Progressive Development of a Distributed Ductile Shear Zone beneath the Patagonian Retroarc Fold-Thrust Belt, Chile Open Access

The Andes form the second highest mountain belt on the planet, exemplifying orogenic processes at an ocean-continent convergent margin. The magnitude and structural style of Cenozoic crustal shortening and uplift, however, are spatially and temporally variable along the strike of the orogen [1–4] and related to preexisting or external geologic conditions such as the sediment thickness in Mesozoic basins ([5]; cf. [6, 7]), lithospheric thickness [8, 9], and/or the dip of the subducting plate [10–12]. Although upper crustal fold-thrust belts (FTBs) accommodate substantial horizontal shortening [13], they can also be spatially separated from the regions of the highest topography [14], raising the question of how middle-to-lower crustal ductile thickening may be mechanically and kinematically coupled with brittle deformation of the upper crust to generate significant topography (e.g., [9, 15–17]). In this paper, we document the progressive evolution of distributed ductile shear and crustal thickening that formed in the middle crust beneath the Patagonian-Fuegian retroarc FTB.

The Patagonian Andes record a period of Late Cretaceous– Paleogene crustal contraction that resulted in the closure and inversion of a marginal basin, the Rocas Verdes Basin (RVB), and subsequent formation of the Magallanes foreland basin and Patagonian retroarc FTB [18–27]. A belt of metamorphic rocks, the Cordillera Darwin Metamorphic Core Complex (CDMC; [26, 28, 29]), records deformation and moderate to high pressure metamorphism that occurred during the Late Cretaceous–Paleogene orogeny [25, 26, 30, 31]. The CDMC is thought to represent either a pre-Jurassic accretionary complex [26, 32] or a succession of passive margin deposits [28] before it was metamorphosed during the Andean Orogeny. Rocks of the CDMC are tectonically juxtaposed against volcanic and sedimentary rocks of the RVB and Magallanes foreland basin within the Patagonian-Fuegian retroarc FTB [25, 31, 33, 34]. This juxtaposition provides a unique opportunity to study possible structural relationships between polyphase ductile deformation that occurs in the metamorphic basement and the development of an overlying retroarc FTB in the upper crust.

The study areas, located along the western margin of the CDMC (Figure 1), preserve both unconformable and sheared contacts between the CDMC and overlying volcanic and volcanoclastic strata of the RVB. At the northern locale, the contact is sheared and rocks of CDMC are imbricated with the rocks of RVB to form the hinterland domain of the Patagonian FTB. Retrodeformed line-balanced cross sections of the Patagonian FTB predict that the sheared contact is part of a regional décollement that detaches the Patagonian FTB from the underlying schist of the CDMC [35, 36]. Previous studies have identified similar shear zones that accommodated Late Cretaceous craton-vergent thrusting of Rocas Verdes rocks onto the continental margin elsewhere in southern Patagonia, indicating that the décollement may be a regional structure [18, 25, 26, 37]. Herein, we test the hypothesis that the regional décollement below the Patagonian retroarc FTB is rooted in a distributed ductile shear zone within the polydeformed basement of the CDMC. Our results exemplify how distributed ductile subdécollement thickening of the middle crust can be kinematically linked with the development of an overlying retroarc FTB and document an example of the down-dip end of a major décollement horizon transitioning to a ductile shear zone below a retroarc FTB in an Andean-style setting.

The CDMC is a topographic high that defines the apex of the Fuegian Andes. It is a ~5000 km² metamorphic massif of pelite, psammite, and orthogneiss of lower greenschist to upper amphibolite grade [25, 26, 28, 29, 34, 38]. Detrital zircon populations from metasedimentary rocks of the CDMC yield both Ordovician-Devonian and Carboniferous-Permian maximum depositional ages [28, 39] and are interpreted to record Paleozoic sedimentation along the continental margin of Gondwana [28]. Paleozoic sedimentary rocks of the CDMC are entirely overprinted by Cretaceous–Paleocene metamorphism and deformation that defines the Andean Orogeny in the southernmost Andes [25, 26, 30, 31, 34, 40].

Near the Beagle Channel (Figure 1), upper amphibolite-grade mineral assemblages that include sillimanite, kyanite, staurolite, and garnet record Late Cretaceous peak metamorphic conditions of $P=12$ kbar and $T=620$°C [30, 31]. The margins of the CDMC typically consist of chlorite-, biotite-, and garnet-bearing lower-greenschist facies pelitic and psammitic rocks, thus forming an antiformal dome with the highest metamorphic grade exposed in the core of the orogen [25, 26, 30]. A suite of Late Jurassic granitic orthogneiss known as the Darwin Granite intrudes pelitic and psammitic rocks of the CDMC and displays equivalent amphibolite-grade metamorphic mineral assemblages as the host rock.

At least three generations of Late Cretaceous structures that record crustal shortening and thickening during the Andean Orogeny pervasively affect orthogneiss and schist of the CDMC [25, 26]. Late Cretaceous–Paleogene deformation is characterized by bivergent folding and thrusting of high-grade rocks in the core of the orogen that thickened and uplifted the hinterland of the orogen, accommodating ~70% horizontal shortening ([25]; see also [26, 32]). Pseudosection modeling and in situ monazite geochronology of retrogressed mineral assemblages from second-generation fabrics within the high-grade core of the CDMC show that exhumation was underway by ~73 Ma [31] and is consistent with $P‐T‐t$ pathways calculated for sillimanite-, kyanite-, staurolite-, and garnet-bearing mineral assemblages from the CDMC presented by Kohn et al. [34]. Uplift and exhumation of the CDMC continued into the Paleogene partly accomplished by uplift in the hanging walls of thick-skinned, basement involved reverse faults [34, 39, 41].

An Upper Jurassic–Lower Cretaceous succession of volcanic, volcanoclastic, and marine sedimentary rocks unconformably overlies and surrounds the CDMC (Figure 1); this contact is sheared in most locations. Silicic volcanic and volcanoclastic rocks known as the Tobífera Formation record widespread silicic volcanism associated with Late Jurassic continental rifting and breakup of Gondwana [32, 42–47]. A coeval suite of bimodal, calc-alkaline intrusive rocks that form part of the Patagonian Batholith records the earliest pulses of arc magmatism along the Pacific margin of Gondwana in the southernmost Andes [48–51]. Jurassic extension resulted in the formation of the RVB marginal basin located between the Patagonian arc and the extended continental margin [52, 53]. Backarc extension in the RVB led to the generation of new quasioceanic crust that now comprises the Sarmiento and Tortuga complexes of southern Chile (Figure 1, [19, 49, 54]) and the Larsen Harbour Complex of South Georgia [55]. A thick succession (>2 km) of Upper Jurassic–Lower Cretaceous marine hemipelagic mudstones and turbidites was deposited in the RVB [52, 56–62].

Crustal thickening recorded by structures and prograde metamorphic textures in the CDMC (e.g., [25, 26, 30, 31, 34, 63]) was contemporaneous with the development of the Magallanes Foreland basin and Patagonian FTB (Figure 1, e.g., [40, 56–58, 64–66]). Imbricated, northeast-vergent first-generation ductile thrusts near the Beagle Channel placed the basaltic floor of the RVB onto the continental margin. Upper Cretaceous posttectonic granitic plutonic rocks and pegmatite dikes (~86 Ma, [25, 50]; see also [55]) intrude and cut these early thrusts, indicating that obduction of the RVB floor occurred by the Late Cretaceous. Continued shortening resulted in the development of a thin-skinned thrust belt that propagated into the Magallanes Foreland basin by the Late Cretaceous [24, 33, 36, 40, 67–71]. A second pulse of deformation characterized by thick-skinned faults dissected the FTB and uplifted basement rocks of the CDMC and Rocas Verdes terrane, aiding the exhumation, denudation, and resedimentation of these rocks into the Magallanes foreland basin by the Paleogene [25, 33, 36, 41, 57, 63, 65, 66, 68, 71].

