9 Surrounding the Narragansett Basin are basement rocks that show varying effects of Alleghanian and/or older deformation. Adjacent to the northwestern part of the Narragansett Basin, the Neoproterozoic Blackstone Group (Fig. 1b) was deformed and metamorphosed syn-tectonically with the intrusion of the 599 ± 2 Ma Esmond Plutonic Suite (Thompson et al. 2010), indicating Neoproterozoic tectonism (Dreier 1985). The Blackstone Group also shows younger deformation and associated metamorphism that affected the ~370 Ma Rhode Island Quincy Granite (Hermes and Zartman 1985), indicating Alleghanian tectonism (Dreier 1985). Alleghanian structures recorded in rocks of the Blackstone Group include NNE-trending folds with axial planar biotite-grade foliation and superimposed ENE-trending folds and crenulations (Dreier 1985). Further west of the Narragansett Basin, the Devonian Scituate Granite (Fig. 1b), which intruded the Esmond-Dedham terrane, is deformed only at its western margin along the proposed Hope Valley shear zone (Goldstein 1986; Getty and Gromet 1988).

10 East of the Narragansett Basin in the area of New Bedford, MA (Fig. 1b), Murray et al. (1990) documented penetratively deformed rocks that are lithologically similar to Hope Valley terrane rocks, and suggested that Hope Valley terrane may also be exposed east of the Narragansett Basin in that area. If so, the Esmond-Dedham terrane on which the Narragansett Basin was deposited would be a thrust slice. However, Hermes and Zartman (1992) concluded that most of the deformation in this area is Neoproterozoic.

11 South of the Narragansett Basin, Neoproterozoic metasedimentary rocks that crop out on Newport Neck (Fig. 1b) show several phases of Neoproterozoic and some possibly younger deformation (Skehan and Rast 1990). The Price Neck Formation records two generations of ENE-trending folds (Skehan and Rast 1990), which are Neoproterozoic based on the relationship to the 595 ± 12 Ma Cliff Walk Granite (Smith 1978). Within the adjacent Proterozoic Newport Neck Formation, NW- to N-trending, tight to isoclinal, west-verging F₁ folds with slaty S₁ cleavage have been refolded by N-trending F₂ folds with an associated dominant S₂ cleavage and N-trending broad upright F₃. The age of these structures has been interpreted as either Neoproterozoic (Skehan et al. 1981) or post-Cambrian (Skehan and Rast 1990). The latter interpretation is based on the observation that the Cambrian rocks were overthrust by the Neoproterozoic rocks prior to F₂; however, Webster et al. (1986) and Landing (1996) interpreted the boundary as an overturned unconformity rather than a thrust.

12 In previous work on the Cambrian rocks in Beavertail State Park, three formations were recognized in the Conanicut Group, from oldest to youngest, the Jamestown, Fort Burnside and Dutch Island Harbor formations (Skehan et al. 1981; Skehan and Rast 1990). The Jamestown Formation contains the Beavertail Point, Hull Cove, and Lionhead members, and the Fort Burnside Formation contains the Short Point and Taylor Point members; no members were recognized in the Dutch Island Harbor Formation. Earlier workers considered these rocks to be of Pennsylvanian age (e.g., Nichols 1956; Quinn 1971), until the discovery of Middle Cambrian trilobites of Acado-Baltic affinities in the basal Jamestown Formation (Skehan et al. 1978). Dikes in the area were described by Nichols (1956) as lamprophyre (see also Skehan and Rast 1990).

13 Previous field research by Murray and Skehan (1979) recognized NE-trending, tight to isoclinal F₁ folds with axial planar, slaty cleavage refolded by NNE-trending, open and overturned F₂ folds. A dominant, W-dipping S₂ axial planar cleavage transects F₁ folds. Murray and Skehan (1979) also postulated later NE-trending F₃ folds, and describe a NW-trending mineral lineation on S₂ and kink bands. Skehan et al. (1981), however, interpreted the earliest folds and numerous “tectonic slides” as soft sediment slump-related structures. Skehan and Rast (1990) repeated this interpretation and further proposed that thrust faults with eastward transport occur on west-dipping S₂ cleavage planes that offset upright F₁ folds and likely caused the folding and boudinage of quartz veins. Most authors (e.g., Murray and Skehan 1979; Burks 1981; Skehan and Rast 1990) attributed all but the first generation structures to the Alleghanian orogeny. Skehan and Rast (1990) suggested that the Cambrian rocks were thrust northwestward over the Pennsylvanian rocks along the BHSZ during D₁ and that NW- to N-directed thrusting was responsible for the deformation of both the Cambrian rocks and Pennsylvanian rocks of the Narragansett Basin.

