4.1 Fracture characterization

Observed fractures within the study area comprise joints, extension fractures, shear fractures and faults (Figure 3). The trace lengths of joints within the study area range from a few centimetres to more than 50 m (Figure 3a). Shear fractures in the area are recognized from abutting, obliquely trending extension fractures, which were generated due to slip along the hosting shear fracture and characteristically occur close to the termination of the shear fractures and define tip-damage zones (DZ; Figure 3b). An example of a fault, composed of sub-parallel shear fractures interlinked within step-over areas, is presented in Figure 3c. A characteristic feature for this type of fault is the presence of step-over DZs, where moderately dipping extension fractures abut the overlapping shear fractures, which constrain the width of the overstepping zone. The largest observed faults of this type exceed 120 m in length. Other types of observed faults comprise (i) up to 15-m-wide, laterally >200-m-long zones, characterized by densely spaced extension fractures abutting the parallel shear fractures at the fault margins (Figure 3d), and (ii) distinct faults with gouge-bearing core domains up to 20-cm-wide and abutting extension fractures within the DZ (Figure 3e; see also Figures 5c–e). The trace lengths of the faults with distinct fault-rock bearing cores are problematic to estimate, as they typically extend beyond the boundaries of the aerial photos. Joints typically lack any infill materials, but some N–S joints have infills of clastic sedimentary material and/or quartz (Figure 3f).

Joints occurring outside the deformation zones characteristically display orthogonal geometrical arrangements, as shown by the distinct clustering to sub-vertical, approximately N–S and E–W and gently NNE–dipping orientation maxima on the stereographic projections (Figure 2c). Shear fractures and faults generally follow the orthogonal joint patterns, as they define E–W and N–S-trending, sub-vertical orientation maxima on the stereograms (Figure 2c). The E–W set shows larger faults comprising wider DZ fracturing (Figures 2b and 3e–e), whereas the N–S shear fractures are less common, and they also typically occur as isolated features having weak linkages with parallel shear fractures. The orientations of the extension fractures deviate from the dominant orthogonal E–W and N–S trends of the other joints and faults as they define (i) a relatively tight belt-shaped girdle representing sub-vertical, NE–SW-trending fractures and (ii) a more scattered set representing E–W to WNW–ESE-trending fractures on the stereogram (Figure 2c). Within the latter set, several moderately NE-dipping planes correspond to the extension fractures within the linking-type DZs (Figure 3c).

4.2 Bedrock surface topography characterization, with linkages to fractures and faults

Topographic signatures of the bedrock surface outside the deformation zones are characterized by smooth and polished stoss sides (Figure 4a) and steep margins on one or two opposing lateral sides (Figure 4b–e). The stoss sides typically show sub-horizontal to gentle dips, following one of the dominant joint orientations, whereas the lee side and frequently also one other margin have been defined by the orthogonally arranged, E–W and N–S-trending joint sets. The edges of the sub-horizontal portions of the bedrock surfaces are rounded close to the vertical edges (Figure 4c,e), and localized erosion along the intersecting joint orientations indicates continued abrasion after the removal of some joint-bound bedrock blocks (Figure 4d).

Within and close to faults, the topographic signatures are more heterogeneous, as characterized by localized depressions of variable morphologies. Step-over zones linking parallel segments of larger shear fractures display slightly elongate but laterally limited depressions, which are bounded by the sub-vertical to vertical shear fractures on one side and by the moderately dipping extension fractures on the other side (Figure 5a,e). Faults characterized by two larger bounding shear fractures, and numerous variably oriented extension fractures in between, define laterally continuous depressions (Figure 5b,e). These depressions are relatively symmetrical, have sharp edges and show blocky and irregular morphologies along the base of the depressions. The largest fault with a gouge-bearing core domain (Figure 3e) displays the most pronounced topographic signatures, characterized by a narrow (0.5–1 m) gorge, which can be followed for >100 m from the shoreline to the inland area (Figure 5c,e). The hanging-wall side of the fault defines an up to 7-m-high overhanging wall with a distinct elevation contrast between the different sides of the fault (Figure 5e). The hanging-wall side of the fault is relatively intact with low fracture density, whereas the foot-wall side is populated by a dense set of moderately NE-dipping extension fractures (Figure 5d), which locally govern the attitude of the bedrock surface (sloping towards the sea on the right-hand side of Figure 5c).

