Extreme strain in the form of flattening or constriction during noncoaxial shear in ductile shear zones provides a record of ductile thrust system dynamics and the overall tectonic evolution. Within the Moine Nappe in northern Scotland, between the Ben Hope and Moine thrusts, the Strathan Conglomerate displays apparent strain partitioning with extreme flattening (e.g., laterally extensive sheets of deformed pebbles with aspect ratios of 134:113:1 and 88–92% estimated thinning) adjacent to the overlying Ben Hope Thrust and extreme constriction (e.g., rods with aspect ratios of 21:4:1 and estimated extension of 1000%) lower in the nappe package. We demonstrate that partitioning of strain is between its intensity and how deformation is manifested. Field, microstructural, and crystallographic orientation data from this study indicate that both areas were deformed by WNW-directed noncoaxial shear and coaxial flattening under amphibolite-facies conditions. Adjacent to the Ben Hope thrust, flattening was pervasive during nonco-axial shear, whereas beneath and within the Moine Nappe package, polyphase folding dominated. There, early, large-scale folds (F2) rotated into the transport direction. Subsequent transport-parallel (F3) folds and tubular sheath folds formed on the F2 limbs and were dismembered to form rods. No evidence of constriction is observed; instead, pervasive noncoaxial shear was accompanied by minor flattening under decreasing temperature conditions. Thus, these S-tectonites in the Moine Nappe are the result of concentrated flattening of pebbles into sheets during WNW-directed shear, whereas the L-tectonites result from heterogeneously distributed shear and folding, coupled with minor flattening, which produced rods without constriction.
Partitioning of strain along major mid-crustal thrusts, such as the Moine Thrust in northern Scotland, is key to understanding the dynamics involved in the formation and movement of ductile fault nappes that extend for hundreds of kilometers in length. Modeling the geodynamics of such thrust systems and determining the overall tectonic evolution requires knowing how strain is accommodated. The type of strain impacts estimates of displacement, thrust propagation rates, mechanisms of fold formation, estimates of temperature conditions during deformation (by modifying crystallographic preferred orientations [CPOs]), and many other factors. Deformation can be manifested in many ways (Fig. 1), pure and/or simple-shear plane strain or 3-D, non-coaxial general shear and/or coaxial flattening or constrictional strain.
Despite extensive work on the Moine Thrust zone, disagreement on the type of strain persists (e.g., Law et al., 1984, 1986; Alsop and Holdsworth, 2002; Strine and Wojtal, 2004; Lusk and Platt, 2020). For example, within the lower grade Moine Thrust zone, Law et al. (1984, 1986) report coaxial deformation within the thrust package and thinning from extensional flow and noncoaxial shear proximal to thrusts. For the Moine Thrust zone and Nappe, Thigpen et al. (2010) document foliation-normal shortening of 35–55%. Lusk and Platt (2020), however, argued for plane strain and dominantly simple shear with no volume loss or thinning above the Moine Thrust, based on quartz c-axis patterns and strain compatibility issues. Farther east, a vorticity study of greenschist-facies rocks along the Ben Hope Thrust also found that most samples support dominantly plane-strain simple shear (Graziani et al., 2021). Non-plane–strain deformation in shear zones results in strain compatibility issues (Fig. 1), although differential shear along the length of a shear zone can cause local flattening and constriction that accommodates the space problems (e.g., Coward and Kim, 1981).
The studies referenced above for the Moine and Ben Hope thrusts rely predominantly on micro-structures and/or CPO data to determine the type of strain. Another long-employed approach is to use deformed conglomerates to quantify the amount and type of strain and the structural evolution of a deformed region (e.g., Hossack, 1968; Siddans, 1983; Mosher, 1987; Moriyama and Wallis, 2002; Dasgupta et al., 2012). Within the Moine Nappe of northernmost Scotland near the Kyle of Tongue, deformation of the Strathan Conglomerate within the ductile thrust package beneath the Ben Hope Thrust apparently varies from extreme flattening to extreme constriction in outcrops ~3 km apart (map view) and provides the opportunity to explore how strain is partitioned and varies spatially in both intensity and type of deformation within a mid-crustal ductile thrust nappe.
The Strathan Conglomerate crops out beneath the Ben Hope Thrust at Strathan Bay (Mendum, 1976) and is repeated several times by folding and thrusting between Strathan Bay and Ben Hutig (Holdsworth, 1989; Fig. 2). In Strathan Bay, the pre-dominantly quartzite and feldspathic conglomerate pebbles have been deformed into thin, laterally extensive sheets (S-tectonites; Fig. 3). Mendum (1976) conducted a strain study on the conglomerate in Strathan Bay and calculated shortening or flattening normal to the foliation of 89–92%. He suggests that the unit was originally 400 m thick prior to shortening to its current 40 m thickness.
In contrast, on the summit of Ben Hutig, ~3 km to the west of Strathan Bay and structurally lower in the nappe package (350 m to 1.2 km lower, depending on projection of thrust; Fig. 2), the conglomerate pebbles crop out as transport-parallel rods (L-tectonites) or wavy lenses with an exaggerated length in comparison to the other two dimensions (Fig. 4). Wood (1973) interpreted the transport-parallel folds and rods present on Ben Hutig as being the result of constrictional strain, the classic method for L-tectonite formation. That study calculated extension of the pebbles in the x-direction on the order of 1000% and shortening in the y- and z-directions on the order of 70% and 75%, respectively. The mean deformation ellipsoid calculated by Wood (1973) is extremely prolate (25:1:0.9). The rods and lenses on Ben Hutig are spatially associated with, and parallel to, fold axes with the same WNW-trending thrust transport direction as that of the Moine Nappe (Peach et al., 1907; Wilson, 1953; Mendum, 1976).
In this paper, we present field, microstructural, and crystallographic orientation data that document the processes involved in the deformation of the Strathan Conglomerate at both localities to address the following questions. Why is the deformation of the conglomerate manifested so differently, despite the similar kinematic framework and close proximity? How is this difference recorded by the microstructures and CPOs? What type of strain is recorded at each locality? Are the rods the result of constrictional strain, and can they be used to measure strain? Lastly, we present a model for the formation of the structures in both areas and discuss the implications of our results for ductile thrusts and nappes.
TECTONIC AND STRUCTURAL SETTING
During the lower Paleozoic, the collision of Baltica with the Laurentian margin of East Greenland caused the Scandian orogeny (435–425 Ma) (Strachan et al., 2002, 2020). In northern Scotland, east-dipping ductile thrusts imbricated high-grade nappes and interleaved Archaean–Palaeoproterozoic Lewisianoid basement and Neoproterozoic Moine meta sedimentary rocks (Holdsworth et al., 1994). Scandian thrust nappes decrease in metamorphic grade westward from amphibolite to lower greenschist facies, and each nappe contains its own stratigraphic, deformation, and metamorphic history (see Barr et al., 1986; Strachan and Holdsworth, 1988; Alsop and Holdsworth, 1993; Alsop et al., 1996; Holdsworth et al., 1994, 2001, 2006, 2007; Strachan et al., 2002, 2010; Thigpen et al., 2010, 2013, 2021; Krabbendam et al., 2011; Mako et al., 2019; Law et al., 2021). Thrusts generally propagated westward, stacking nappes (Skinsdale, Swordly, Naver, and Moine nappes from east to west in northernmost mainland Scotland), which overrode the foreland along the Moine Thrust (Strachan et al., 2020, and references therein), forming a generally westward-propagating foreland thrust belt (e.g., McClay and Coward, 1981).