To document the tectonic evolution of the structurally lowest part of the Patagonian–Fuegian FTB, we mapped (1 : 50,000 scale) the contact between polydeformed rocks of the CDMC and the structurally overlying FTB in two locations that encompass the basement-cover contact: (1) within the high-grade core of the orogen near Seno Martínez and (2) toward the foreland on Peninsula Brunswick where the CDMC is juxtaposed above the Patagonian–Fuegian FTB (Figures 1–3). Outcrop-scale (10 m²) structural observations were made to interpret the structural significance of polyphase folding and high-strain zones within with the CDMC and their kinematic relationship with the structurally overlying FTB. Field observations were combined with microstructural and CPO analyses of thin sections from the CDMC to interpret the kinematic development and deformation conditions of the shear zones.

Empirical and experimental studies have provided information on the relationship between common quartz CPOs in shear zones and active slip systems, strain type (plane strain, flattening, or constriction), the degree of noncoaxial deformation (kinematic vorticity), and the activity of recrystallization mechanisms [72–83]. Quartz CPOs have been examined in the context of crystal slip systems, recrystallization mechanisms, and dislocation creep regimes over a large range of deformation temperatures (~270 to ~700°C; see [77] for a review). Two mechanisms for estimating temperatures of deformation are widely used, and we use a combination of the two in this paper.

The quartz recrystallization microstructure thermometer was proposed by Stipp et al. [83, 84] based on a study of the Tonale shear zone and can be related to quartz recrystallization “regimes” identified by experimental work [85]. In this method, temperatures are estimated using dominant recrystallization processes: grain boundary bulging (regime 1), subgrain rotation (regime 2), and grain boundary migration recrystallization. Two types are observed for the latter, the transition from subgrain rotation to grain boundary migration (regime 3) and high-temperature grain boundary migration.

The fabric opening-angle thermometer uses cross-girdle patterns in quartz $c$-axis pole figures [86–88] for estimating deformation temperatures (with ±50°C precision). Quartz $c$-axis opening angles increase with deformation temperature as prism [c] slip becomes increasingly important over basal <a> slip ([89]; review in [77]).

Both of these thermometers assume that temperature is the only controlling mechanism, though Faleiros et al. [86] present a multiple regression opening-angle thermometer calibration that is also sensitive to pressure. However, it is widely recognized from both experimental (e.g., [85]) and natural (e.g., [90, 91]) studies that both quartz $c$-axis opening angles and quartz slip-systems inferred from CPO patterns and recrystallization microstructures are also sensitive to other factors, including hydrolytic weakening (e.g., [92]: [93]), strain rate (e.g., [84, 85, 94, 95]), and type of 3D strain [72, 80, 82]. Thus, there may be significant variation in the deformation temperatures at which a given range of opening angles or quartz slip systems occurs between differing shear zones (e.g., [90, 91]; review in [77, 86]). We use petrologic estimates of deformation temperatures to calibrate our results.

Mineralogical and quartz textural observations, as well as CPO analyses, were performed on thin sections cut perpendicular to foliation and parallel to the most prominent quartz stretching lineation for analyses of S/L fabrics from dynamically recrystallized quartz-rich shear zones (i.e., $XZ$ thin sections). In micaceous domains, thin sections were cut perpendicular to the kink fold axes. Because the deformation was generally noncoaxial, the $XZ$ sections may not represent the plane normal to maximum vorticity for each sample. For consistency, we conducted analyses using the $XZ$ sections to compare fabrics that are parallel to the bulk top-northeast transport direction as determined from field observations (discussed in Sections 4 and 5).

Quartz CPOs were measured using the HKL Nordlys Electron Backscatter Detector (EBSD) on the Phillips XL30 Environmental Scanning Electron Microscope at the University of Texas at Austin Department of Geological Sciences. All analyses were conducted using a working distance between 18 and 25 mm and an accelerating voltage of 25–30 kV. Background conditions and diffraction patterns were automatically collected using the AztecHKL EBSD software package by Oxford Instruments. Only indexed diffraction patterns with a mean angular $deviation<1°$ were accepted. Data were collected in quartz-rich domains using a variable step size that was appropriate for the grain size of each sample. The data were reduced using the HKL Channel 5 software to one point per grain with a misorientation angle of 10° to define grain boundaries. $c$-axis and $a$-axis pole figures were plotted using Pfch5 [96] and contoured with multiples of uniform distribution (MUD). Only plots with $c$-axis pole figures that yield $MUD>2$ were considered.

Seno Martínez is a north-northeast-trending fiord located along the western margin of the Cordillera Darwin in Tierra del Fuego (Figures 1 and 2). Bedrock exposures of the CDMC crop out along the shoreline as a chlorite, biotite, and locally garnet-bearing pelitic and psammitic schists of lower-greenschist to lower-amphibolite facies (Figure 2). The southern end of the fiord preserves plutonic rocks of the Patagonian Batholith as well as epidote-actinolite-bearing schist that have been interpreted as part of the RVB seafloor [18, 25, 26]. The contact between the batholith and the RVB terrane is not exposed and is inferred to be an intrusive contact. Along the eastern shore of Seno Fontane (Figure 2), epidote-actinolite schist of the RVB terrane is in tectonic contact with the underlying pelitic and psammitic schists of the CDMC. Toward the north of the study area, the mouth of Seno Martínez merges with a northwest-trending fiord called Seno Keats. On the northern shore of Seno Keats at Bahía Angelito (Figure 2), quartz- and chlorite-bearing schist of the CDMC is in tectonic contact with the overlying Jurassic volcanic strata of the Tobífera Formation. Thus, pelitic and psammitic schists of the CDMC exposed along Seno Martínez form a structural culmination bound on the north by Jurassic rocks of the Tobífera Formation and the south by mafic schist of the RVB terrane (Figure 2). Pelitic and psammitic schists of the CDMC, exposed along Seno Martínez in the core of the culmination, record at least four generations of ductile noncylindrical folds that are attributed to four deformation phases (D₁–D₄, discussed below, Figure 4) and form the Magallanes shear zone (MSZ, defined herein).

Along strike ~110 km to the northwest from Seno Martínez, rocks of the CDMC are exposed on Peninsula Brunswick where they are tectonically interleaved with Jurassic rocks of the RVB (Tobífera Formation). Near Bahía Fortesque (Figures 1 and 3), a ~1 km thick ductile high-strain zone known as the Bahía Fortesque shear zone (BFSZ) juxtaposes the CDMC above the Tobífera Formation [36]. Near Estuarios Silva Palma (>10 km north of Bahía Fortesque; Figure 3), the CDMC is exposed in the hanging wall of an out-of-sequence fault that places it above rocks of the Patagonian FTB and Upper Cretaceous to Paleocene foreland basin strata. In both locations, chlorite-quartz schist of the CDMC displays several generations of deformation ([36]; discussed in Sections 4.7 and 4.8) that are similar to those at Seno Martínez (described in the next section).