4.3 Extrapolation of the outcrop-scale results to less-exposed adjoining areas (semi-regional scale)

This section documents our interpretation how the results from the detail area (Figure 2) can be used to provide a geologically motivated structural interpretation within a semi-regional scale, allowing usage of lower-resolution datasets and minimizing the time required for the structural characterization of a larger area. We emphasize that our semi-regional scale structural interpretation (Figure 6) is based solely on the topographic signatures of the LiDAR data and did not include validation through field observations. For this reason, uncertainties and some sampling bias between the low-lying coastal and higher inland areas may be present, and the interpretation should be considered a first-pass model that can be validated and improved with focussed field observations.

The dominant E–W and N–S trends of faults and larger shear fractures and the resulting topographic depressions observed at detailed scales (Figures 3-5) are also present within the semi-regional scale LiDAR DEM as distinct linear depressions of variable sizes (Figure 6a). Consequently, it is justified that we extrapolate our findings into wider scale to provide insight into structural controls in the development of landscape-scale topography. The magnitude of the elevation difference between the topographic highs and lows associated with semi-regional N–S and E–W-trending elongate depressions is approximately similar, but the topographic sections have contrasting styles (Figure 6b). The E–W section, at right angles to the ice-flow direction, displays relatively smooth morphological transitions from lows to highs, and the highest areas are characteristically rounded, which results in a continuous wave-form geometry with approximately 500-m wavelengths and 15- to 20-m amplitudes. In contrast, the ice-flow-parallel N-S section shows a less systematic morphology characterized by relatively flat domains at variable elevations, bounded by steep slopes associated with sub-vertical faults, particularly within the northern part of the section.

The above observations support the outcrop-scale findings that the brittle N–S-trending structures are less well developed than the E–W-trending fault systems. As a map-view summary, the following viewpoints can be highlighted: The development of narrow and elongate, fjord-like N–S-trending topographic lows is related to the combined effect of glacial abrasion along sub-horizontal bedrock surfaces and quarrying of flow-parallel joints (‘I’ in Figure 6a), without significant contributions from larger faults or fault zones. For some of such ‘fjords’ (or megagrooves), the abrupt termination and the associated steep to sub-vertical E–W slope appear to be controlled by a distinct discontinuity (referring to any unspecified fracture; IIa). These discontinuities have very limited lateral continuity outside the fjord termination, or they have a very limited effect on the DEM morphology (IIb); we attribute them to jointing instead of faulting.

Even though we found no direct evidence for the presence of N–S-trending faults or fault zones, comparable to the E–W-trending faults in the study area (Figures 3c–e and 5), it is likely that some well-defined elongate depressions (III) are the surface expression of N–S faults or fault zones. This is even more likely, as the interpreted fault at III spatially coincides with a narrow strait separating the Geta Island and the landmasses further west. Some of the approximately E–W-trending faults recognized within the detailed area (Figure 5c) can be linked to distinct topographic depressions (IV). Extrapolated further towards the south, these features define a network of anastomosing faults, with an overall E–W trend, but locally associated with NE–ENE-trending splay-type terminations (IV) and linkages between the E–W-trending segments (V). The distinct E–W-trending topographic feature at VI defines the boundary between the lowland in the north and higher elevation bedrock in the south, and we tentatively attribute it to one of the largest faults within the study area. A similarly significant, but more NE–SW-trending, curved and discontinuous fault (VII) bounds the topographic high along its SE margin. The frequently observed fault-related extension fractures (e.g., Figure 5e) are evident within the NE corner of the Geta Island (VIII), where they are associated with the E–W faults and locally define the geometry of the shoreline. The spatially restricted occurrence of these extension fractures suggest that the E–W faults terminate close to the shoreline.