Within the Moine Nappe in northernmost Scotland, near the Kyle of Tongue, Scandian-aged thrusts strike roughly NNE and display a WNW transport direction (Holdsworth et al., 2001, and references therein). These thrusts overprint a bedding-parallel penetrative fabric associated with garnet-grade metamorphism (Holdsworth et al., 2001) (D1) that formed during the Neoproterozoic (Strachan et al., 2002). Structural mapping and analysis of multiple fold generations have demonstrated large-scale sheath folding (F2) associated with thrusting (Holdsworth, 1990; Alsop and Holdsworth, 2002, 2004). These authors documented asymmetric isoclinal F2 folds with hinges at low angles to the transport direction that show opposite vergence across a series of large-scale culminations and depressions. The F2 folds have curved fold hinges that they attribute to rotation into the transport direction (Fig. 5A) caused by differential flow, surging, or slackening of movement during overall WNW transport. Two types of minor F3 folds are recognized, which are closed to isoclinal folds with highly curved hinges or small-scale sheath folds (F3a) and open-to-closed folds oriented parallel or subparallel to the lineation that preclude rotation (F3b) because of their interlimb angle. The F3b folds are parallel to and associated with rods at Ben Hutig.
The metasedimentary rocks were metamorphosed to amphibolite-facies conditions, with Mg-Fe thermometry of garnet amphibolites yielding temperatures of 572° C+/−50° C on the flank of Ben Hutig in the footwall of the Achininver Thrust and 601° C +/−50° C ~750 m to the east, above the Ben Hope Thrust (at Port Vasgo) (Thigpen et al., 2013; Fig. 2). Eastward, the Ben Hope Thrust sheet contains higher grade metamorphic rocks, which reach 655 °C (Thigpen et al., 2013) at the thrust contact with the even higher grade rocks of the Naver Nappe (>700 °C; Ashley et al., 2015; Thigpen et al., 2021, and references therein). On Ben Hutig, Thigpen et al. (2013) estimated a deformation temperature of 535 °C using c-axis fabric-opening angles from a quartz rod but noted that this temperature may be underestimated because of reported constrictional strain at this site.
The Strathan Conglomerate is Neoproterozoic in age and unconformably deposited on Archaean–Paleoproterozoic Lewisian gneisses (Mendum, 1976; Holdsworth, 1989; Holdsworth et al., 2001), which crop out locally in Strathan Bay. The conglomerate pebbles are derived primarily from the Lewisian basement (Holdsworth et al., 2001) with quartzite, psammite, vein quartz, pegmatite, and feldspathic lithologies (Mendum, 1976). This basal conglomerate is overlain by psammites and pelites that contain cross-bedding, convolute bedding, channel structures, and graded bedding, which indicates deposition in shallow water (Holdsworth, 1989, 1990; Holdsworth et al., 2001). In Strathan Bay, the conglomerate is located immediately below the Ben Hope Thrust, and the unit is ~40 m thick. On Ben Hutig, the conglomerate is thinner (0.5–1 m) and interbedded with equivalent layers of psammite.
Wilson (1953) considered the rods at Ben Hutig to be folded and dismembered quartz veins, and Mendum (1976), although recognizing conglomerate at Ben Hutig, viewed most rods as quartz veins. The presence of Strathan Conglomerate on Ben Hutig was confirmed and mapped by Holdsworth (1989) and Holdsworth et al. (2001) and by this study. Although numerous quartz veins are observed on Ben Hutig, most are late crosscutting all structures or associated with a late-stage, upright F4 fold. Some veins are parallel to the S2 foliation and folded by F3; however, the most pronounced veins (Fig. 4A; e.g., fig. 8 of Wilson, 1953) cut the foliation for much of their length and can be readily distinguished from pebbles (see Field Observations section). Quartz veins indicate some fluid-assisted, diffusive mass transfer during deformation; however, the pebbles themselves do not preserve features characteristic of pressure solution in the field or thin section (cf. Mosher, 1976, 1978, 1981).
Field observations, structural data, and oriented samples were collected and analyzed at Strathan Bay and Ben Hutig with a focus on the conglomerate, and structural maps were constructed at scales of 1:2500 and 1:1250, respectively (Collier, 2006). At Strathan Bay, mapping was restricted to the footwall side of the bay, where the Strathan Conglomerate crops out. At Ben Hutig, mapping was restricted to the zones of outcrop, which occur along the southern ridge and the four knolls surrounding the summit. Axial ratios were also calculated for the pebbles at Strathan Bay and Ben Hutig. (Note that the original size of the pebbles is unknown and may range from pebbles to cobbles or even boulders in some cases; we will refer to these as pebbles throughout.) These were determined by measuring the three dimensions of the pebbles on perpendicular joint faces at a total of 10 locations, half at Strathan Bay and half at Ben Hutig. Approximately 30 pebbles were measured and averaged at each location. No measurements of deformed veins were made.
Oriented samples taken from Strathan Bay and Ben Hutig were sectioned to examine the fabrics and textures. For each sample collected, two orthogonal-oriented thin sections were cut, one parallel and one perpendicular to the lineation or fold axes, and both perpendicular to the foliation, for a total of 35 thin sections for Strathan Bay and 60 thin sections for Ben Hutig. When not parallel, lineations on pebble surfaces that record internal pebble deformation were preferred over mineral lineations within foliation. In addition, 71 thin sections from non-oriented samples collected during preliminary fieldwork by Mosher (1978) were made in three perpendicular orientations: perpendicular to the foliation, either parallel or perpendicular to the lineation or fold axes, and parallel to the foliation and lineation or fold axes.
Five representative thin sections from each location were prepared for electron backscatter diffraction (EBSD) analysis: from Strathan Bay, two thin sections (SB2E and SB3A) adjacent to the Ben Hope Thrust and three thin sections (SB5E, SB6A, and SB7E) from beneath it; from Ben Hutig, one thin section (BH1E) from outcrop with very thin and laterally extensive sheets, one thin section (BH10A) from a quartz rod, two thin sections (BH14BA and BH14CA) from F3 hinges, and one thin section (BH18A) from an F3 limb. Thin sections ending in “A” are lineation-parallel, and those ending in “E” are lineation-perpendicular. Data are plotted in planes other than the one in which they were collected by 90° rotations around the X, Y, and Z axes. EBSD data were collected using an Oxford Instruments Nordlys nano detector on a JEOL 6490LV SEM at the University of Texas Geomaterials Characterization and Imaging Facility in the Department of Geological Sciences. EBSD data were collected at 15 kV and 2–5 nA using a focused beam from uncoated samples with a chamber pressure of 20 Pa. Map step size ranged between 2 µm and 3.5 µm for maps focused on quartz, and 10 µm for maps focused on feldspar. Step sizes were chosen based on qualitative estimates of grain sizes from optical petrography to ensure that each grain was measured more than three times. Diffraction patterns were collected and indexed using 2×2 binning and Refined Accuracy in Aztec 3.1 and cleaned according to the best practices of Prior et al. (2009). All post-processing was done using the MATLAB toolbox MTEX 5.7.0. Orientation data were smoothed following the procedure described in Hielscher et al. (2019). Orientation distribution function half-widths were set to 8° for inter-sample comparability. The misorientation index (M-index) of fabric strength based on the distribution of uncorrelated misorientation angles was calculated for quartz orientation distributions according to Skemer et al. (2005). Crystallographic vorticity axes (CVA) were calculated in accordance with Michels et al. (2015) and plotted in the structural XZ plane. Quartz opening angle thermometry was calculated using an assumed pressure of 5.5 kbar (see Law et al., 2021, and Thigpen et al., 2021, for a summary of regional pressure-temperature work) applied to the thermometer of Faleiros et al. (2016).