The two combined study areas are divided into six structural domains on the basis of overprinting relationships and intensity of specific structures (domains I–IV, Figure 2; domains V–VI, Figure 3). In the following sections, we describe the macro- and microscopic structures preserved within each domain and then correlate them throughout the study area (Figure 4).

The southeastern end of Seno Fontane (site 1005, Figure 2) exposes outcrops of metavolcanic rocks that are predominantly epidote-actinolite-grade greenschist facies schist (Figure 5). The schist mineral assemblage is epidote, actinolite, (±) hornblende, feldspar, quartz, pyrite, and hematite (Figures 5(c) and 5(d)). Diabase dikes intrude the schist and are commonly boudinaged or display pinch and swell structures (Figure 5(a)). In some locations, metatuffs are preserved that dip gently toward the west-southwest (Figure 5(b)). Here, a subsolidus schistose foliation (S₁) is defined by the alignment of actinolite (±) hornblende and recrystallized quartz (±) feldspar layers that also dip gently toward the west-southwest. A poorly developed recrystallized quartz lineation (L₁) plunges gently toward the west-northwest (Figure 2). On the basis of their probable volcanic origin and intermediate-mafic composition mineral assemblage, we interpret the epidote-actinolite schist as reflecting part of the oceanic floor of the Jurassic RVB (cf. [26]).

Northward for about 5 km, the epidote-actinolite schist is tectonically interleaved with pelitic chlorite- and muscovite-bearing schists, interpreted as part of the CDMC. The pelitic schists contain a pervasive composite schistose foliation (S₂) defined by the alignment of muscovite, quartz, and chlorite. Compositional layering defined by micaceous, fine-grained quartz and feldspar domains and coarse-grained quartz-rich domains (S₁) forms intrafolial, rootless, and isoclinal folds (F₂) that are transposed parallel to and partly define the composite macroscopic foliation. S₁ in the epidote-actinolite schists of the RVB terrane (sites 1005, 1007, and 1009) is subparallel to the composite S₂ fabric of the CDMC and considered equivalent (Figure 4). Therefore, S₁ in the CDMC schist must reflect an older deformation (D₁; cf. “pre-Andean” fabric of [26]) than the prominent composite foliation that is shared by the epidote-actinolite and pelitic schist (D₂); D₂ fabrics are interpreted to record the earliest Andean deformation. Intrafolial folds (F₂) are unique to the pelitic schist and not expressed in the epidote-actinolite schist. Where the pelitic and epidote-actinolite schists are interleaved, they are sheared and preserve a schistose foliation (S₂) that dips moderately southwest and contains a prominent dynamically recrystallized quartz stretching lineation that plunges southwest (S₂/L₂, Figures 6(a), 6(c), and 6(e)). Diabase dikes that intrude the pelitic and epidote-actinolite schists are dismembered, forming boudin lenses enveloped by S₂ (Figure 6(b)); boudin neck extension directions do not appear to have a preferred orientation among the few outcrops where they were observed (Figures 6(b) and 6(d)). Epidote-actinolite schist is cut out by a thrust fault north of site 1009 (Figure 2).

Pelitic schist of the CDMC in the northern part of domain I (sites 1009–1012) preserves two additional phases of folding also not expressed in the epidote-actinolite schist of the RVB terrane (Figure 4). Sets of tight, upright, inclined, and reclined folds (F₃) refold the pervasive S₂ schistose foliation (Figures 6(e) and 6(g)). F₃ folds generally do not exhibit a well-developed axial planar foliation. F₃ fold axes plunge shallowly either northwest-southeast or east-west (Figure 6(g)). In some locations, F₃ fold axes are subparallel with the L₂ quartz stretching lineation (Figure 6(e)). Asymmetric, upright kink bands (F₄) fold the S₂ surface (Figure 6(f)) and overprint F₃. The axial surfaces of F₄ kink bands are well developed in phyllosilicate-rich domains and dominantly dip steeply toward the west-southwest (Figure 6(h)). A subsidiary set of kink bands dips steeply toward the east-northeast, conjugate to the west-southwest dipping set. Kink band axes plunge shallowly toward the north and south (Figure 6(h)). F₄ kink folds are asymmetric and show senses of rotation that are compatible with horizontal shortening (i.e., top-northeast on southwest-dipping S₄ surfaces or top-southwest on northeast-dipping S₄ surfaces).

Rocks exposed along Seno Oryan are located in the core of a regional antiform (F₄, Figure 2) and display the highest metamorphic grade in the study area. Garnet-bearing upper greenschist facies pelitic schists exposed in the antiform core record the superposition of three generations of folds (F₂–F₄, Figures 2, 4, and 7, sites 1009–1018). Here, the pervasive macroscopic foliation (S₂) is defined by the same composite foliation as in domain I and contains a prominent quartz stretching lineation (L₂, Figures 7(e) and 7(f)). From Seno Oryan to the northern end of domain II (site 1021; Figure 2), S₂ and L₂ are folded by tight and isoclinal folds (F₃, Figures 7(a)–7(d)); L₂ and poles to S₂ define a southwest-northeast-trending girdle (Figure 7(g)). Here, F₃ folds occur as upright, reclined, and recumbent folds depending on how they are refolded by F₄ (Figures 7(a)–7(d)). F₃ fold axes plunge moderately toward the west-northwest and southeast or steeply toward the east. Axial surfaces to F₃ folds (S₃) dip steeply and moderately toward the north and east, indicating both upright-horizontal and reclined orientations for F₃ folds (Figure 7(h)), as they are folded by F₄. Quartz lineations (L₂) plunge shallowly toward the west-northwest and are locally parallel to F₃ axes (Figures 7(g) and 7(h)). In high shear strain zones, F₃ form sheath folds (Figures 7(c) and 7(d)). In outcrops where the limbs of F₃ sheath folds are well preserved, a prominent quartz mineral lineation exists on the folded S₂ surface (Figure 7(c)). Here, sheath folds (F₃) refold the S₂/L₂ surface, have isoclinal limbs, and form long axes that are subparallel to the local stretching lineation (L₂, Figure 7(c)) and plot along the same girdle as S₂ and L₂ (Figure 7(g)). Everywhere in domain II, the S₂/L₂ surface and F₃ folds are refolded by upright, steeply plunging tight folds (F₄) that form both Ramsay type II (mushroom) (Figure 7(b)) and type III (fishhook) (Figure 7(a)) superposed fold patterns [97]. F₄ folds have near-vertical, north-striking axial surfaces and steeply north- and south-plunging fold axes and thus form reclined folds (Figures 7(a), 7(b), and 7(i)).

In domain III (Figure 2), quartz-, chlorite-, and muscovite-bearing schists contain the same pervasive composite schistose foliation (S₂) and recrystallized quartz mineral lineation (L₂, Figure 4). Here, regional, gently inclined isoclinal folds (F₃) refold the S₂/L₂ surface and have amplitudes on the order of 100 s of meters (e.g., Figure 8(a)). Poles to refolded S₂/L₂ surfaces form a southwest-northeast-trending girdle (Figure 8(c)). S₃ axial surfaces dip gently toward the north and south, and F₃ axes plunge shallowly toward the west and southeast (Figure 8(d)). Everywhere in domain III, the S₂/L₂ surface and F₃ folds are overprinted by sets of kink folds (F₄) that have steeply east-northeast- or northeast-dipping axial surfaces (S₄) and fold axes that plunge both steeply and shallowly toward the south-southeast (Figures 8(a), 8(b), and 8(e)). F₄ kink folds are best expressed in micaceous layers where they display top-southwest sense of rotation and a well-developed northeast-dipping axial surface (S₄) that crosscuts S₂ (Figure 8(b)).