5.1 The relationship of bedrock fractures and glacial erosion mechanisms

Recent studies have shed light towards understanding coastal erosion processes, and the findings indicate that fracture-controlled quarrying is the dominant erosion process (Buchanan et al., 2020). Major coastal erosion through quarrying typically requires rocks characterized by distinct layering, including mechanically weak layers that allow detachment of larger blocks (Prémaillon et al., 2021). Further, as the glacial striations are preserved on the oucrop surfaces, we conclude that the potential effect of coastal erosion in the Geta Island is limited to intersections of the fault damage zones and shorelines (with joint-bound block size of less than 2 m, e.g., right side of Figure 3d), while the bulk of erosion relates to glacial processes. Our results show clear linkages between the tectonic structures (faults and fault zones) and the surface topography of the bedrock. A such, the results are in line with the models in which the pre-existing fractures are the dominant mechanical discontinuities used in quarrying (Hooyer et al., 2012; Lane et al., 2015) and indicate that the generation of syn-glacial fractures is very limited. The tectonic origin of those fractures that contributed to the formation of the bedrock surface topography allows recognition of three specific patterns: (i) The attitudes of the extension fractures within the fault damage zones (tip DZ or linking DZ) have gentler dips than the shear fractures or joints occurring outside the fault zones, and (ii) fracture densities within the DZs are higher than within the surrounding areas, and (iii) fracture connectivity within the DZs is higher than within the surrounding areas.

The gentler fracture attitudes within the fault DZs have resulted in asymmetric topographic depressions characterized by sub-vertical to vertical slopes associated with the shear fractures or faults and moderately-dipping slopes associated with extension fractures (Figure 5a,c–e). Together, these features result in geometries comparable to the cliffs and overdeepenings in areas where the dip direction of the primary bedding in sedimentary rocks is opposite to the slope of the glacial valley (Kelly et al., 2014). By contrast, bedding surfaces parallel with the glacial valley slope (Kelly et al., 2014) or dipping towards the ice flow (Lane et al., 2015) have generated smoother topographic signatures characterized by abrasion. The higher fracture density and connectivity within the fault zones contribute to the fragmentation of bedrock into small block sizes and hence promote quarrying as the main mechanism of glacial erosion, as concluded by (Dühnforth et al., 2010). As the fragmented fault zones are topographically lower than surrounding abraded areas, it is evident that quarrying has been more effective than abrasion in eroding the bedrock. This finding is in conflict with the model in which abrasion dominates over other subglacial erosional processes beneath fast-flowing glaciers (Herman et al., 2015; Yanites & Ehlers, 2016) but is in line with other studies indicating quarrying dominates over abrasion (e.g., Loso et al., 2004), with particular reference to effective subglacial quarrying during the final stages of glacial cycles, when abundant subglacial meltwater was available to aid in quarrying (Glasser et al., 2020). The most pronounced glacial erosion could be achieved by glacial ripping, involving hydraulic jacking of pre-existing bedrock fractures by overpressurized sub-glacial meltwaters (Hall et al., 2020), but no widespread bedrock fragmentation associated with sandy or silty fracture infills was found in the present study area. For this reason, we attribute the zones of the most pronounced glacial erosion to quarrying, localized by the presence of pre-glacial faults and fault zones.