FIELD OBSERVATIONS AND DISCUSSION
In Strathan Bay, the Strathan Conglomerate lies immediately below the Ben Hope Thrust and is dominated by S-tectonites. The thick, massive conglomerate layers contain pebbles that have been deformed into very thin and laterally extensive sheets that record variable amounts of strain (Fig. 3). Ben Hutig is located ~3 km west of the Ben Hope Thrust in Strathan Bay, and thinner, interlayered conglomerate and psammite are exposed structurally lower in the nappe package (Fig. 2). Pebbles are dominated by pronounced L-tectonites and folds, rather than the extreme S-tectonites of Strathan Bay. Although both areas display the same structural elements, they are manifested differently in the field. Below, we describe in detail the structures and relationships in the two areas. Although we only observed Scandian age deformation, we will use the established nomenclature for structural generations for this portion of northern Scotland (cf. Holdsworth, 1990) rather than that of Collier (2006).
Strathan Bay Field Observations
In Strathan Bay, quartzite, granite, and feldspathic quartzite pebbles have been deformed into very thin and laterally extensive sheets (Fig. 3) that define an east-dipping, very well developed planar fabric, S2, which is generally parallel to bedding (S0) (Figs. 6A and 6B; Table 1). Aligned micas and planar quartz grains define S2 in the psammite. A well-developed and constantly oriented lineation, L2, is present within the S2 fabric (Figs. 3D, 3E, 6A, and 6B; Table 1). L2 is defined by aligned elongate micas in the psammite and between conglomerate pebbles and nearly parallel fine-scale (≤ 1mm) rodding on pebble surfaces that is oriented 5–10° from the mineral lineation.
The extent to which the pebbles of all compositions have been deformed into sheets varies along the length of the thrust and perpendicular to the sheets. Along the length of the thrust, deformation varies from zones of low strain (28:15:1 axial ratio) to zones of very high strain (134:113:1 axial ratio; Figs. 2 and 3). Perpendicular to the sheets, high-strain and low-strain zones are transitional and gradually change in strain intensity.
Only a few F3 and F4 folds, and no F2 folds, were seen in Strathan Bay. The rare F3 folds deform bedding, S2 (including planar conglomerate pebbles), and L2. The orientation of F3 axes varies along the length of the bay, possibly because of isolated sheath folding or late-stage broad folds (F4) (Figs. 6A and 6C; Table 1). A weakly developed axial planar foliation, S3, defined by micas, is present in some of the larger folds.
Rare, isolated, late-stage upright gentle to open folds and crenulations in Strathan Bay, classified as F4 folds, differ from F3 in style but are not directly seen to fold F3 axes or axial planes. Given the scale of these features, and the observation that the fold axes generally lie within the average S0/S2 plane, it is most likely that the orientation of the anisotropy is the primary control on their orientation (Figs. 6A and 6D; Table 1).
Ben Hutig Field Observations
Polyphase folding of the interlayered conglomerate and psammite is the dominant structure on Ben Hutig (Fig. 7). Interlayered white, dark, and pale gray quartzite, granite, and feldspathic quartzite pebbles form pronounced L-tectonites parallel to F3 fold axes (Fig. 4) rather than the extreme S-tectonites of Strathan Bay.
S2 observed in the conglomerate of Ben Hutig is much different from that at Strathan Bay, though it is the same in the psammite. The pebbles have a lens to elliptical shape when viewed on surfaces perpendicular to F2, regardless of their orientations parallel or oblique to bedding (Figs. 7A and 7B), and have a significantly longer dimension (24:4:1 axial ratio) when viewed parallel to F2 axes (Figs. 2 and 4). The intermediate and long dimensions of the pebbles are well aligned and define S2 (Figs. 4B, 7A, and 7B). As at Strathan Bay, a similarly defined, well-developed L2 lineation is present within the S2 fabric (Fig. 4D, inset in panel F), though no difference in orientation is noted between the mineral and rodding lineation.
Unlike Strathan Bay, evidence for large-scale, isoclinal F2 folds is common. S0 and S2 are generally parallel (Figs. 7A and 8A), which is indicative of isoclinal fold limbs, but locally they are oriented at a high angle (Figs. 7B–7D) in fold-hinge regions. Most F2 axes (measured by the intersection of S0 and S2) are parallel to the L2 lineation (Figs. 8A and 9A; Table 1).
Abundant F3 folds with axes parallel to the long dimension of the pebbles range in scale from centimeters to tens of meters in amplitude and wavelength. (Figs. 4B, 4C, 7C, and 7D). F3 folds are inclined to overturned or recumbent, open to closed, with some locally tight. S0, F2 folds, S2, and L2 are folded by F3 (Figs. 4B, 4C, 4F inset, 7C, 7D, and 9A–9C), and interaction between larger-scale F2 and F3 folds is observed (Figs. 7C and 7D). Some F3 folds have sinuous axes that varied in trend by a maximum of ~10° (Figs. 4D and 4F inset). A moderately well-developed axial planar foliation, S3, is locally observed. The S3 foliation is defined by aligned micas in the psammite and between conglomerate pebbles.
A few centimeter-scale F3 folds have a sheath fold geometry in both the conglomerate and the psammite as viewed on NNW-striking joint surfaces (Figs. 10B–10D). Sheath fold noses generally are not exposed, however, though their SE-dipping axial planes are compatible with a westward direction for the hinges (Figs. 8B and 9B; Table 1). One N-trending isoclinal fold of S0 and S2 with a 2 m amplitude was interpreted as a possible hinge zone of a larger sheath fold.
One late stage, upright, gentle F4 fold, ~20 m in wavelength, and associated parasitic folds and crenulations, was mapped on Ben Hutig (Figs. 8B and 9D). In outcrop, this F4 fold shows essentially no change in bedding thickness, has undeformed quartz veins parallel to its axial plane, folds S0 and S2, and crenulates S2 and S3.
The difference between pebble shapes between Strathan Bay and Ben Hutig is striking. Only one isolated section (30 m2) in the northeast corner of Ben Hutig contains very thin and laterally extensive sheets of pebbles that are not folded, similar to those in Strathan Bay. Elsewhere, pebbles form lens shapes that are elongate in the third dimension parallel to F3, which folds them to varying degrees, or rods parallel to F3 axes.