Within the CDMC pelitic schist, micaceous domains are interlayered with dynamically recrystallized quartz domains that define the composite S₂ foliation. In micaceous domains, kinematic indicators include C$′$ shear bands and asymmetric muscovite sigma tails on garnet porphyroblasts (Figures 9(a) and 9(b)). In thin sections cut perpendicular to S₂ and parallel to L₂, both types of kinematic indicators show top-northeast sense of shear. Curved, synkinematic quartz inclusion trails are present in garnet porphyroblast cores and are consistent with dextral (top-NE) rotation. Faint inclusion trails in some garnet rims are continuous with S₂ and also indicate top-northeast shear (⁠$XZ$ plane, Figures 9(c) and 9(d)). In quartz-rich lithologies of the CDMC, alternating layers of dynamically recrystallized fine- and coarse-grained quartz define S₂ (Figures 10(a)–10(c)). In coarse-grained layers, recrystallized quartz grains are typically elongate parallel to S₂ and have seriate, interlobate, to ameboid grain boundaries (“SIA” in Figure 10). Quartz grains are commonly internally deformed, containing subgrain boundaries, and new recrystallized polygonal grains have similar shapes and sizes as the new grains (e.g., Figures 10(a)–10(c)). Fine-grained layers consist of polyphase aggregates of quartz, feldspar, and chlorite. Although quartz is also recrystallized in fine-grained layers, pinning microstructures associated with the other phases appear to have inhibited grain growth (e.g., fine-grained, polyphase domains in Figures 10(a) and 10(b)). Quartz textures from S₂ fabrics indicate both subgrain rotation and grain boundary migration recrystallization consistent with regime 3 dislocation creep [85] and the subgrain rotation recrystallization to grain boundary migration transition of Stipp et al. [83]. In samples collected near the core of the F₄ antiform at Seno Oryan (Figure 2), some quartz grains also exhibit chessboard extinction indicative of high-temperature dislocation creep (Figure 10(c)). Quartz textures do not exhibit evidence for a separate foliation associated with F₃ folds, consistent with the observation that macroscopic F₃ folds are not associated with a pervasive axial planar cleavage (i.e., Section 4.2).

In rocks where F₄ kink bands are well expressed (e.g., domain III), the composite schistose foliation (S₂) is pervasively overprinted by F₄ kink folds (Figure 10(e)), and quartz exhibits a prominent shape-preferred orientation that is oblique to S₂ (~65°) and parallel to the axial surface of F₄ (S₄, Figures 10(d)–10(f)). Recrystallized and elongate quartz grains defining S₄ contain parallel well-developed subgrain boundaries, indicating subgrain rotation recrystallization (Figures 10(d)–10(f)). These quartz grain boundaries also are locally irregular and cuspate-shaped (Figures 10(d) and 10(f)), providing textural evidence of grain boundary migration recrystallization during S₄. Quartz textures from S₄ domains indicate both subgrain rotation and grain boundary migration recrystallization consistent with regime 3 dislocation creep [85] and/or the subgrain rotation to grain boundary migration transition of Stipp et al. [84]. Microstructural kinematic indicators parallel to S₄ are ambiguous; however, asymmetric kink folds that define the macroscopic S₄ foliation always show senses of rotation that are compatible with horizontal shortening (i.e., top-southwest on northeast dipping surfaces and vice versa, Figure 6(h)).

Along the north shore of Seno Keats in Bahía Angelito, the quartz- and chlorite-bearing pelitic schist of the CDMC is in tectonic contact with the Tobífera Formation (Figures 2 and 11). Here, a 4 m thick conglomerate with pebble- and cobble-sized clasts supported by a muddy matrix unconformably overlies the CDMC, while the conglomerates, in turn, are overlain by 1 m of turbidite sandstones (Figures 11(a)–11(c)). The conglomerates and sandstones dip southwest and are interpreted to represent a basal clastic unit within the Tobífera Formation that is described elsewhere in Tierra del Fuego (cf. [32, 98]).

The unconformable contact is sheared and cut by several small thrust faults. Five meters below the contact (site 1024, Figure 2), pelitic schist of the CDMC is observed in the footwall of a thrust where it preserves structures similar to those in domains I and III in Seno Martínez (Figures 4, 11(a), and 11(c)). Here, a pervasive schistose foliation (S₂) contains a recrystallized quartz stretching lineation (L₂), and the S₂/L₂ surface is folded (F₃), forming a north-trending girdle (Figure 11(d)). F₃ folds are isoclinal and upright with steeply north-northwest- and south-southeast-dipping axial surfaces that form a conjugate set. F₃ fold axes are horizontal and trend toward the east-northeast and west-southwest (Figure 11(e)). Directly below the contact (<5 m), the CDMC is sheared and displays schistose foliation (S₂) that dips dominantly toward the southwest (Figure 11(c)). Southwest-dipping C-S fabrics that are parallel to S₂ indicate top-northeast sense of shear. In some locations, lenses of conglomerate are enveloped by the basement schist and are bound on all sides by small thrust faults (Figures 11(a) and 11(c)). Both the conglomeratic lenses and sandstones contain a slaty cleavage defined by pressure solution seams that are coplanar with the foliation in the pelitic schist (Figures 11(a)–11(c) and 11(f)). The brittle thrust faults crosscut the ductile D₂–D₃ fabrics in the schist, record slip along the contact, and are grouped with D₄ structures. Quartz microstructures from the pelitic schist of the CDMC in domain IV are similar to those of domains I–III and are not discussed further. The shear zone within the pelitic schist of the CDMC below the contact is inferred to be coeval with D₂ and D₃ deformation in domains I–III.

One hundred meters to the southwest, the conglomerate and sandstone deposits are conformably overlain by volcanoclastic units of the Tobífera Formation (Figure 11(g)). Here, white, silicic ignimbrites overlie dark-grey ash beds and dip southwest. The ash beds display a slaty cleavage defined by pressure solution seams and flattened quartz and feldspar lapilli. Bedding is not distinguishable within the ash beds. The slaty cleavage dips toward the southwest and is overprinted by a set of crenulation cleavages. The crenulation cleavage dips steeply toward the east-northeast, and crenulation lineations are horizontal, trending north-northwest and south-southeast (Figure 11(h)). The crenulation cleavage is asymmetric and shows a top-to-the-southwest sense of rotation, indicating contraction. The ash deposits are cut by a set of small, brittle thrust faults that form a southwest-dipping imbricate fan (Figure 11(g)). Fault planes dip shallowly toward the south-southwest; one set contains a slickenline that plunges obliquely toward the west-southwest, indicating a thrust-left sense of slip (Figure 11(i)). Sets of tensile quartz veins occur in close proximity to the thrusts and dip towards the east-northeast (Figures 11(g) and 11(j)). Some quartz veins are sigmoidal and indicate a top-northeast sense of rotation. Although inaccessible, bedding in the Tobífera Formation is well exposed on the mountainside northeast of Bahía Angelito where it dips toward the northeast and is gently folded to form a regional syncline, suggesting that the sheared contact at the base of the Tobífera Formation is imbricated and folded (Figure 2).