14 The southern part of the Pennsylvanian Narragansett Basin (southern Rhode Island) shows multiple phases of deformation and Barrovian-type metamorphism that ranges in grade from sillimanite in the west to chlorite and biotite in the east (Mosher 1983). The metasedimentary rocks of the northern part of the basin (northern Rhode Island and Massachusetts) and adjacent Norfolk Basin (Massachusetts) are at a lower metamorphic grade and less intensely deformed with upright, open, ENE-trending folds that parallel the northern basin margins (Quinn and Oliver 1962; Lyons et al. 1976; Cazier 1987).

15 In the southern part of the basin, the first phase of deformation (D₁) formed open to isoclinal F₁ folds and an axial planar fabric (S₁). F₁ folds generally trend NNE (to N) and verge westward; associated NNE-striking thrusts show westward transport (Mosher 1983, 1987). Rocks deformed at high metamorphic grades preserve two distinct fold generations, F_1a and F_1b, and axial planar foliations, S_1a and S_1b, respectively (Henderson and Mosher 1983; Reck and Mosher 1988; Snoke and Mosher 1989; Mahler-Cogswell and Mosher 1994). S_1A and S_1B are penetrative foliations observed across the basin. At high metamorphic grades the foliations are defined by parallel muscovite, biotite, and elongate, ductilely deformed quartz (Mahler-Cogswell and Mosher 1994) and at low metamorphic grades by pressure-solution cleavage with fine-grained, parallel muscovite and chlorite and elongate pressure-solved quartz (Farrens 1982). The second phase of deformation (D₂) formed open to tight F₂ folds of bedding, S_1A and _1B, and an axial planar S₂ crenulation cleavage enhanced by pressure solution. F₂ folds trend NNW to NNE and verge eastward (Farrens 1982; Mosher 1983). Peak metamorphism occurred after D₁, and generally prior to D₂ (Mosher 1983; Snoke and Mosher 1989).

16 D₃ and D₄ are primarily developed in the NE-trending, vertical Beaverhead shear zone (BHSZ) (Mosher 1983; Mosher and Berryhill 1991; Burks and Mosher 1996), although the effects are also more broadly distributed in the high-grade metamorphic rocks (Reck and Mosher 1988; Mahler-Cogswell and Mosher 1994). Within the zone, NNE-trending sinistral strike-, oblique-, and dip-slip shear zones and faults are related to D₃. F₃ folds generally trend NE to E, and an associated S₃ foliation was observed only at the highest grades (Mosher 1983; Mosher and Berryhill 1991; Mahler-Cogswell and Mosher 1994). During D₃, clockwise-younging superposed crenulations and mesoscopic folds formed in discrete zones (Mosher 1983; Mosher and Berryhill 1991; Burks and Mosher 1996). Kink bands displaying normal motion are spatially related to faults striking N to NNE, and are overprinted by D₄ structures (Mosher and Berryhill 1991).

17 Dextral NE-striking strike-slip and ENE-striking oblique- and dip-slip shear zones and faults are related to D₄. During D₄, pre-existing fabrics and structures are rotated into the BHSZ, and superposed anti-clockwise-younging crenulations and mesoscopic folds formed within the BHSZ and overprint D₃ structures (Mosher 1983; Mosher and Berryhill 1991; Burks and Mosher 1996). At the highest metamorphic grades, N- to NE-trending open F₄ folds were observed to reorient S₂ and F₃ folds (Reck and Mosher 1988; Mahler-Cogswell and Mosher 1994). Kink bands spatially related to E- to ESE-trending faults show normal motion (Mosher and Berryhill 1991). All across the southern part of the basin, centimetre- to decametre-scale boudins record N-S extension late during D₄. This boudinage likely formed synchronously with the ENE-striking dip-slip fault motion (Farrens 1982; Mosher 1983; Mosher and Berryhill 1991). Additionally, NNW- to NNE-striking normal faults related to D₃ and D₄, or to Triassic rifting, juxtapose rocks of different metamorphic grades (Mosher 1983).