The lens-shaped pebbles, when viewed perpendicular to their long dimension, show pinch-and-swell features or boudinage. Many lensshaped pebbles taper on both ends, elliptical pebbles commonly show thin, folded tails, and some form near perfect cylinders (rods, see below; Fig. 4). Some pebbles are sheared along small-scale shear zones parallel to horizontal bedding between psammite and conglomerate, with an apparent top-to-the-SW motion (Fig. 7B). Elsewhere, pebble tails are deflected by small-scale shear zones oblique to the F3 axes (Fig. 10A). These features, coupled with the one isolated outcrop of sheeted pebbles, suggest that some of the conglomerate may have undergone a similar deformation path to Strathan Bay and formed sheets before being dismembered. Not all of the pebbles were once laterally continuous sheets, however, because where S2 is oblique to bedding near F2 fold hinges and thin pebble layers are interlayered with psammite, pebbles viewed perpendicular to the transport direction have lens shapes that clearly terminate at the layer boundary (Fig. 7A). (Note that the pure white quartzite pebbles are interlayered with gray quartzite and pink feldspathic quartzite pebbles and have the same lens shapes and dimensions; they do not cut across the psammite bed as would be expected if these were veins.)
Rods aligned approximately parallel to F3 axes are a dominant feature on Ben Hutig. Rods occur in clusters or are interspersed with the folded, lens-shaped pebbles on both limbs and hinges of larger F3 folds. Most rods are near-perfect cylinders or slightly elliptical cylinders, but a few display tails, which is reminiscent of highly attenuated layers or limbs (Figs. 4D–4F, 10A, and 10D). In some locations, the rods and tails were associated with small shear zones nearly perpendicular to the average orientation of the rods (Figs. 9B and 10A). Rods are not folded by F3 folds; however, some of the rods contain a twisted or corkscrewed interior fabric seen when cutting slabs during thin-section preparation. Similar to F3 axes, rods are locally sinuous along their length, with about a 10° swing in the trend of their long axes (Fig. 4D, F inset), and are wrapped by the L2 lineation. As noted previously, locally, quartz veins have been deformed into rods, but most rods in the areas mapped for this study are conglomerate pebbles (Figs. 4D–4F). For example, Figure 4E shows rods and dismembered F3 folds of quartzites that are clearly confined to a conglomerate layer, and none extend into the adjacent psammite as would be expected if these were folded veins.
Most planar elements (S0, S2, S3, sheath fold planes, and shear zones) generally fall on a NNW-striking great circle dipping ~70° W, which suggests folding on an axis plunging 20° ENE (Figs. 9A–9C). Map and stereonet analysis indicate that the large, upright F4, despite its axis being in an appropriate orientation, was not responsible for the dispersion of F3/S3 (Fig. 9C). Although locally this F4 reoriented S0 and S2, most of the change in orientation across the area is the result of F3 folding. All linear elements (F2 axes, L2, F3 axes, F3 pebble axes, and rods) primarily lie along an ~N-striking great circle and have a general plunge of 15–20° toward the NE, E, and SE. (See Figs. 8 and 9 and Table 1 for specifics.)
Discussion of Field Relations
Below the Ben Hope Thrust at Strathan Bay, the presence of conglomerate pebbles forming sheets demonstrates that flattening has occurred, and the variable degree of strain perpendicular to the thrust indicates that flattening was accompanied by shearing. The widths of the sheets perpendicular to the transport direction (~1–2 m at high strains) are too large to represent the original pebble dimensions (see Figs. 3D and 3E). The lineation defined by aligned micas and very fine scale (≤1 mm) rods on the pebble surfaces, and the slight difference between them, most likely reflects variation in movement over time.
The exposure of conglomerates at the northern end of Ben Hutig with the same sheet-like shapes as at Strathan Bay suggests that much of the conglomerate on Ben Hutig was first deformed in a manner similar to the conglomerate near the Ben Hope Thrust, but most likely at low–medium strains based on comparison with those at Strathan Bay. The pinch and swell structures, elongate lens-shaped clasts, boudinage, and tails off many rods also support modification of low–moderate, strain sheet-shaped conglomerate pebbles. With the somewhat lower temperatures at Ben Hutig compared with Strathan Bay (noted by Thigpen et al., 2013), the difference in rheology of pebbles (quartzite versus granite) would be more pronounced. The few locations within F2 fold hinges, where pebbles in thin, interbedded pebble and psammite layers are oblique to bedding, indicate that not all pebbles were deformed into extensive sheets and dismembered.
The spatial association of the F3 fold axes and rods, which trend parallel to the WNW transport direction, presence of tails on the ends of rods, and the few small, parallel sheath folds indicate that the rods evolved from or at the same time as the F3 folds. The L2 lineation is folded by F3 folds and is wrapped around the rods, which indicates rotation during formation. Small-scale shear zones oblique to the transport direction also indicate potential dismemberment of sheets and F3 folds that contributes to the twisting of these linear structures. The spread of F3 axes within a shallow, east-dipping plane is compatible with an overall tubular sheath fold geometry for F3.
The irregularities within the conglomerate (e.g., boudins, lenses, and pebbles with a different rheology) provided perturbations on which sheath folds could nucleate, as described by Cobbold and Quinquis (1980). The extreme length of the rods and F3 fold axes and predominantly open-to-closed, fold-interlimb angles, however, are more compatible with the formation of tubular sheath folds, where folds form in that orientation on variably oriented, preexisting F2 limbs, as described by Skjernaa (1989), Mies (1991) and Boettcher and Mosher (1998; Fig. 5B). The alternative explanation, that during shearing shortening occurred perpendicular to the transport direction and parallel to the shear plane (i.e., constrictional strain; e.g., Fletcher and Bartley, 1994), is unlikely because of recumbent nature of many F3 and the shallow dips of nearly all F3 axial planes and S3.
The detailed structural analysis described in this paper is in overall agreement with the regional interpretations of Holdsworth (1989, 1990) and Alsop and Holdsworth (1993, 2002) for the study area. We observe isoclinal F2 folds with a swing of axes from SE- to ENE-plunging, which is compatible with rotation into the transport direction as they propose. One of their fold culminations projects into the Ben Hutig area and potentially could be our large-scale F4 shown in Figures 8B and 9D. We do not, however, observe any change in dip direction of the fold axial planes across this axis as would be predicted, and the spread of F3 axes within a shallow, east-dipping plane is demonstrably unrelated to the larger-scale F4 fold structure.
The F3 axes, rods, and few sheath fold axes trend parallel to the transport direction; however, as discussed above, their generally open-to-closed, interlimb angles preclude significant rotation, which was also noted by Alsop and Holdsworth (2002). We propose that as the F3 folds formed, some were initially somewhat oblique to the shear direction and were rotated or twisted around their axes.
The strain intensity determined by Lusk and Platt (2020; p. 6) for metaconglomerates in this general area yield a strain intensity (D = [(ln (X/Y)2 + ln (Y/Z)2]1/2) of between 2.5 and 2.3. Our strain analysis is compatible overall with these results, yielding intensity values of 3.13–2.58 at Strathan Bay and 2.27 at Ben Hutig.