At Bahía Fortesque (Figure 1), the CDMC is imbricated with the Tobífera Formation along a ~1 km thick contractional shear zone known as the Bahía Fortesque shear zone (BFSZ, Figure 3; [36]). The shear zone is defined by a southwest-dipping, composite tectonic foliation (S_1-2) that contains a prominent southwest-plunging, recrystallized quartz lineation (L₂, Figure 4). Macroscopic kinematic indicators include C$′$ shear bands and asymmetric, isoclinal, and intrafolial folds (F₂ and F₃) that all indicate top-northeast shear (see additional description in [36]). In thin section, the S₁ foliation is defined by compositional layering between medium-grained quartz-rich domains and fine-grained quartz- and phyllosilicate-rich domains (Figures 12(a) and 12(b)). S₁ is commonly folded by intrafolial isoclinal folds (F₂) that are transposed parallel to and help to define the composite foliation (S_1-2, Figure 12(a)). The composite foliation (S_1-2) is folded by isoclinal recumbent folds (F₃; Figure 12(b)) with coplanar axial surfaces. A sample collected from below the BFSZ (11016) preserves upright, tight F₃ folds that do not have a pervasive S₃ (Figure 12(d)), indicating qualitatively lower strain than samples from within the BFSZ (cf. Figures 12(a)–12(c)). The foliation numbering scheme presented here is modified from that of Betka et al. [36] by $Sn+1$ because the early composite intrafolial folds (F₂) of S₁ compositional layering were not recognized in that study of outcrop-scale structural fabrics.

Quartz occurs in layers parallel to S_1-2 in the BFSZ and has undergone substantial recrystallization that resulted in the formation of new grains (Figures 12(a)–12(c)). The new grains are commonly polygonal and have approximately uniform grain sizes and have similar shapes and sizes to subgrains (Figures 12(a)–12(d)). Some grain boundaries are irregular and cuspate-shaped (Figure 12(a)). Quartz microstructures from the BFSZ are dominantly characteristic of subgrain rotation recrystallization and regime 2 dislocation creep of Hirth and Tullis [85] and/or the subgrain rotation regime of Stipp et al. [83]. In some locations, layers of recrystallized, elongate grains (S₄) oblique to S₂ indicate top-northeast shear (Figure 12(a)). These elongate grains are subparallel to the axial planes of steeply dipping kink folds (F₄) that are superimposed on S₂ fabrics [36]. Feldspar grains are not common, but where observed, they are not substantially recrystallized (Figure 12(c)).

The CDMC crops out in the hanging wall of a basement-involved reverse fault at Estuario Silva Palma, ~10 km north of Bahía Fortesque (Figure 3; described in [36]). Here, outcrops of the CDMC are structurally below the BFSZ and are juxtaposed above imbricated strata of the Magallanes fold-thrust belt along a south-dipping reverse fault (Figure 3; [36]). The schist is deformed by two generations of structures similar to those at Bahía Fortesque (Figure 4). A pervasive schistose foliation (S_1-2) contains a prominent quartz stretching lineation (L₂), and both structures are folded by tight F₃ folds (see Figure 13 in Betka et al. [36]). S_1-2 is also defined by compositional layering between medium-grained quartz-rich domains and fine-grained quartz and phyllosilicate-rich domains. In contrast to Bahía Fortesque, here, the F₃ folds are tight but are not transposed to form a composite foliation suggesting qualitatively lower strain in Estuario Silva Palma. In quartz-rich domains, quartz occurs in layers that are extensively recrystallized to form new polygonal grains of approximately uniform grain size that is similar to the shape and size of the subgrains (Figure 12(d)). In the hinges of F₃ folds, recrystallized quartz grains are elongate parallel to the axial plane (S₃) of the folds (Figure 12(d)). Some grain boundaries are irregular and cuspate-shaped (Figure 12(d)). Quartz microstructures from the CDMC in Estuario Silva Palma also are indicative of subgrain rotation recrystallization and regime 2 dislocation creep of Hirth and Tullis [85] and/or the subgrain rotation regime of Stipp et al. [84]. Microstructural shear sense indicators are commonly absent or ambiguous in samples from Estuario Silva Palma.

Field and microstructural observations indicate that there are two prominent foliations (S₂ and S₄) in all of the structural domains (I–IV) within the MSZ near Seno Martinez and one prominent foliation (S_1-2 in domains V and VI) within the BFSZ on Peninsula Brunswick. To better characterize the deformation temperatures, type of strain, and sense of shear during the formation of each foliation, we measured the CPOs of dynamically recrystallized quartz in ten samples from S₂ domains in the MSZ near Seno Martínez (sample locations on Figure 2) and nine samples from the S_1-2 composite foliation of the BFSZ (sample locations on Figure 3). Quartz $c$-axis opening angles were measured on samples that yielded well-defined cross-girdle patterns. All analyses were conducted on $XZ$ thin sections. For S₂ fabrics, $X$ is parallel to L₂ and $Z$ is perpendicular to S₂. The quartz CPO results and interpretation of the slip system(s), sense of shear, deformation temperatures, and type of strain from each sample are presented in Figures 13, 14, and 15 and in Tables 1–3. The results are summarized below and discussed with the outcrop-scale structural observations in the following section (Section 6).

From domain I where rocks of the RVB terrane and the CDMC are tectonically interleaved (Figure 2), three samples were analyzed: 1006A from the CDMC and 1007A and 1009A from the RVB (Figure 13, Table 1). Here, the CDMC shows coaxial deformation with dominantly basal <a> slip (⁠$OA=59°$⁠). The RVB samples show top to the northeast sense of shear, consistent with outcrop shear sense indicators at this location, and both basal <a>, rhomb <a>, and prism <a> slip (where present, $OA=44°$⁠).

In the northern part of domain I in the CDMC, three samples (1010CA, 1011A, and 1012A) from garnet-bearing schist were analyzed and reflect a higher metamorphic grade (Figures 2 and 13, Table 1). Two samples show top to the west-northwest or northeast sense of shear with basal <a>, rhomb <a>, and prism<a> slip (OA 70–77°). The sample farthest north (1012A) shows noncoaxial flattening with top to the west-southwest sense of shear and evidence of probable increasing temperatures with both basal <a> and prism [c] (⁠$OA=132°$⁠).

In domain II, sample 1015A is from garnet-bearing schist in the core of the F4 antiform, and 1018C and 1018EA are from quartz-chlorite-muscovite schist on the northern flank (Figures 2 and 13, Table 1). Sample 1015A shows top to the northwest sense of shear and predominantly prism [c] slip (Figure 13(g)). Of the other two samples, one (1018C) shows predominantly prism [c] slip (⁠$OA=~138°$⁠) with top to the south sense of shear, and the other (1018EA) records predominantly basal <a> and lessor prism <a> and rhomb <a> slip with top to the northeast sense of shear.

One sample (1021BA) was collected from the northern end of domain II, where sheath folds are present (Figures 2 and 7). This sample shows top to the north sense of shear with both basal <a> and predominantly prism <a> slip (⁠$OA=40°$⁠), similar to the samples from the southern end of domain I (i.e., Figures 13(a)–13(e), Table 1) and compatible with cooler D₂ deformation temperatures than in the southern end of domain II. The presence of prism [c] slip and larger opening angles (OA) in samples 1015A and 1018C is consistent with the location of those samples in a structurally lower position in the core of the F₄ antiform. These locations include garnet-bearing schist with synkinematic inclusion trails (Figure 9) and textural evidence (i.e., chessboard extinction, Figure 10(c)) for high-temperature dislocation creep that occurs in Seno Oryan near the core of the F₄ antiform. S₂ CPOs in domain III are pervasively overprinted by S₄; no samples from domains III or IV displayed quality S₂ data.