19 On a map scale, bedding locally changes dip over 10–20 m and forms upright folds and monoclines that are transected by S₂ (Figs. 2, 3). Because these folds and monoclines were not observed to fold S₂, they are interpreted to be the earliest fold phase (F₁). Map-scale F₁ folds and monoclines deform bedding on a wavelength of 3–15 m and amplitude of 1–2 m; some axial planes dip moderately to the east. F₁ interlimb angles range from open to isoclinal as a result of refolding by the later F₂ fold phase. The range of F₁ inter-limb fold angles (0–90°) is locally dependent on the relative size and position of individual F₁ and F₂ fold hinges and limbs.

20 In outcrop, the first fold generation (F₁) consists of symmetrical and asymmetrical folds that deform bedding (S₀). F₁ folds are typically observed on S₂ surfaces that transect S₁ axial surfaces and F₁ hinge lines, suggesting that F₁ is non-coaxial with F₂ (Figs. 4–6). Common wavelengths and amplitudes for F₁ folds in outcrop range from 5–15 cm. F₁ folds were not observed in the Lionhead or Hull Cove members of the Jamestown Formation, but were observed in all other stratigraphic units. In the Beavertail Point member of the Jamestown Formation and the Fort Burnside Formation, F₁ folds are open to isoclinal. In these stratigraphic units, quartzite layers form parallel folds, but in mica-rich layers the hinges of F₁ folds have thickened (Fig. 5). F₁ folds in the Dutch Island Harbor Formation are cylindrical with tight to locally open inter-limb angles (Figs. 3, 6). Few F₁ axes were measured because most were visible only on S₂ foliation surfaces. However, F₁ axes that could be measured have steep to shallow plunges and generally fall on a very broad N-S girdle (Fig. 7a). In addition, measured axial planes associated with F₁ are typically upright and strike N-S. Poles to F₁ axial planes fall on a diffuse E-W girdle (Fig. 7a) and define a SSE-trending fold axis, consistent with F₂ fold axes.

21 Early faults developed synchronously with F₁, as indicated by the observation on an S₂ surface of centimetre to decimetre offset along one limb of an F₁ fold (Fig. 8). Additionally, early faults complicate stratigraphic relationships by juxtaposing formations in the wrong stratigraphic order. Bedding is truncated across the early faults that disrupt the stratigraphy, confirming that they are faults, not stratigraphic contacts (Fig. 9). For example, folded bedding planes of the Short Point member are truncated by an early-stage fault and juxtaposed with the Beavertail Point member (Fig. 9). These faults are folded by F₂, but not F₁ and are interpreted as syn-F₁ faults; similar to the intraformational faults observed on S₂ surfaces. No thrust faults are observed on S₂ cleavage planes, and all disruption of F₁ folds is attributed to syn-F₁ faults.

22 The S₁ fabric forms a continuous cleavage defined by the alignment of platy minerals. S₁ is axial planar to F₁ folds and is poorly developed in outcrop except within and near F₁ fold hinges of mica-rich layers (Fig. 5). It is rarely present in quartzite layers. Locally, S₁ is better developed than S₂, but only within mica-rich layers of the Beavertail Point member and Dutch Island Harbor Formation.

23 In thin section, compositional layering (S₀) is typically recognizable as light layers, composed of quartz and carbonate, with minor amounts of fine muscovite alternating with dark layers made up of fine muscovite and lesser quartz (Fig. 10). Both layers contain chlorite, and siliceous and/or siderite nodules. Pyrite crystals are most abundant in the Dutch Island Harbor Formation, and less abundant in other formations. Graphite is more abundant in darker mica-rich layers, but is also present in lighter layers. Features interpreted to be microstylolites are locally present sub-parallel to S₀, but could also be crenulated graphite layers.

24 S₁ is a penetrative metamorphic fabric in thin section, defined principally by a very well-developed alignment of fine muscovite. Some larger blades of muscovite and minor amounts of chlorite also define this fabric (Fig. 10). Where present, S₁ is generally sub-parallel to bedding except in the hinges of F₁ folds. S₁ is also preserved as mineral inclusions in siderite and siliceous nodules that overgrew S₁ (Fig. 11). In siderite nodules, S₁ is defined by aligned inclusions of pyrite, white mica, rare chlorite, and elongate quartz. In siliceous nodules, S₁ is defined by aligned inclusions of mostly elongate quartz, minor fine muscovite, and rare chlorite (Fig. 12). S₁ is everywhere evident in mica-rich layers, but S₁ is typically not well developed in quartz-rich layers where it is defined by rare muscovite grains.