Microstructural and Crystallographic Data Observations and Discussion
Strathan Bay Conglomerate
The S2 foliation is defined by the orientation of the planar pebbles and large micas parallel in the matrix and within pebbles. Most micas within pebbles are fine-grained (<20 µm) and completely enclosed within individual quartz grains (Fig. 11A), which is indicative of high-temperature grain boundary migration (GBM) recrystallization (GBM II microstructures of Stipp et al., 2002a, 2002b).
In quartzites and feldspathic quartzites, quartz forms coarse-grained (250 µm to 1.75 mm), amoeboid-shaped grains with highly irregular boundaries (Figs. 11A, 11B, and 11D). These irregular-shaped grains were observed to be oblate within the foliation plane in some thin sections cut parallel and perpendicular to lineation and parallel to the foliation. In higher strain areas, quartz is finer grained (~200 µm). A c-axis preferred orientation is well developed in the quartz grains (Fig. 11A and see below). Overall, both types of quartzites are completely recrystallized and locally exhibit recovery microstructures. Subgrains and new, recrystallized grains are equant with similar shapes and sizes in most sections, indicating the local operation of subgrain rotation recrystallization (SGR; Fig. 11B). Patches of new grains share similar extinction positions in cross-polarized light that appear to outline larger preexisting grains (Figs. 11A, B).
Granitic pebbles are composed of equal proportions of dynamically recrystallized quartz and feldspar with micas included within the pebbles and some quartz grains (Fig. 11C). Recrystallized quartz and feldspar grains are usually of similar sizes (~100 µm) and equant, or at higher strains are elongate parallel to the foliation. Subgrains are rare, but they are present in both quartz and feldspar.
The feldspathic quartzite pebbles (10–15% feldspar; Fig. 11D) contain patches of completely recrystallized, fine-grained feldspar (<13 µm), similar to the granitic pebbles. A few feldspar grains have subgrains, and some patches of new feldspar grains share similar extinction and define larger preexisting grains, which suggests recrystallization by SGR. In highly strained pebbles, the quartz forms elongate bands between patches of fine-grained feldspar (50–100 µm) that vary from irregular to very elongate parallel to S2 (Fig. 11D). The overall appearance of these pebbles resembles high-temperature striped gneisses (see, e.g., Hippertt et al., 2001). In hand sample, the feldspar forms pink streaks within the pale gray quartz.
Most quartz grains in Strathan Bay pebbles, however, are equigranular and do not have an elongate shape parallel to the pebbles, a relationship that might be expected from the extremely flattened pebble shape. Locally, new, elongate SGR quartz grains and elongate subgrains are either oriented oblique to or perpendicular to the length of the pebbles. Obliquely oriented grains tilted to the NW and dipping ~20° to 45° SE, consistent with a top-to-the-NW shear sense, are observed in some sections cut parallel to the lineation. Elongate recrystallized grains and subgrains oriented perpendicular to S2 are usually restricted to the hinges of the few F3 folds. These grains define an axial planar S3 and are best seen in thin-sections oriented perpendicular to the lineation and F3 axes. Random or multiple orientations of elongate subgrains and recrystallized grains that formed by SGR, none of which are parallel to the pebbles, are observed in a few thin sections perpendicular to the lineation. Matrix micas in the hinges of F3 folds are recrystallized crenulations and are either roughly parallel to the recrystallized quartz grains, helping to define the S3 foliation, or oriented sub-parallel to the margins of the pebbles.
Ben Hutig Conglomerate
Microstructures and textures are similar to those at Strathan Bay, but with some important differences. Overall, the quartz in quartzite and feldspathic quartzite pebbles ranges from 125 µm to 625 µm in the coarse-grained pebbles and 25–125 µm in the finer grained pebbles. The change in grain size appears to be related to the percentage of mica present between grains, with far fewer micas located within grains, which suggests lower temperatures and GBM I recrystallization (Stipp et al., 2002a, 2002b). Coarse-grained, amoeboid-shaped quartz grains parallel or somewhat oblique to the foliation were observed in thin sections cut parallel to the F3 fold axes and rods. Overprinting these were subgrains, recrystallized grains produced by SGR, and discontinuous undulatory extinction bands oriented at up to 60° to the foliation, and in a few cases they were perpendicular (Fig. 12C). In the feldspathic quartzites, the patches of fine-grained feldspar occur both as irregular masses and locally as strips parallel to S2 when viewed in thin sections cut parallel to F3 axes and rods (Fig. 12C). The quartz crystal preferred orientation is stronger at Ben Hutig than it is at Strathan Bay (see below).
Granitic pebbles at Ben Hutig are less common than they are at Strathan Bay, but have similar textures with generally equant, intermixed quartz and feldspar. Unlike granite pebbles at Strathan Bay, subgrains are fairly common in both quartz and feldspar, along with discontinuous undulatory extinction. In the feldspathic quartzite pebbles at Ben Hutig, both quartz and feldspar grains also show continuous undulatory extinction, and feldspar has fewer subgrains. The black and dark gray quartzite pebbles have abundant very fine-grained heavy minerals (some magnetite or ilmenite, hematite, titanite, zircon, and other very fine, high-relief grains) that commonly decorate grain boundaries. Locally, heavy minerals appear to define an original bedding within the quartzite, and the quartz grain size is much finer.
The largest differences at Ben Hutig from Strathan Bay are observed in thin sections cut perpendicular to the F3 axes, where abundant elongate quartz grains recrystallized by subgrain rotation are oriented mostly parallel to F3 axial planes, defining S3 (Figs. 12A, 12B, and 13). Moderately abundant elongate subgrains and bands of discontinuous undulatory extinction within quartz parallel these recrystallized grains (Fig. 12A). Patches of these grains share similar extinction positions, defining preexisting larger grains, especially in the finer-grained quartzites. Elongate, recrystallized quartz grain-shape fabrics, which are usually perpendicular to the pebbles in the hinges (and axial planar), are locally a little more parallel to one limb or slightly askew (Figs. 12A and 12B). In individual folds, elongate grains show fanning, and folds of pebbles in adjacent layers and along the same layer show somewhat different orientations of elongate grains (Fig. 13A). In the limbs of F3 folds, the recrystallized quartz is oblique to the pebble boundaries and in places changes direction, almost defining a fold within the pebbles. In one refolded fold, the elongate grains defining S3 cut across the axial planes of an earlier fold (Fig. 13B). In feldspathic quartzites, the patches of feldspar grains locally form elongate lenses parallel to elongate quartz within F3 hinges. Micas in the matrix are aligned parallel to S3 or form recrystallized crenulations around the outer arc of the folds (Fig. 13).
Several thin sections cut parallel to L2 show some degree of folding perpendicular to the F3 axes. These folds have the same characteristics of F3 folds, folding both pebbles and matrix and having elongate, recrystallized quartz parallel to axial planes that is locally parallel to one limb. Locally, however, the elongate, recrystallized quartz grainshape fabric is seen completely folded by these folds. These folds provide further evidence of progressive sheath folding on the small F3 scale, even though the sheath itself was not observed. In other slides cut parallel to fold axes, elongate, recrystallized quartz and micas are parallel to the axis and pebble, further defining S3.