At Bahía Fortesque (domain V), samples 11105A-D of CDMC are from within the high-strain zone (Figures 14(a)–14(d)) and sample 11016 (Figure 14(e)) is from below it. All samples show top to the northeast sense of shear with predominantly basal <a> slip along with variable amounts of prism <a> slip (⁠$OA=50$–70°), consistent with cooler deformation temperatures than fabrics from domains I to II (Table 2).

In domain VI near Estuario Silva Palma, samples 11053 and 11054A–C (Figure 15, Table 3) were collected from the CDMC below the BFSZ and show top to the northeast sense of shear or, in one sample coaxial deformation, with basal <a>, prism <a>, and variable rhomb <a> slip (⁠$OA=~58°$⁠, ~55°, and ~45°), consistent with the deformation conditions of domain V.

Structures in the CDMC near Seno Martínez and Peninsula Brunswick share many common elements that allow regional correlation on the basis of their stratigraphic position, sense of shear, and sequence of deformation. Most importantly, the BFSZ, the D₂–D₃ high-strain zones of the MSZ, and the Bahía Angelito fault zone each accommodated top-northeast thrusting of rocks from the RVB terrane (Tobífera Formation at Bahías Fortesque and Angelito, mafic-intermediate volcanic rocks of the RVB seafloor at Seno Martínez) over the CDMC schist (Figures 2 and 3). In each of these localities, the earliest macroscopic structure in the CDMC is a pervasive schistose foliation that contains a prominent down-dip quartz mineral lineation that is refolded (i.e., S₂/L₂ at Seno Martínez and S_1-2 composite and L₂ at Bahía Fortesque and Estuario Silva Palma) and displays dominantly top-northeast thrust shear-sense indicators. Similarly, D₃ is characterized by the modification of D₂ structures into northeast-vergent tight and/or isoclinal noncylindrical folds that are interpreted to reflect progressive deformation during top-northeast late D₂ and D₃ shearing. Qualitative strain gradients defined by the tightness of D₃ folds and/or occurrence of F₃ sheath folds indicate D₃ high-strain zones in domain II at Seno Martínez and at Bahía Fortesque (i.e., [36]). D₄ structures are characterized everywhere by upright or steeply inclined kink folds and/or tight folds, indicating a regional period of northeast-southwest horizontal shortening that is superimposed on D₂/D₃ structures (discussed below, Figure 4).

(1) D₂ High-Strain Zones. Quartz textures and CPOs from D₂ fabrics in domains I and II indicate an apparent increase in deformation temperature structurally downsection from the tectonized contact with the RVB terrane toward the core of the D₄ antiform (domain I, Figures 2 and 16) and a decrease northward through domains II–VI. In the southern part of domain I (DI, RVB in Figure 16), quartz mylonite preserves basal and prism <a> slip systems, $c$-axis opening angles between ~45 and 60°, and quartz textures show evidence of subgrain rotation and grain boundary migration recrystallization (regime 3, [85]), altogether consistent with deformation temperatures of ~300–500°C (Figure 16; cf. [83, 86]).

Near the core of the D₄ antiform (Seno Oryan, Figure 2), prism [c] slip is also present in D₂ fabrics within garnet-bearing schist, and the $c$-axis opening angles increase to 70–138°, indicating warmer deformation temperatures (~500 to >650°, DI+gt and south DII in Figure 16, cf. [83, 86]). The anomalously large opening angles (132° and 138°, samples 1012A and 1018C) and presence of prism [c] slip yield deformation temperatures of 865 and $893±50°C$⁠, respectively, using the opening angle thermometer [86]. These temperatures are incompatible with the garnet-greenschist to lower amphibolite mineral assemblage of the schist and peak metamorphic temperatures [30] (Figures 12, 13, and 16). Additionally, at such high temperatures, the quartz textures would be expected to reflect widespread high-temperature grain boundary migration recrystallization (large amoeboid grains, GBM II of [83]) and partial melt. We infer that the presence of water or slower strain rates, and potentially the increasing importance of flattening strain in the core of the MSZ, might have locally favored the transition to prism [c] slip, increasing the opening angle (e.g., [77, 86]). We thus report the temperature range of 500 to $>650±50$°C constrained by samples 1010CA and 1011A for D₂ fabrics within the garnet-bearing region of domain I and the southern part of domain II (Figure 16).

In the northern part of domain II, prism [c] slip is absent, and the $c$-axis opening angles decrease to ~40°, indicating deformation temperatures of $~325±50°C$ (north DII in Figure 16), equivalent to domain I and consistent with the observed D₄ doming of the garnet isograd between domain I and the northern part of domain II (Figure 16).

Although the observed quartz slip systems and microstructures across the entire area are sensitive to other variables (i.e., strain type, strain rate, and hydrolytic weakening), the estimated deformation temperatures are further supported by garnet-chlorite and garnet-biotite thermobarometry carried out by Kohn et al. [30] on two samples collected near Seno Oryan (Figure 2). Their results yielded “peak” pressure and temperature conditions of ~9–10 kbars and ~545–570°C, respectively, in good agreement with our results from garnet-bearing fabrics in domain I (Figure 16).

Microstructural kinematic indicators from the CDMC in domains I and II include C-S and C$′$ shear bands and asymmetric tails on synkinematic garnet porphyroblasts that are compatible with top-northeast shear during D₂ (Figure 9). Supporting these observations, pole figures from D₂ recrystallized quartz fabrics in domains I and II are commonly asymmetric and indicate either dominantly coaxial or top-northeast sense of shear (Figures 13(b), 13(c), 13(e), 13(i), and 13(j)). However, quartz CPOs from several samples collected in domains I and II have asymmetries that indicate a different sense of shear, either top-west or northwest (Figures 13(d), 13(f), and 13(g)) or top-south (Figure 13(h)). These samples probably reflect either reorientation of the D₂ fabrics as they were folded around later (F₃ and F₄) folds or local conjugate shearing under coaxial strain. The dominance of top-NE kinematic indicators from microstructures and quartz CPOs is interpreted to reflect bulk top-northeast shear during D₂ in domains I and II. Any effect of recrystallization during F₃ folding on CPOs from S₂ fabrics is unclear; however, on the basis of subparallelism of F₃ sheath folds with the L₂ stretching lineation (i.e., Section 4.3), we infer that the D₂ and D₃ transport directions were parallel to each other and that F₃ recrystallization would probably have enhanced S₂ CPOs.

(2) D₄ Shear Zones. Coarse-grained quartz layers commonly contain obliquely recrystallized grains that form a prominent shape preferred orientation aligned parallel to internal subgrain boundaries. These aligned grains are at a high angle to S₂ and parallel S₄ crenulations in micaceous layers and the axial planes of F₄ folds. Quartz recrystallization textures from D₄ everywhere in domains I, II, and III provide evidence of regime 3 dislocation creep [85] and subgrain rotation recrystallization to grain boundary migration transition of Stipp et al. [83] (Figures 10(d)–10(f)). Although D₄ temperatures are difficult to constrain, quartz microstructures, slip systems, and CPOs (see [35]) from D₄ structures in the MSZ (domains I–III) are indicative of upper greenschist to amphibolite grade deformation temperatures, probably equivalent to the warmer range of temperatures during D₂ (⁠$~500±100$°C). These D₄ temperatures were probably obtained after prolonged (~86–73 Ma, [31]) crustal thickening during D₂ and early D₃ (cf. [25]).