25 The dominant fold generation, F₂, deforms bedding into E-verging, overturned, and locally recumbent folds. Associated with F₂ is an axial planar cleavage (S₂), which is commonly the most pervasive fabric in outcrop (Figs. 9, 13). F₂ folds bedding, F₁, S₁, and the early faults that truncate F₁ fold limbs and disrupt the stratigraphic order. F₂ folds are typically asymmetrical with inter-limb angles ranging from open to tight. Centimetre to decimetre-scale F₂ folds are commonly parasitic to metre-scale F₂ folds and have wavelengths of 0.5 cm to 5 m with amplitudes ranging from 0.5 cm to 5 m. F₂ fold axes trend NNE to NNW and plunge shallowly (0–10°) north and south (Fig. 7a). Many F₂ folds have periclinal fold axes that split or die out along the hinge line and are doubly plunging, which most likely causes the variable plunge of F₂ fold axes (Fig. 7a). Bedding poles fall on a broad E-W girdle, consistent with F₂ fold axes orientations (Fig. 7a). In general, F₂ folds have a similar trend as F₁ folds, but are typically non-coaxial because of the generally steeper plunge of F₁ folds (Fig. 7a). In outcrop, interference patterns of F₁ and F₂ are rare, but where present confirm non-coaxial fold generations as a result of a difference in plunge (Fig. 4).

26 In two localities along the southeastern coastline of the field area, the Beavertail Point member of the Jamestown Formation is in contact with the Dutch Island Harbor Formation and folded by F₁ and F₂ (Figs. 2, 4). Skehan et al. (1981) interpreted this contact as an early fault associated with F₁ on the basis of their stratigraphic interpretation (Beavertail Point member as the oldest unit and the Dutch Island Harbor formation as the youngest unit). However, in our field and petrographic observations we found no evidence of brecciation along this boundary or any other evidence of faulting (Fig. 4). Thus we interpret it as a stratigraphic contact between the Dutch Island Harbor Formation and the Beavertail Point member of the Jamestown Formation. With this interpretation, we suggest that the major structure cropping out along the southeast coastline of the field area is a synformal, E-vergent F₂ fold cored by the Beavertail Point member (Fig. 3).

27 In addition to folds of bedding, S₂ is axial planar to folds of two different styles of quartz veins, indicating that the veins were emplaced prior to, and later folded by, F₂ folds. En échelon quartz-filled extension fractures are deformed into tight to isoclinal folds. Such veins are typically 1–3 cm in thickness, and when unfolded range in length from 0.5–2 m. Other quartz veins folded by F₂ were likely planar structures first before being folded by F₂. Such veins are typical where S₂ is shallowly dipping and are recumbently folded, with inter-limb angles that are closed to isoclinal. These folds have wavelengths varying from 1 cm up to 3 m, amplitudes ranging from 1 cm to 4 m, and vein thicknesses of 1–4 cm.

28 The penetrative S₂ cleavage is present across the entire field site (Fig. 1c). In outcrop, S₂ forms a smooth, continuous pressure solution cleavage (e.g., Figs. 9, 13). This cleavage is axial planar to F₂ and transects F₁ axial planes (Figs. 3–6), verifying that it formed after F₁. S₂ generally strikes NNE and is shallowly to moderately west dipping (Fig. 7b).

29 In thin section, the S₂ fabric is a crenulation cleavage of S₁ enhanced by pressure solution and microstylolites. Within mica-rich layers, S₂ forms tight to isoclinal crenulations of S1 that are parasitic to mm-scale folds of S₀ and S₁ (Fig. 10). Seams of hematite and graphite are parallel to and truncate limbs of the crenulations (Fig. 10). Pervasive pressure solution along the crenulation limbs has resulted in a decrease of the inter-limb angle from that observed in outcrop. On the limbs of mm-scale folds, parasitic folds of S₁ are asymmetric, whereas in the hinge they are symmetric. Wavelengths for these folds range from 0.08 mm to 0.5 mm, with amplitudes between 0.1 mm to 0.5 mm. S₂ is better developed in mica-rich layers than quartz-rich layers, where S₂ is best observed as microstylolites. In mica-rich layers, blades of white mica bend around fold hinges, but in quartz-rich layers individual white mica and muscovite grains have been recrystallized in fold hinges. The very close spacing of S₂ indicates why S₂ is commonly more pervasive than S₁ at the outcrop scale (Fig. 10).