Rods examined in thin sections cut perpendicular to the lineation or fold axes appear as roughly circular groups of grains (Fig. 13A). The grains are usually elongate, but the orientation of elongation is variable within and between rods, which suggests they moved independently, with generally less pronounced elongation in the centers (Fig. 13A). Mica inclusions are limited inside the rods, and rods usually appear to be somewhat wrapped by the micas in the matrix around them. In most cases, where tails are present on the rods, the elongate quartz in the tail is parallel to the tail itself and to the micas immediately surrounding it. Where a tail is pinched off, the quartz is more equant. When viewed in thin sections cut parallel to the lineation or fold axes, the quartz grains maintain a fairly well-developed elongation. In this view, the elongation is usually oriented at 45° to the pebble and dips to the southeast (Fig. 12C), but a few pebbles show additional elongation orientations, including parallel to the rod.
Crystallographic Data for Strathan Bay and Ben Hutig
EBSD data were collected from five sections cut perpendicular to foliation at Strathan Bay and five at Ben Hutig and are summarized in Table 2. At Strathan Bay, two samples were cut parallel to lineation (XZ plane; SB3A and SB6A) and three were cut perpendicular to lineation (YZ plane; SB2E, SB5E, and SB7E); at Ben Hutig, four samples were cut parallel to lineation (XZ plane; BH10, BH14BA, BH14CA, and BH18A) and one perpendicular to lineation (YZ plane; BH1E). Quartz grains measured from the XY and YZ planes are equally elongate parallel to the foliation, with aspect ratios of 1.94 ± 0.42 (Strathan Bay) and 2.03 ± 0.31 (Ben Hutig). The misorientation index (M-index) values for crystal fabric strength based on the distribution of uncorrelated misorientation angles (Skemer et al., 2005) are relatively weak for Strathan Bay and an order of magnitude higher for Ben Hutig quartzites (Table 2).
When all crystal data for both areas are rotated into the XZ plane, asymmetric quartz c-axis girdles and a-axis point maxima oriented perpendicular to the girdle indicate a top-to-the-WNW shear sense associated with rhomb <a>, prism <a>, and basal <a> slip (Fig. 14). Nearly all of the quartz data indicate plane-strain conditions and demonstrate that the observed lineation is parallel to the maximum principal stretch direction. Quartz c-axes from a sample taken from beneath the Ben Hope Thrust (SB2E) at Strathan Bay, however, form a very diffuse small circle that is characteristic of flattening with an apparent opening angle of 75° and an estimated deformation temperature of 570 °C, using the thermometer of Faleiros et al. (2016). Another sample from beneath the thrust (SB6A) has a quartz c-axis, Type I cross girdle, which is typical of plane strain, but with a distinct external fabric asymmetry girdle indicating a top-to-the-WNW shear sense. The quartz fabric opening angle of 65° indicates a deformation temperature of 516 °C. Quartz c-axes from three Ben Hutig samples (BH10A, BH14BA, and BH14CA) have remnants of small circle girdles, which when reconstructed suggest apparent opening angles of 57°, 66°, and 49°, and temperatures of 466 °C, 523 °C, and 399 °C, respectively, though this interpretation is highly speculative.
Quartz grain orientation spread (GOS) values and average first-order, kernelaveraged misorientation per grain (gKam) values for the quartzites at Strathan Bay are lower than those at Ben Hutig (Table 2), which suggests that quartz grains display little intragranular deformation at Strathan Bay and more at Ben Hutig. At Ben Hutig, SGR is supported by the steeper lattice curvature gradients represented by the average gKam values, a positive deviation in high-angle grain boundaries with <50° misorientations compared to an untextured quartz polycrystal (Fig. 15D), and by numerous observed subgrain boundaries parallel to recrystallized grain boundaries. Strathan Bay quartzites also show an increased number of high-angle grain boundaries with a <50° misorientation, but to a much lesser extent (Fig. 15C).
At both localities, a striking difference occurs between the quartz subgrain misorientation axes and high-angle grain boundary misorientation axes. Subgrain misorientation axes indicate that prism <a> slip caused the formation of those boundaries (e.g., Figs. 15A, 15B, and 15E). If most quartz grains were the result of SGR, where subgrain boundaries are progressively rotated via prism <a> dislocations into high-angle boundaries, the resulting high-angle grain boundaries should show the same misorientation axis. In contrast, Strathan Bay quartz lattices correlated across most high-angle grain boundaries have dominantly rhomb <a> and the uncommon slip system pyramidal <c+a> misorientation axes, and in some cases minor basal <a>, prism <a>, and/or prism <c> (Fig. 15A). Thus, most high-angle grain boundaries are not the product of SGR and prism <a> slip and may be unrelated to slip on these systems. The Ben Hutig quartz lattices show the same misorientation axes, but with more prism <a> than is observed at Strathan Bay, which indicates that more grains are the product of SGR (Fig. 15B).
We also conducted CVA analysis, which uses EBSD data to determine the local grain-scale kinematic vorticity axis for a given sample based on the change in orientation of individual crystal lattices during shearing (Michels et al., 2015). When the rotation axes of all grains measured within a sample are plotted together, their relationship to the foliation and lineation are analyzed. Vorticity axes that lie in the plane of the foliation record simple shear (progressive noncoaxial deformation) if the axes are perpendicular to the lineation and record pure shear (progressive coaxial deformation) if the axes are parallel to the lineation; if oblique, they record triclinic shear (Kruckenberg et al., 2019, and references therein). These grain-scale vorticity axes are assumed to be rapidly reset and represent the last deformation affecting the lattice orientations (Kruckenberg et al., 2019).
For Strathan Bay samples, the vorticity axes for quartz and feldspar generally do not lie in the foliation plane, which suggests an intermediate condition between progressive coaxial and non-coaxial shear strain (Fig. 16). Only the SB3A and SB6A quartz vorticity axis maxima lie roughly within the foliation plane. SB3A maxima are primarily at a high angle to the lineation, which implies noncoaxial shear, whereas the primary SB6A maximum is nearly parallel to the lineation, which implies coaxial deformation. SB5E quartz and feldspar also show maxima on the foliation plane nearly parallel to the lineation, which implies coaxial deformation, although the data are widely spread across the stereonet. SB2E quartz, SB6A feldspar, and SB7E quartz and feldspar maxima lie primarily off the foliation plane but are closer to being parallel to the lineation, which indicates a component of coaxial deformation. Note that SB6A quartz axes differ from the feldspar axes, indicating decoupling of the feldspar and quartz during deformation. We interpret the complex vorticity plots for the Strathan Bay samples to show that the lattice orientations record coaxial flattening during non-coaxial shear.
CVA results for Ben Hutig samples are similar to those of Strathan Bay samples in not recording only coaxial or noncoaxial deformation, although overall they indicate more noncoaxial shear (Fig. 16). BH10A and BH14CA quartz vorticity axis maxima are roughly within the foliation plane and oblique to the lineation, which implies noncoaxial shear. BH14BA has two distinct quartz vorticity axis maxima, one within the foliation nearly perpendicular to the lineation d the pinning of grain boundaries, particularly in the black to dark gray quartzite pebbles, is indicative of GBM I recrystallization. This type of recrystallization and overall smaller grain sizes indicates somewhat lower temperatures (500–550 °C; Stipp et al., 2002a, 2002b) than at Strathan Bay, though presence of some SGR in feldspar indicates temperatures at the high end of that range. Similar to Strathan Bay, the strong quartz preferred orientation is compatible with a top-to-the-WNW shear sense and slip on rhomb <a>, prism <a>, and basal <a> slip planes (Fig. 14B). The quartz c-axis plots for samples of F3 fold hinges and a rod (BH10A, BH12BA, and BH14CA), however, show remnants of small circles with relatively large apparent opening angles (49–66°), which may record early flattening that has been significantly overprinted by subsequent noncoaxial shear. The sample from the sheeted pebble outcrop and from an F3 fold limb (BH1E and BH18A) only show evidence of WNW-directed, noncoaxial shear. The CVA analyses also support dominantly noncoaxial shear but with a coaxial component; the complexity of these vorticity axis plots may reflect the folding and twisting of these fabrics.