In contrast to Seno Martínez, quartz CPO results from the BFSZ (Figures 14(a)–14(d)) indicate ubiquitous basal <a> slip with minor evidence for the occurrence of prism <a> slip (i.e., Figure 14(b)) and $c$-axis opening angles range from ~50 to 70°. The sense of shear from CPO patterns is top-northeast in every sample (Figures 12 and 14), in agreement with the macroscopic kinematic indicators from the shear zone (see [36]). Quartz recrystallization textures, slip systems, and $c$-axis opening angles within the BFSZ are indicative of regime 2 dislocation creep of Hirth and Tullis [85] and/or the SGR recrystallization regime of Stipp et al. [84], suggesting deformation temperatures between 375 and 575°C (DV in Figure 16). In samples of the CDMC taken near Estuario Silva Palma, below the BFSZ (domain VI), quartz textural observations, CPOs, and $c$-axis opening angles are indicative of deformation temperatures between ~300° and 475° C (DVI in Figure 16). These estimates are in good agreement with recent pseudosection modeling of mylonitized siliceous metatuff from the Tobífera thrust measured by Muller et al. [40] in two locations between Estuario Silva Palma and Canal Jeronimo. They report peak pressure and temperature conditions of $P=3$–4 kbar and $T=300$–340°C near Estuario Silva Palma and $P=5.3$–6.3 kbar at $T=420$–460°C near Canal Jeronimo (Figure 16).

Progressive, top-northeast D₂/D₃ shearing appears to have occurred regionally near the top of the CDMC at the contact with the volcanic and volcanoclastic rocks of the RVB terrane. At Seno Martínez, Bahía Fortesque, and Bahía Angelito, the contact is sheared and separates polydeformed and metamorphosed CDMC schist from overlying rocks of the RVB terrane that do not express the same degree of deformation and superposed folding. These observations support the interpretation of the MSZ and BFSZ as regional structures that define a décollement at the base of the Patagonian FTB (Figure 16; cf. [36]). The décollement occurs at a stratigraphic and rheological contact between relatively strong ignimbrite and volcanic units of the Tobífera Formation and RVB seafloor with weaker pelitic schists of the underlying CDMC. Deformed rocks of the CDMC within and below the décollement have a similar top-northeast sense of shear during D_2-3 as first-generation structures within the Patagonian FTB near the foreland.

Defined here for the first time, the Magallanes décollement comprises the MSZ near Seno Martínez, the fault zone at Bahía Angelito, and the BFSZ on Peninsula Brunswick and accommodated top-northeast thrusting of the RVB onto the continental margin (i.e., CDMC). In the hinterland part of the orogen, the MSZ comprises several kilometer-wide high-strain zones that accommodated the underthrusting and ductile thickening of the continental margin (CDMC). Toward the foreland, the décollement localized forming the ~1 km thick BFSZ that ramped up section across the brittle-ductile transition (i.e., Bahía Angelito fault zone and Canal Jeronimo fault zones, Figure 16) and transferred displacement into the nascent Patagonian FTB (cf. [25, 36]). Structural observations reported by Betka et al. [36] document brittle-ductile thrust fault zones near Canal Jeronimo, ~75 km northwest of the BFSZ, that uplift the CDMC and imbricate the Tobífera Formation. The northwest decreasing depth of exposure preserves progressively shallower levels of the Magallanes décollement where it crosses the paleo-BDT between Seno Martínez and Canal Jeronimo (Figure 16). Within and below the MSZ, shortening was accommodated by crystal plasticity, while above it and toward the foreland by formation of the FTB. Quartz microstructures and CPOs presented in this study indicate that the deformation temperatures associated with D₂–D₄ stages of the décollement within the MSZ were warmer (~500° to >650°C in the high-strain core of the MSZ) than within and below the BFSZ on Peninsula Brunswick ~40 km in the transport direction (northeast) toward the foreland (375–575°C in the BFSZ), suggesting that the décollement dipped shallowly toward the hinterland (Figures 16 and 17) during its formation.

Integrating the new data presented in this study with previous work [25], we discuss the tectonic significance of the Magallanes décollement and polyphase deformation of the CDMC with respect to the development of the southern Andes (Figure 17). In the Late Jurassic, the RVB existed on the Pacific side of the continental margin of South America but inboard of the Patagonian arc (e.g., [52, 53, 56]). The onset of the Andean Orogeny in the southernmost Andes is defined by the closure and inversion of the RVB [25, 26, 52, 56]. Obduction of the mafic seafloor of the RVB and underthrusting and metamorphism of the continental margin were underway by ~86 Ma (Figure 17(b); [25]). Obduction of the RVB terrane was partly facilitated by the development of the Magallanes décollement that formed near the basement-cover contact. At this time, pre-Andean fabrics (i.e., S₁ at Seno Martínez) were transposed into parallelism with the pervasive composite schistose foliation (S₂) and prominent quartz stretching lineation (L₂) associated with D₂ fabrics near Seno Martínez. Displacement was transferred toward the northeast resulting in the shearing and imbrication of the Tobífera Formation, development of the Bahía Fortesque high-strain zone, and the earliest first-generation thrusts of the Patagonian fold-thrust belt [36]. The D₂ fabrics from the Magallanes décollement are interpreted to be contemporaneous and correlative with first-generation obduction structures described by Klepeis et al. [25] along the Beagle Channel to the south, as well as with the Canal de las Montañas Shear Zone along strike to the north [18]. This interpretation is consistent with previous tectonic studies (e.g., [25, 26]) which indicate that the onset of underthrusting of the continental margin beneath the RVB terrane occurred regionally during the Late Cretaceous.

Continued shortening and underthrusting of the continental margin resulted in the progressive deformation, tectonic burial, and synkinematic metamorphism of the CDMC schists and the transposition of third-generation folds (F₃, D₃) into D₂ structures (cf. [25]). Within the Seno Martínez D₃ high-strain zone (Figure 4, domains I and II, Figure 16), F₃ folds are expressed as both isoclinal and sheath folds that are both compatible with southwest-northeast shortening and bulk top-northeast shearing. Below the D₃ high-strain zone at Seno Martínez, F₃ folds occurred as regional recumbent folds that thickened the CDMC (domain III). Displacement during D₃ was transferred toward the northeast along the Magallanes décollement toward the Bahía Fortesque high-strain zone where D₃ is characterized by the isoclinal folding and shearing of the composite foliation (local S₂/L₂, Figures 4 and 16). The Bahía Fortesque high-strain zone ramped up section and fed shortening into the Patagonian fold-thrust belt (Figures 16 and 17; [35, 36]).