30 The S₂ fabric wraps around the siderite and siliceous nodules (Figs. 11, 12), and pressure shadows composed of mostly muscovite, elongate quartz, and calcite with minor amounts of chlorite and rare clinozoisite have grown off of them parallel to S₂ (Fig. 12). Clusters of pyrite crystals that grew prior to S₂ also have pressure shadows composed primarily of mica, quartz, and iron carbonate aligned parallel to S₂.

31 The youngest fold generation, F₃, is recorded by map-scale gentle to open folding of S₂. This fold generation is inferred from changes in S₂ orientation along continuous exposure on the coastline of Beavertail State Park (Fig. 1c). On the western coastline of the park, the S₂ fabric dips moderately to the west, whereas it is shallowly dipping on the eastern coastline (Figs. 1–3). Poles to S₂ and F₂ axial planes spread along a WNW to ESE-striking girdle defining a NNE-trending fold axis (Fig. 7b), consistent with a shallowly plunging NNE-trending F₃ fold axis. Although defined by the change in S₂ orientation, F₂ and F₁ folds, S₁ and early faults are also observed to be folded by this generation. For example, F₂ folds change from overturned to recumbent across the F₃ folds (Figs. 2, 3).

32 In thin section, mica-rich layers of the Short Point member contain post-S₂ crenulations. These crenulations formed only on highly pressure-solved limbs of S₂ crenulations along very mica-rich bands. The crenulations are at moderate to high angles to S₂ (Fig. 14) with wavelengths of <0.075 mm and up to 0.15 mm in amplitude. Multiple sets of differently oriented, and locally cross-cutting, crenulations are observed within the same thin section. Locally, crenulations curve from low to high angles relative to S₂, and some are localized between siderite nodules, indicating that shearing may have accompanied their development. No spatial relationship to F₃ is observed; hence we term these S _C , because their relationship to post-F₂ deformation is unknown.

33 Boudinage commonly warps bedding and the S₂ fabric. Boudins are generally 1–3 m in length, but locally are up to 15 m. Most boudins are rounded with flat edges at their necks. Quartzite layers acted competently, whereas phyllitic layers were more ductile. Within completely mica-rich areas, the layers display foliation boudinage and blocky quartz veins are common in neck regions. The trend of several boudins displays predominantly N-S extension.

34 After the formation of F₂ and S₂ and before brittle faulting, another generation of quartz and/or siderite veins formed. The veins vary from blocky to fibrous and some contain chlorite as well. Styles of veining include quartz and/or siderite filled en échelon extension fractures, planar quartz and siderite veins, and large blocks of vein quartz with minor chlorite. The latter blocks range in size from 1–3 m ² and are typically separated by 5–15 m. Because vein quartz is commonly present at the necks of several small-scale boudins, this observation suggests that these blocks are necks of large-scale boudins. In addition, sigmoidal and planar en échelon quartz veins cross cut S₂. Arrays of these veins typically range from 0.5–1 m but are up to 10 m in length with widths of 15–30 cm to 5 m. Spacing between en échelon fractures is typically 1–3 cm, but can be up to a metre. Within an array, thicknesses of individual veins range from 1–15 cm. Planar veins that commonly cut S₂ vary from 0.1 cm to 10 cm in thickness. Planar veins display a variety of orientations, but generally strike NW and dip moderately SW suggesting NE-SW extension (Fig. 15a). The wide variety of quartz veins that were emplaced prior to, during, and after F₂ indicates that silica was mobile throughout deformation.

35 A 1–2 m thick lamprophyre dyke in Beavertail State Park cross-cuts S₂ and all fold generations (Fig. 1c). The dyke is composed of quartz, carbonate (replacement of quartz), feldspar, and muscovite with minor pyrite. It strikes roughly NW-SE and dips rather shallowly (~30°) to the SW and is not reoriented along with S₂ by the map-scale F₃ fold, indicating its emplacement post-dated F₃. Its orientation is similar to that of most planar quartz veins and suggests that it may be associated with the NE-SW extension related to the veins. Locally, the dyke displays pinch and swell type boudinage over 3–4 m. Also, it is cut by brittle faults and pervasively by several quartz and calcite veins.