The dominance of subgrains and new SGR grains axial planar to F3 folds provide evidence of protracted SGR recrystallization during D3 acting on the larger grains that formed during earlier GBM I. Elongate SGR quartz grains oriented axial planar to recumbent and inclined F3 folds indicate a component of flattening during D3. Axis rotation or corkscrewing during D3 is also indicated by the slight offset of axial planar, elongate quartz grains in several F3 hinges, fanning of S3 within individual folds and between folds and rods in adjacent layers, refolded folds, and progressive folding in perpendicular orientations of elongate quartz grains. This progressive twisting and folding is supported by the presence of discontinuous and continuous undulatory extinction in F3 hinges and within rods, which is not observed at Strathan Bay. Because the rods share characteristics of the F3 folds, it is inferred that they formed during D3 as well. The dominance of SGR microstructures indicates that F3/S3 deformation was more pervasive and at a lower temperature, though the recrystallized grain size indicates temperatures at the higher end of the SGR temperature range (~490–535 °C). The lower temperatures are compatible with decreasing temperatures traced westward toward the Moine Thrust (Fig. 2) and the 535 °C estimate by Thigpen et al. (2013) using a Ben Hutig quartz rod, c-axis opening angle. Unlike at Strathan Bay, that deformation continued as temperatures decreased is further indicated by subgrains and discontinuous and continuous undulatory extinction. Also, grain boundaries of quartz and feldspar previously recrystallized by SGR are irregular and embayed, which indicates bulging recrystallization at low temperatures (300–400 °C; Stipp et al., 2002a, 2002b). Similar to the microstructures recorded at Strathan Bay, micas on Ben Hutig were also recrystallized.
DISCUSSION OF STRAIN, STRAIN PARTITIONING, AND ROD FORMATION
The type of strain at each locality is best documented by the field relations, as discussed previously. At Strathan Bay, the conglomerate pebbles underwent extreme flattening and formed sheets with variable shear strain perpendicular to the plane of flattening. This 3-D strain requires both thinning of the ductile thrust zone and extension both parallel and perpendicular to the transport direction and impacts any estimate of translation along the thrust. At Ben Hutig, a single location of sheet-like pebbles similar to that at Strathan Bay, plus pinch and swell structures, boudinage, and dismemberment of sheet-like pebbles across the mapped area, suggest a similar initial flattening deformation but at lower strains during the formation of S2 and rotation of F2 into the transport direction. The locally varying orientation of the anisotropy (S0/S2) at angles to the shear plane resulted in the formation of tubular folds as well as sheath folds caused by the perturbations from the dismembered pebbles. The open-to-closed fold interlimb angles indicate these transport parallel F3 folds did not rotate into this position, and recumbent and shallowly dipping axial planes demonstrate that these F3 folds were not the result of constriction. The rods in the area mapped are clearly dismembered F3 hinges and thus also did not form due to constriction and cannot be used to estimate 3-D strain as was done by Wood (1973).
In terms of microstructures, the primary difference observed is GBM II at Strathan Bay and GBM I at Ben Hutig, which indicate higher temperatures near the Ben Hope Thrust, and a stronger overprint of SGR and incomplete recovery at Ben Hutig, resulting from pervasive F3 folding and rod formation and continued deformation as temperatures decreased. This continued noncoaxial shear resulted in an order-of-magnitude higher quartz misorientation index at Ben Hutig (M index) than at Strathan Bay.
Quartz crystallographic data from the two areas are remarkably similar, and the same slip systems are inferred to be active in both areas. All CPOs indicate WNW-directed shear, except for one Strathan Bay sample taken adjacent and immediately beneath the Ben Hope Thrust that indicates flattening (SB2E). Another Strathan Bay sample may indicate some flattening (SB6A), and three Ben Hutig samples possibly have remnants of early flattening but are highly overprinted by shear (BH10A, BH14SA, and BH14CA). The CPOs indicate that the magnitude of shear strain along this shear zone was, for the most part, able to overprint any flattening strains and that the lineation is parallel to the maximum stretch direction for these samples. None of the measured quartz c-axis fabrics contain well-defined cleft-girdles indicative of constriction, such as was observed by Xypolias et al. (2013) for the Aetopetra shear zone on Evia Island, Greece. The quartz crystallographic data primarily record the WNW-directed shear and apparent plane strain. The kinematic vorticity axes determined by CVA analysis, calculated from the rotation axes of individual grains, provide a more sensitive measure of the type of strain. For both areas, the CVAs indicate both progressive coaxial and noncoaxial shear, with more coaxial strain at Strathan Bay, and the complex patterns we observe suggest that they record more than just the final deformation.
In summary, our study documents components of coaxial flattening and noncoaxial shear that are manifested differently along and beneath the Ben Hope Thrust. Although the type of strain affecting the conglomerate pebbles is apparently the same overall, the relative magnitudes and resultant structures are distinctly different. Several factors lead to this partitioning. (1) The rheological difference between thin, alternating layers of conglomerate and psammite and more massive conglomerate clearly plays an important role. At Ben Hutig, conglomerate within the interlayered psammite would readily form buckle folds, which with increasing shear strain would rotate into the transport direction. (2) The location of the conglomerate relative to the thrust is another critical factor; the massive conglomerate at Strathan Bay is adjacent to the thrust itself, whereas the thinner conglomerate layers at Ben Hutig lie within the nappe package well beneath the thrust. (3) Temperature must also have played an important role. The rocks at Strathan Bay were at higher temperatures and were overthrust by even higher temperature nappes. In addition to noncoaxial shear, the rocks there responded to the mass of the overlying nappe with significant flattening accommodated by dislocation creep and GBM II. Later, F3/S3 deformation at lower temperatures was much less prevalent, as evidenced by minor SGR grain development affecting the dominant GBM II textures. These observations indicate that these rocks exposed at Strathan Bay most likely stayed hotter for a longer period of time. Traced toward the west, approaching the Moine Thrust, temperatures were lower, and at Ben Hutig, deformation was accommodated by the folding of interlayered metasedimentary rocks, the rotation of fold axes into the transport direction, and dislocation creep and GBM I recrystallization of quartz. As temperatures decreased at Ben Hutig, continued noncoaxial shear led to the formation of tubular sheath folds, and their dismemberment led to the formation of rods and SGR-dominated dynamic recrystallization. Coaxial flattening deformation is observed on Ben Hutig but was significantly less intense than at Strathan Bay.
