-glacier of the borehole network (Fig. ), and bergschrunds and crevasses provide englacial entry points for meltwater draining from these still firn-covered areas into the temperate zone below, sustaining it through latent-heat release. At Chessjen, the transient temperature jump observed in mid-July (Sect. ) provides indirect support that such pathways remain active.

The situation at Alphubel is more nuanced. While the firn-free terminus is cold and the firn-covered upper glacier is temperate, consistent with the pattern seen at Chessjen and Hohsaas, the underlying thermal history likely differs. The largely cold-based terminus is difficult to reconcile with a firn-free period of only 5–10 years, and the cold-temperate-cold vertical structure observed, where a temperate layer is sandwiched between a cold surface layer and a cold bed, suggests a fundamentally different thermal history at this site. Historical imagery from 1930 shows Alphubel as a much thicker hanging glacier with a prominent frontal break-off (Fig. S18), a morphology more characteristic of cold or polythermal high-elevation hanging glaciers with limited basal sliding . This is in contrast to the much larger valley glacier systems represented by Chessjen and Hohsaas, as documented in the historical imagery. We therefore interpret the cold terminus as largely predating the current period of firn loss, with the temperate core in the upper glacier representing a relatively recent feature sustained by meltwater percolation from the remaining firn cover above. A similar transition from a cold to a polythermal regime has been documented at Whymper hanging glacier on the Grande Jorasses south face .

The conceptual diagram in Fig.  illustrates our interpretation of how the thermal structure of small Alpine glaciers may have evolved in response to climate change and the associated firn cover retreat over the last decades. In the past (e.g., 1980s), the glaciers studied here were likely entirely temperate. Extensive firn cover, sustained by a cooler climate, provided an effective heat source through meltwater percolation and latent heat release upon refreezing . Active glacier dynamics promoted crevassing, creating englacial entry points for meltwater and further supporting temperate conditions throughout the ice column via cryo-hydrologic warming . Greater ice flow velocities also enhanced strain heating, providing an additional internal heat source .

Sustained atmospheric warming has driven firn retreat, progressively reducing latent heat input from meltwater percolation and refreezing. Where firn is lost, meltwater runs off at the surface rather than percolating into the pore space. The associated latent heat release upon refreezing is therefore lost . Simultaneously, ongoing glacier recession reduces ice flow velocities and with it strain heating, while also limiting crevassing and the availability of englacial meltwater pathways. In the absence of firn, both the loss of latent heat from refreezing and the reduced insulation allow the winter cold wave to penetrate the ice surface more effectively, initiating the development of a cold surface layer and, over time, cold ice. The thermal response to firn loss is neither instantaneous nor spatially uniform. This is illustrated by the contrast between profiles L1 and L2 at Hohsaas, which are only 20–30 m apart but show markedly different thermal structures (Fig. a). At L1, ice thickness (15–17 m) is comparable to the seasonally-influenced layer, leaving insufficient ice thickness for a temperate core to persist even where firn-loss history is similar to L2; at L2, the greater ice thickness (>30 m) provides enough insulated depth for temperate ice to remain despite ongoing cooling. This contrast highlights how ice thickness modulates the pace and completeness of thermal response to firn loss, independent of the surface forcing itself. Elevation, topography, residual meltwater supply, and the thermal inertia of the ice all modulate how quickly and how completely cold conditions develop following firn loss. In some cases, atmospheric warming may already offset the cooling initiated by firn loss, for instance where accelerated surface ablation removes cold ice faster than the winter cold wave can produce it. Alphubel may represent an exception to this thermal history, reminding us of the complexity of thermal responses to climate change. As a high-elevation hanging glacier, it likely harboured cold or polythermal conditions at its terminus even historically, in contrast to Chessjen and Hohsaas which, as much larger valley glacier systems, were likely fully temperate (Fig. S18).

The results presented here suggest that polythermal structures may be more widespread among small Alpine glaciers than previously recognized. The existing observational ice temperature record is strongly biased toward larger glaciers and high-altitude sites , largely overlooking smaller glaciers at lower elevations where firn is progressively being lost and polythermal structures may be actively emerging. Regional-scale thermal modeling, informed by the growing body of englacial temperature observations such as those compiled in the glenglat database and the records presented here, offers a path toward assessing the current distribution of polythermal structures and projecting their evolution under continued atmospheric warming. Such projections are particularly relevant given the hazard implications of polythermal configurations.

Polythermal glaciers are not dangerous per se. The thermal structures identified in this study are however consistent with a configuration known to promote meltwater accumulation beyond drainage capacity: a frozen terminus combined with a temperate upper glacier . During drilling at Alphubel, the only site with direct evidence for a fully frozen terminus, pressurised water escaped from borehole AH1TT at approximately 15 m depth, producing a water fountain 3 to 5 m high. This indicates that water pressure locally exceeded the ice overburden, implying that meltwater had been trapped sub- or englacially and could not drain. If such conditions occur at the glacier bed, basal friction may be substantially reduced locally . A broadly similar configuration, combining a cold glacier tongue with pressurised englacial water, has been discussed in the context of the 2022 Marmolada collapse . Whether similar configurations exist at other small Alpine glaciers remains unknown, but the number of potentially affected sites may be substantial.

To give a rough sense of how many glaciers share the size and elevation characteristics of our study sites, we filtered the SGI2023 for glaciers smaller than 0.5 km2 within an elevation range of 2900 to 3800 m a.s.l. The choice of 2900 m a.s.l. as a lower bound roughly excludes elevations where surface ablation dominates over cold ice formation, as suggested by our observations at Sex Rouge. Approximately 515 glaciers in Switzerland meet these criteria, a number that is likely an underestimate, as small hanging glaciers at high elevations are frequently mapped as part of glacier accumulation areas. This underscores the need for a systematic assessment of the thermal structure of small Alpine glaciers and its implications for hazard potential.