36 Brittle faults and related fractures with no demonstrable offset cut all fold generations (F₁, F₂, and F₃), their associated fabrics (S₁ and S₂), the dyke and all other structures. Faults and related fractures strike N to NE and ENE to ESE (Figs. 2, 15b). Sense of motion on faults was determined from quartz and chlorite slickenlines, deflection of S₂ offset beds and/or displacement on spatially associated kink bands (see below). Minor sinistral and dextral motions were observed, but normal motion is predominant (Fig. 15c). T-axes generally trend NW-SE and P-axes have a variety of orientations, but are mostly steeply plunging. These data indicate normal motion with extension in the NW-SE direction. Although previous maps (e.g., Skehan et al. 1981) and those from this study (also see Carter 2008) show a large number of steeply dipping “faults” in these same orientations; most do not show demonstrable offset of structures, unit contacts or layers, and commonly structures continue across the “fault” with no offset. These features are visually prominent because of the intense wave erosion along this coastline, but without evidence of offset (or with evidence of no measurable offset) should best be termed fractures. The largest offset measured on any of the faults is on the Beavertail fault of Skehan et al. (1981), which has no more than 5 m of normal throw based on offset of the prominent NW-trending lamprophyre dyke.

37 Kink bands of S₂ are spatially associated with the brittle faults (Fig. 15b) and vary from a few to multiple arrays covering areas up to a metre wide. Kink bands range from 5 mm to 10 cm across, though on average are 1–3 cm wide. They typically strike sub-parallel to faults, trending NNE to ENE. In some localities, kink bands adjacent to faults are oriented oblique to the fault by ~10–30° and form en échelon arrays indicating oblique motion; some indicate faults with sinistral and others with dextral components of motion. Almost all kink bands record a NW side up motion of no more than 1–2 cm, consistent with the NE-strike of normal faults, predicted from the T and P axes.

38 The vertical Beaverhead shear zone (BHSZ), located in the northwestern part of the field area, juxtaposes the Cambrian formations with Pennsylvanian formations of the Narragansett Basin and affects the S₂ fabric (Fig. 1c). Near the BHSZ, the strike of S₂ in the Cambrian rocks shows progressive clockwise rotation (Fig. 1c). The deflection of S₂ as it approaches the BHSZ is consistent with dextral motion and indicates that S₂ developed prior to the last recorded dextral motion on the zone (Mosher 1983; Mosher and Berryhill 1991; Burks and Mosher 1996).

39 The dominant metamorphic mineral is muscovite which is aligned parallel to S₁. The only metamorphic index mineral present is chlorite, and it is either sub-parallel or at high angles to S₁. In addition, rare tails of chlorite and clinozoisite have grown parallel to S₂ off pyrite, siderite, and siliceous nodules. Round grains that resemble isolated sand grains are composed of chlorite. The freshness and distinct grain shape of the chlorite suggest that it replaced a preexisting detrital mineral of a different composition during metamorphism. Given the absence of biotite, the metamorphic grade of this area could not have been higher than lower greenschist facies, and peak temperatures occurred during the formation of S₁, and continued during the development of S₂. Whole rock ⁴⁰ Ar/ ³⁹ Ar cooling ages for muscovite which defines the S₁ fabric in the Cambrian rocks ranges from 240 to 260 Ma (Dallmeyer 1982).

40 Because the Pennsylvanian metasedimentary rocks are intensely deformed by D₃ and D₄ within the Beaverhead shear zone and to a lesser extent along other faults and shear zones, it is appropriate to compare Pennsylvanian rocks on Conanicut and Aquidneck islands (Fig. 1b) that lie outside the BHSZ with the Cambrian rocks south of the shear zone. Garnet- to staurolite-grade metasedimentary rocks on Conanicut Island to the north are geographically closest, but rocks on Aquidneck Island to the east are at the same metamorphic grade as the Cambrian rocks and hence more directly comparable. The summary below is based on extensive research by S. Mosher and her students (also see Farrens 1982; Henderson and Mosher 1983, and summary in Snoke and Mosher 1989).

41 In Pennsylvanian rocks on Conanicut Island, the first deformation produced F _1a and F _1b folds and associated S _1a and S _1b axial planar foliations, whereas on Aquidneck Island only one generation (F₁ and S₁) was produced. F _1b and F₁ are generally tight to isoclinal, NNE-trending, WNW-verging, shallowly plunging folds (Fig. 7c, Table 1). S _1a,b and S₁ are defined by very well-developed parallel arrangement of muscovite and chlorite in phyllite and schist and pressure solution cleavage in more quartz-rich rocks. Open to tight, NNE-trending, shallowly plunging, ESE-verging F₂ folds deform S _1a,b and S₁ and are nearly coaxial to F _1b and F₁, respectively (Fig. 7c, Table 1). F₂ folds have an associated axial planar crenulation cleavage, S₂, enhanced by pressure solution. F₂ folds are more numerous on Aquidneck Island. In the higher metamorphic grade Pennsylvanian rocks on Conanicut Island, garnet, staurolite, and biotite grew between D₁ and D₂.