Temperature appears to be the main control on the amount of coaxial deformation. It is possible, however, that strain rate or water content varied between the two locations. Strain rate would most likely be higher along the thrust, but GBM II rather than SGR microstructures are dominant, which is the opposite of what one would expect with increased strain rate. Water contents within the pebbles and conglomerate itself should have been similar. An influx of aqueous fluids along the thrust (Holdsworth, 1989; Holdsworth and Grant, 1990) is possible; however, the high temperatures documented here during shearing most likely preclude substantial fluid influx, but the effect cannot be ruled out. We propose that the Ben Hope Thrust placed much hotter, thick nappes on top of the Strathan Conglomerate, maintaining higher temperatures for a longer period of time, and allowing the mass of the overlying nappes to cause significant flattening in the underlying Moine Nappe. This is similar to the model for dynamic spreading within the Ben Hope Thrust plane proposed by Holdsworth and Grant (1990).
Model for Structural Evolution
Using the field, microstructural, and crystallographic data, we propose the following model to explain strain partitioning within the Strathan Conglomerate and this part of the Moine Nappe. The thick, massive conglomerate adjacent to the overlying Ben Hope Thrust at Strathan Bay underwent flattening due to the overburden of the overriding Ben Hope Thrust sheet and noncoaxial shear throughout the deformation process, aided by high temperatures (550–650 °C) during amphibolite-facies metamorphism (Fig. 17A). The L2 lineation parallels the WNW transport direction of the Ben Hope Thrust as substantiated by the CPO data. Dislocation creep and lattice reorientation on prism <a> and rhomb <a> planes, concurrent with or followed by easy GBM II, accommodated S2 shearing deformation and flattening of pebbles into highly oblate sheets. Both flattening and noncoaxial shear are supported by the CPO data. As temperatures decreased somewhat, SGR partially overprinted the earlier microstructures and minor folding occurred, all of which were associated with top-to-the-WNW shear sense.
Deformation of thinner interlayered conglomerate and psammite within the nappe package below the Ben Hope Thrust at Ben Hutig is dominated by multiple generations of folds. During emplacement of the overlying Ben Hope Thrust sheet, F2 folds initiated at a high angle to the transport direction as a result of noncoaxial shear and rheological differences between the interlayered psammite and conglomerate (Fig. 17B). Pebbles were flattened into quasi-sheets, defining S2 axial planar to F2 folds. As thrusting continued, progressive noncoaxial shear caused partial to complete rotation of F2 axes into parallelism with the transport direction as indicated by the variation in F2 axis orientations and their inferred isoclinal shape. Quasi-sheet–like pebbles underwent extension and local boudinage and may have become more elongate parallel to the transport direction during this rotation. Internal deformation was again accommodated by dislocation creep and lattice reorientation on prism <a> and rhomb <a> quartz slip systems concurrent with or followed by easy GBM I under lower temperatures (500–550 °C) than at Strathan Bay. CPO data support noncoaxial, top-to-the-WNW shearing, and vestiges of earlier coaxial deformation are observed.
During continued noncoaxial shear and decreasing temperatures (535–490 °C), sheath folds and F3 folds with axes parallel to the transport direction formed, as indicated by the open-to-closed, F3 interlimb angles, which would have approached an isoclinal shape (similar to F2) if they had undergone extreme rotation (Fig. 17C). These tubular folds would form because of the oblique orientation of the pre-existing S2/S0 anisotropy to the shear plane (Figs. 5B, 17B, and 17C). Local perturbations caused by deformed and boudinaged pebbles of different rheologies would cause the formation of smallscale F3 and sheath folds. Formation of F3/S3 was accompanied by SGR forming elongate, recrystallized quartz grains oriented parallel to the axial planes, which indicates a decrease in temperature. The inclined to recumbent parasitic folds with SGR grains parallel to the axial planes indicate some degree of flattening at this time. Progressive folding and apparent rotation and corkscrewing of elongate, recrystallized grains is observed in thin section. The presence of discontinuous and continuous undulatory extinction and many subgrains indicate that deformation continued as temperatures decreased. An alternative would be if strain rate increased during ongoing deformation (Law, 2014).
Rods formed in areas where the S2 fabric, consisting of quasi-sheets, was pulled up into F3 sheath folds to form parasitic folds as indicated by the close spatial relationship of rods and F3 (Fig. 18). The hinges of these parasitic folds thickened, and the limbs attenuated along small-scale shears seen in association with the rods. These parasitic-fold F3 axes rotated and lengthened as they were pulled along the sheath fold as shown by the reoriented and variable axial planar fabrics and by the wrapping of the lineation around the F3 axes and rods. The attenuated limbs evolved into tails and eventually disappeared, resulting in a highly elongate rod oriented parallel to the transport direction with a distorted internal fabric. Coeval shear along the length of the F3 axes and rods is indicated by the oblique, rotationally recrystallized grains indicating top-to-the-WNW shearing and local sheath folding.
Field, microstructural, and crystal fabric analyses of the Strathan Conglomerate in northernmost Scotland document the partitioning of strain intensity and how deformation is manifested between rocks adjacent to the Ben Hope Thrust and beneath it within the Moine Nappe.
No evidence for constriction is observed, and the strain ratios measured at Ben Hutig by Wood (1973) are only apparent strains. The extremely long rods are dismembered tubular fold axes that formed in that orientation on lithologic and structural an isotropies oriented at angles to the shear plane, in this case preexisting folds that had rotated into the shear direction. Their extreme length is not due to elongation, and the two short dimensions do not reflect shortening in those directions. Thus, when studying other ductile shear zones, careful analysis is needed to determine the type of strain and deformation path.
This study documents that the northern portion of the Ben Hope ductile thrust zone at Strathan Bay underwent extensive flattening, requiring extension in all dimensions within the shear plane and significant thinning, up to 89% to 92%, as proposed by Mendum (1976). The magnitude of this flattening strain and thinning needs to be considered in any tectonic model for the ductile nappe bounding thrusts in northern Scotland or other mid- to deep-crustal shear zones.
Strain at both conglomerate localities records noncoaxial shear. The observed strain partitioning is between intense flattening at Strathan Bay versus folding, rotation of fold axes, and tubular sheath folding accompanied by minor flattening at Ben Hutig. We observed that CPO fabrics associated with extensive flattening strains can be nearly or completely overprinted by significant shear strains. In complexly deformed shear zones, CVA, which plots the poles along which each individual crystal lattice rotates, can detect a combination of progressive coaxial and noncoaxial deformation not readily observed in CPOs. This study also indicates that CVAs can record progressive deformation and retain evidence of earlier rotations.
The combination of noncoaxial and coaxial shear documented here illustrates the importance of substantiating how much strain a ductile shear zone has undergone in the third dimension (i.e., along orogenic strike).
Portions of this paper are adapted from Sarah Collier’s Master of Science thesis; she thanks Miriam Barquero-Molina for help with fieldwork, Ben Davis for the illustrations in Figures 17 and 18, and William D. Carlson and Ian Dalziel for review of the thesis. We thank Rick Law, Robert Holdsworth, and an anonymous reviewer for substantial reviews that greatly improved the paper. Financial support for this research was provided by the Jackson School of Geosciences, the Geology Foundation of the University of Texas at Austin, and the Geological Society of America.