As discussed earlier (Sections 4.2 and 4.3), the N–S-trending fractures occurring parallel to the ice flow were locally used by glacial quarrying (Figure 4b,d,e) but did not have a major role in defining equally continuous elongate topographic depressions comparable to those associated with the E–W-trending faults. Consequently, the N–S-trending fjord-like topographic valleys (or mega scours) in the Geta Island are megagrooves (Bradwell et al., 2008; Eyles, 2012; Stokes & Clark, 2003), which are pathways for enhanced glacial flow that developed during successive glacial advances (Roberts et al., 2010). Megagrooves typically develop in well-stratified or layered rocks (Krabbendam & Bradwell, 2011), and their orientation may be controlled by the alignment of the lithological units and contacts (Roberts et al., 2010) or fractures (Bradwell et al., 2008). As s layering or stratification is absent from the bedrock of the present study area, it is likely that bedrock fractures had the primary contribution to the development of the observed megagrooves. Consequently, we suggest that the N–S-trending ice-flow-parallel fractures contributed to the process of lateral plucking (Krabbendam & Bradwell, 2011), which at small scales provided some control over the width of the roches moutonnées (Rastas & Seppälä, 1981) and at larger scales defined the margins of the larger elongate ice-flow-parallel fjord-like topographic depressions (Figure 6). Moreover, we infer that the sub-horizontal ice-bed parallel fractures—which are poorly constrained in this study—played a role in the effectivity of glacial quarrying (Woodard et al., 2019) and ice-bed morphology variation.

5.2 Implications to ice flow dynamics-driven glacier erosion in landscape evolution

Bedrock along the northern coast of the Åland Islands has been subjected to prolonged streamlining during successive glaciations, with particular contribution from the fast-flowing Baltic Ice stream during early Holocene (Greenwood et al., 2017). The flat-lying polished rock surfaces, distinct glacial striae and whalebacks are compatible with basal debris-driven abrasion under hard-bed conditions as the prevailing erosion mechanism during the Late Pleistocene (Glasser & Bennett, 2004). The scarcity of short distance transport of boulders and boulder fields on Åland further indicates that glacial quarrying, jacking or ripping (Hall et al., 2020; Krabbendam et al., 2022) were not the regionally dominant erosion mechanisms. As such, conditions of glacial erosion in Åland indicate glacial conditions of thick and fast, constantly sliding ice (e.g., Evans, 1996; Glasser & Bennett, 2004), which is likely associated with a thin clay rich ‘lubrication’ layer of former basinal sediments beneath the streaming glacier. Similar conditions are typical for numerous locations around the FIS and the former Laurentide Ice Sheet region (Bukhari et al., 2021; Eyles, 2012; Eyles et al., 2016; Krabbendam et al., 2016). In areas of homogeneus bedrock lithology, we conclude that the presence of pre-existing fractures and topographic depressions—as recognized in this study—are a result of localized, structurally controlled quarrying that took place under dominantly abrasive erosional conditions. This finding is in line with the conclusions that site-specific fracture patterns is the most important feature in controlling the bedrock erosion and associated bed morphology (Krabbendam et al., 2016; Woodard et al., 2021).

In a regional scale, bed morphology is the main controlling factor over the slip of hard-bedded glaciers (e.g., Eyles, 2012; Kamb, 1970), as it contributes to the resistance of the ice to flow across and around basal obstacles. Bingham et al. (2017) recognized that much of the variability in the basal traction of the West Antarctica glaciers relates to topographical variation at the ice bed in the scale of individual subglacial basins, and such variation may not be adequately constrained with the available geophysical methods. For the presently exposed glacier beds in Northern America and Fennoscandia, structurally controlled bedrock ridges have been shown to slow down or temporarily even stop the glacier as shown by the deposition of end moraines and glacifluvial deltas (Crossen, 1991; Skyttä et al., 2015). Such findings that even moderate (<100 m) vertical changes in the ice-bed topography contribute to the ice flow (Skyttä et al., 2015) further highlight the need for using structural geological approaches in addressing the topographical variations of the ice bed. This applies to understanding the major deformation zones (faults, shear zones and fracture zones) but also the spatial variability in fracture densities within the bedrock, as the latter has been shown to play a role in the development of surge-type glaciers (Crompton et al., 2018).