42 Evidence of D₃ and D₄ are expressed on both islands, but to different extents. On Conanicut Island, numerous NNE-trending sinistral oblique-slip faults are observed. Adjacent to the faults, NE-trending, shallowly plunging, upright to recumbent F₃ folds (Fig. 7d, Table 1) with an axial planar crenulation cleavage is present, along with sinistral tension gashes and sheath folds. On Aquidneck Island, S₁ and S₂ are affected only by E-trending kink bands, although the spread of S₂ along an ESE-trending girdle suggests later open refolding (Fig. 7d, Table 1). On both islands, late-stage boudins with N-S extension affected the metasedimentary rocks. Boudinage is more prevalent on Aquidneck Island where decametre-scale foliation boudinage is observed. In some localities on Aquidneck Island, clockwise younging superposed crenulation cleavages are present. Lastly, on both islands, earlier structures are cut by N-trending, steeply dipping brittle normal faults.

43 Structures and associated foliations and cleavages affecting the Cambrian rocks are very similar in style, orientation, relative timing, and deformation fabrics to those affecting the adjacent Pennsylvanian rocks on Conanicut and Aquidneck islands. Previous workers have suggested correlation of all but the first generation structures, and by implication that F₁ in the Pennsylvanian rocks equates to F₂ in the Cambrian rocks. However, F₁ folds affecting the Pennsylvanian rocks verge westward, whereas F₂ folds in the Cambrian rocks verge eastward. It is highly unlikely that correlative fold generations in adjacent localities would have different vergences. Additionally, F₂ in both Pennsylvanian and Cambrian rocks verges eastward and has similar deformation fabrics. Hence, correlation of F₂ between localities is more plausible. F₁ folds in both areas are tectonic in origin and have the same orientations, styles, and minerals defining the fabrics. Given the proximity of these rocks, it is unlikely that the Cambrian rocks escaped the more penetrative first deformation to affect the Pennsylvanian rocks. Thus we correlate F₁ folds in the Cambrian and Pennsylvanian rocks.

44 Subsequent structures (including faults, kink bands, multiple cross-cutting crenulations, and boudins) are also similar and readily correlated. The gentle to open F₃ folds affecting the Cambrian rocks may be correlative with the F₃ folds associated with D₃ sinistral shear along faults in the basin, but more likely are similar to the later N-trending F₄ folds observed further west at Bonnet Shores in the basin (Fig.1b) associated with D₄ (Reck and Mosher 1988). On the rest of Conanicut and Aquidneck islands, exposure is primarily restricted to N-trending outcrops which are not optimal for viewing gentle to open folds in this orientation, although as noted above, S₂ on Aquidneck Island appears folded on a regional scale. On Bonnet Shores, gentle to open N-trending F₄ folds were observed to reorient the S₂ crenulation cleavage. Normal, sinistral, and dextral faults in the same orientations affect both Cambrian and Pennsylvanian rocks. On the basis of the remarkable similarity of structures in all respects, we propose that all deformation recorded in the Cambrian rocks occurred during the Alleghanian orogeny, or possibly later in the case of normal faults.

45 The relative timing of metamorphism and deformation is the same in both the Pennsylvanian and Cambrian rocks. Additionally, ⁴⁰ Ar/ ³⁹ Ar muscovite dates of about 250 Ma from Aquidneck Island (Pennsylvanian rocks) and from whole-rock phyllite samples from Beavertail State Park (Cambrian rocks), which are at the same metamorphic grade, are the same within error and show similar age spectra (Dallmeyer 1982). All the ⁴⁰ Ar/ ³⁹ Ar muscovite cooling ages are 240 to 260 Ma across the range of metamorphic grades in the Narragansett Basin (Dallmeyer 1982).

46 Our results are more compatible with the preliminary work by Murray and Skehan (1979) and Burks (1981) than later work by Skehan et al. (1981) and Skehan and Rast (1990). Detailed petrographic work described above demonstrates unequivocally that F₁ folds are tectonic structures and S₁ a metamorphic fabric, not soft-sediment slump features. During our mapping at a scale of 1:1000, we observed no thrust faults on S₂ cleavage planes. Also, thrusting along S₂ could not have caused folding of quartz veins as S₂ is axial planar to the folded quartz veins. The early faults we observed are folded by F₂ and are associated with the formation of F₁, truncating one limb of F₁ folds.