Craton-directed (top-northeast) D_2-3 progressive deformation documented in this study is correlative with a phase of north-vergent first-generation thrusting reported near the Beagle Channel that accommodated the partial obduction of the Rocas Verdes terrane onto the continental margin prior to ~86 Ma ([25]; see also, [26]). Our CPO opening angle thermometry results indicate that deformation conditions during D_2-3 increased from ~300–500°C in DI to 500–650°C northward in the core of the subsequent F4 antiform (DI+gt and south DII), suggesting that the Magallanes décollement occurred at ~20–25 km depth (assuming a 25°C/km geothermal gradient). This result is compatible with thermobarometry from Seno Martínez (⁠$P$ ~9–10 kbars and $T$ ~545–570°C; [30]) as well as additional thermobarometry and pseudosection modeling from amphibolite-grade rocks exposed in the high-grade core of the CDMC near the Beagle Channel that were buried to depths ~35 km (⁠$P=12$ kbar and $T=620$°C, [31]; see also [30]) and were structurally below garnet-bearing rocks near Seno Martínez. Cooler deformation temperatures toward the foreland (BFSZ, 375–575°) and recent pseudosection modeling of metatuffs in the Tobífera Formation (⁠$P$ ~3–4 kbar and $T$ ~300–340°C near Estuario Silva Palma, [40]) indicate that the décollement cut up, feeding displacement into the nascent Patagonia-Fuegian FTB above the brittle-ductile transition (Figures 16 and 17).

Fourth-generation structures at Seno Martínez reflect the complete closure of the RVB and collision of the Patagonian arc by the Paleogene (cf. [25, 26]). At Seno Martínez, D₄ is defined by the sets of upright north- and south-dipping kink bands that overprint D₂ and D₃ structures and record horizontal southwest-northeast shortening and subvertical thickening. D₄ strains localized in domain II where upright, isoclinal F₄ folds refold D₂ and D₃ structures and form a high-strain zone that uplifts garnet-grade rocks in the core of a regional F₄ antiform. Folding of the garnet isograd at Seno Martínez correlates with the antiformal doming of amphibolite-grade isograds near the Beagle Channel (i.e., [25, 30]) that resulted in the uplift and partial exhumation of high-grade rocks in the core of the CDMC by the Paleogene (cf. exhumation fabrics (S₂) of [31]; second-generation bivergent structures of [25]; D₂ and D₃ of [26]). This event resulted in Late Cretaceous–Paleogene cooling of the high-grade core of the CDMC [31, 34]. At Seno Martínez, D₄ shortening is interpreted to cause the uplift and partial exhumation CDMC and be coeval with the propagation of out-of-sequence thick-skinned basement-involved reverse faults that cut first-generation structures of the Magallanes décollement and fold-thrust belt (Figure 17(d); cf. [25, 36]). The timing of D₄ is also coincident with the depositional ages of sediments in the Magallanes foreland basin that reflect rapid Paleogene exhumation and denudation of the Cordillera Darwin [41, 66].

It has long been posited that décollements below foreland FTBs are “rooted” down-dip in distributed ductile shear zones (e.g., the Appalachians, [99]; the Moine thrust, [90, 91]; the Monashee detachment and Canadian Cordillera, [100, 101]; the Himalayan-Tibetan orogen, [102, 103]). Ductile flow of the middle and lower crust contributes to orogenic topography and is thought to govern tectonic evolution of convergent orogens (e.g., [102, 104–108]). For example, Lamb [9] and Giambiagi et al. [15, 16] conclude that distributed ductile thickening of the middle and lower crust can account for variation in topographic relief along the arc of the central Andes. In southern Patagonia, the increasing depth of exposure from the Patagonia FTB to the core of the CDMC has exhumed midcrustal structures that record polyphase deformation of the CDMC and document distributed subdécollement plastic strain, providing an opportunity to study these processes directly in an Andean setting.

We find that shortening below the Patagonian-Fuegian FTB was absorbed by the superimposition of several generations of noncylindrical ductile folds and general shear (D₂–D₄; cf. [25, 26, 36, 37]), exemplifying how distributed plastic strain (below the paleo-BDT) is kinematically linked with the growth of a retroarc FTB in the upper crust. The occurrence of both coaxial and noncoaxial shear, as well as flattening strain types, within the several-km wide MSZ demonstrates that subdécollement strain was three-dimensional and widely distributed. These observations underscore the importance of pervasive 3D plastic flow of the middle and lower crust below and toward the hinterland of a retroarc FTB, highlighting the limitations of two-dimensional cross-sectional balancing techniques for estimating the crustal-scale kinematic evolution and tectonic shortening of orogenic wedges (e.g., [16, 100]). Our results from the Patagonian FTB (cf. [25, 31]) document an ancient example of the deformation conditions and progressive evolution of distributed ductile deformation in the hinterland of an orogenic wedge that may be analogous to processes that are occurring below the BDT in active collisional (e.g., [103, 109]) and noncollisional (e.g., [16]) settings.

A newly defined regional ductile shear zone, the Magallanes décollement, accommodated underthrusting of a continental margin below a retroarc FTB during the Late Cretaceous inversion of a marginal basin and formation of Patagonian Andes. The décollement is defined by the transposition of several generations of noncylindrical folds, northeast-vergent L-S tectonites, C-S fabrics, and C$′$-type shear bands that indicate dominantly top-northeast transport. Below the décollement and toward the hinterland, tectonic shortening was accommodated through polyphase folding and associated thickening of the basement schist. Above the décollement and toward the foreland, shortening was accomplished by the formation of a retroarc FTB. The décollement is a regional structure that is exposed in at least three locations over 100 km² of the Patagonian Andes where it separates Late Jurassic and Early Cretaceous rocks of the RVB above from the underlying Paleozoic CDMC. Quartz microstructures, CPOs, and $c$-axis opening angles related to the décollement record deformation temperatures that decrease by ~at least 100°C from hinterland localities (~500 to >650°C) toward the foreland (~375–575°C), indicating a shallow hinterland dip of the ductile décollement. The Magallanes décollement is a well-exposed example of the structural evolution of polyphase distributed ductile shear below and toward the hinterland of a retroarc fold-thrust belt during an “Andean-style” orogeny.

The structural evolution of the décollement occurred across two phases of Late Cretaceous-Paleogene progressive deformation that resulted in the folding, synkinematic metamorphism, and burial of the underlying CDMC (D₂–D₃). These first phases of deformation (D₂–D₃) are interpreted to record the obduction of the Rocas Verdes Terrane and the transfer of displacement into the Patagonian FTB. A later phase of progressive deformation (D₄) records regional southwest–northeast horizontal shortening and vertical extension that contributed to the uplift and doming of the Magallanes décollement. Thrust-controlled vertical thickening enhanced erosion near the hinterland of the orogen, driving exhumation of high-grade rocks of the CDMC (cf. [31, 41, 66]). The latter stage of deformation is also coincident with a period of thick-skinned out of sequence thrusting the cut the décollement and first-generation thrusts in the Patagonian FTB. D₄ shortening is interpreted to reflect complete closure of the RVB and collision of the Patagonian arc with the continental margin.

Data presented in this paper can be obtained by contacting the corresponding author, PB.

This work is part of a Ph.D. dissertation completed by Betka at the University of Texas at Austin and benefited greatly from many discussions with Ian Dalziel, Constantino Mpodozis, Randy Marrett, and Brian Horton.

The authors declare that they have no conflict of interest.

This work was supported by funding from the National Science Foundation (EAR-0635940 to Klepeis) and by grants awarded to Betka from the American Association of Petroleum Geologists Grants in Aid Program and the Jackson School of Geosciences at the University of Texas at Austin. We thank captains Keri Pashuk and Greg Landreth for safe passage and crucial logistical support aboard their vessel Northanger, as well as the captain and Alejandro González on Cabo Tamar. PB thanks Rachel Bernard for generously sharing her Matlab script for plotting and editing pole figures with MTex. Constructive reviews by Kyle Larson, Mauricio Calderón, and Olivier Vanderhaeghe improved this manuscript.