47 We also found no evidence that the Cambrian rocks were thrust northwestward over the Pennsylvanian rocks along the BHSZ during D₁ as proposed by Skehan and Rast (1990). Thrusting and fold vergence in the Narragansett Basin during D₁ was to the west (e.g., Mosher 1983, 1987), not northwest. In addition, detailed mapping of the Cambrian rocks adjacent to the BHSZ and Pennsylvanian rocks on the eastern shore of Conanicut Island adjacent to the trace of the BHSZ by Henderson and Mosher (1983) did not find any evidence for thrusts. On the western shore, no exposures were found near the contact, and no thrusts are observed. Pennsylvanian rocks in Fort Getty State Park lie within the vertical BHSZ and are intensely deformed by both D₃ and D₄.

48 Deformation in the Narragansett Basin is interpreted to have resulted from closure of the basin as a result of final collision of Africa with North America, with first west-directed deformation (D₁) followed by back-folding with eastward vergence (D₂) (Mahler-Cogswell and Mosher 1994). Subsequent interaction between the two continents is interpreted to have caused sinistral shear along NNE-trending zones in the western passage of Narragansett Bay and the western basin margin (D₃), followed by dextral shear along a NE-trending zone, including the BHSZ and southwestern part of the basin (Mahler-Cogswell and Mosher 1994). The deformation of the Cambrian rocks matches that of the basin outside of the intense shear zones, and thus is compatible with this tectonic model.

49 The proposal by Skehan and Rast (1990) and Rast and Skehan (1990) that the BHSZ is a terrane boundary between the Esmond-Dedham and Bulgarmarsh terranes is not substantiated by our work or other work in the basin. Our work also does not support the suggestion of Skehan and Rast (1990) that the Cambrian rocks are part of the Meguma terrane which was thrust over the Pennsylvanian rocks along the BHSZ from some distance to the southeast as a result of a north- to northwest-directed collision during D₁ and followed by a north-directed D₂. No structures in the southern Narragansett Basin or adjacent Cambrian rocks show evidence of NW- or N-directed transport. As discussed above, transport during D₁ in the southern part of the basin was westward and during D₂ was eastward. The northern part of the basin located far from the BHSZ shows ENE-trending fold axes, but they follow the basin margins and were likely caused by closure of the basin and not likely to be related to D₁ or D₂ in the south (Mosher 1983). Additionally, the low-grade, open to tight, NNE-trending, E-verging F₂ folds with associated crenulation cleavage enhanced by pressure solution observed in both Cambrian and Pennsylvanian rocks could not have formed as a result of N-directed thrusting.

50 Moreover, Pennsylvanian rocks of the Narragansett Basin lie on both sides of most of the BHSZ (Fig. 1b), and the grade of metamorphism is similar across the zone. The chlorite-grade Cambrian rocks (this study) are adjacent to biotite-grade Pennsylvanian rocks on Beaverhead at Fort Getty State Park (Burks 1981) and eastern Conanicut Island (Henderson and Mosher 1983); chlorite-grade rocks on Aquidneck Island (Farrens 1982) are adjacent to biotite-grade rocks on Prudence Island (Thomas 1981) (see Fig. 1b for locations). Also, the Pennsylvanian fossils have the same age range (Stephanian) on both sides of the shear zone (Skehan et al. 1986). Thus, no evidence exists that the BHSZ is a major thrust fault. If a major terrane boundary exists underneath the Narragansett Basin, terrane accretion occurred prior to the formation of the basin; otherwise pre-Alleghanian structures would likely be preserved in the Cambrian rocks.

51 We propose that the Cambrian sediments were deposited on older Neoproterozoic Avalonian basement and remained undeformed until the Alleghanian orogeny. We suggest that deformation related to the accretion of this part of Avalonia to composite Laurentia was accommodated in the western margin of the SENE Avalon terrane (i.e., the Hope Valley terrane) and that the Cambrian rocks escaped deformation until the Alleghanian orogeny. The Pennsylvanian sediments of the Narragansett Basin were deposited on both Neoproterozoic Avalonian basement and on Cambrian sedimentary rocks, at least in the southernmost part of the basin.