Age and tectonic setting of the Quinebaug-Marlboro belt and implications for the history of Ganderian crustal fragments in southeastern New England, USA

In New England, the Putnam-Nashoba terrane and its lowermost part, the Quinebaug-Marlboro belt, occupies the easternmost part of Ganderia, where it structurally overlies Avalonia. The presence of arc-related mafic rock geochemistry, positive εNd(500) values, 1.6–1.8 Ga model ages, and a single Cambrian U-Pb zircon age led to the interpretation that the Putnam-Nashoba terrane coincides with the trailing edge of Ganderia (Hibbard et al., 2006; van Staal et al., 2009, 2012; Pollock et al., 2012; Kuiper, 2016; Kay et al., 2017) (Fig. 1). The entire Putnam-Nashoba terrane is interpreted as a Ganderian cover sequence that was separated from its Ganderian basement by tectonic wedging of Avalonia in the Permian during the Alleghanian orogeny (Wintsch et al., 2014). The unique position of the Quinebaug-Marlboro belt thus provides one of the potentially most complete records available for the eastern trailing edge of Ganderia, before the arrival of Avalonia, in the northern Appalachians. Much of that record is obscured, however, and until only recently, the Quinebaug-Marlboro belt was considered to contain “rocks of unknown affinity” (Hibbard et al., 2006). This is in large part due to the challenges in finding samples suitable for accurate geochronology in highly deformed, fault-bounded, upper amphibolite–facies rocks that lack preserved fossils. These challenges create obstacles to establishing long-distance correlations. The paucity of reliable ages and the uncertain origins of the belt have hampered regional correlations within the Appalachians (Hibbard et al., 2006, 2007a, 2007b; Merschat and Hatcher, 2007; van Staal et al., 2009, 2012; Merschat et al., 2010; Kay et al., 2017), and Ganderia is considered one of the most complex peri-Gondwanan terranes (van Staal et al., 2012).

Advances in the unravelling of Ganderia’s complex evolution show that it developed as an ~2500-km-long arc terrane on the margin of Amazonia during the Ediacaran to Cambrian (Nance et al., 2008; van Staal et al., 2012, 2020). In the middle Cambrian, Ganderia departed from the Gondwanan margin at ~45°S paleolatitude and subsequently accreted to the Laurentian margin at ~20°S in the Late Ordovician to close the Iapetus Ocean (van Staal et al., 2012). This transit is thought to have occurred across Iapetus as Ganderia left the northern margin of Gondwana before being accreted onto Laurentia (Keppie et al., 2012).

The Quinebaug-Marlboro belt represents a poorly understood portion of the vast Ganderian arc system (Goldsmith, 1991b; Hibbard et al., 2006). The complex history of the Quinebaug-Marlboro belt and larger Putnam-Nashoba terrane is tied to arc formation and subsequent amalgamation and modification of Ganderian crustal fragments during at least four Paleozoic orogenic events—the Salinic, Acadian, Neo-Acadian, and Alleghanian (Wintsch et al., 1990, 2007; Hepburn et al., 1995, 2014; Acaster and Bickford, 1999; Stroud et al., 2009; Kay et al., 2017).

Assessment of the nature and origin of protoliths of high-grade metamorphic rocks has long been a challenge (e.g., Hyndman, 1985; Passchier et al., 1990; Wintsch et al., 1990; Frye, 1991). Workers have debated for decades as to whether the protoliths of the rocks in these terranes were sedimentary, volcanic, or plutonic (e.g., Wintsch et al., 1990; Goldsmith, 1991b). The absence of primary structures and the parallelism of layering and metamorphic foliation have led to equivocal interpretations. Distinguishing sedimentary protoliths from igneous protoliths is possible in some cases through geochemistry (Wintsch et al., 1990), but the distinction between volcanic and plutonic rocks is less clear. Crustal fragments underlain by high-grade rocks often represent some of the greatest challenges to plate reconstructions. Unravelling the temporal and spatial history of exotic crustal blocks requires integrated studies involving mapping, geochronology, geochemistry, and structural analysis. The combined characterization of whole-rock geochemistry with high-resolution zircon imagery and U-Th-Pb geochronology has become an essential tool for deciphering protoliths, and thus the origins of rocks from high-grade, polymetamorphic, complexly deformed settings (e.g., Aleinikoff et al., 2006, 2007; Kröner et al., 2014).

In this study, we present new sensitive high-resolution ion microprobe (SHRIMP) U-Pb geochronology results from ten samples using zircon, titanite, and monazite and new major, trace, rare-earth element (REE), and isotope geochemistry from 28 samples to synthesize existing work and establish the age, tectonic setting, and regional correlations for these problematic rocks. Structural analysis is added in places where the context is necessary to characterize the timing of deformation, foliation, folding, and intrusion to mineral growth in these polydeformed rocks. The data include the first reliable primary igneous ages for the Quinebaug and Marlboro Formations. This study demonstrates that by integrating comprehensive geochemistry, precise geochronology with high-resolution imagery, structural analysis, and tectonic synthesis it is possible to unravel the age of enigmatic terranes and establish their origins and complex history.

The ~180-km-long by 10-km-wide Quinebaug-Marlboro belt extends from northeastern Massachusetts to southeastern Connecticut (Figs. 1 and 2) where it occupies the eastern and structurally lowest part of the larger “Putnam-Nashoba terrane” or “Nashoba terrane” (Zen et al., 1983; Rodgers, 1985; Zartman et al., 1988; Goldsmith, 1991b; Kay et al., 2017). The Quinebaug-Marlboro belt consists of Lower Paleozoic metamorphosed volcanic, sedimentary, and plutonic rocks interpreted as an extinct peri-Gondwanan arc complex (Wintsch et al., 1990; Goldsmith, 1991a; Hepburn et al., 1995, 2014; Acaster and Bickford, 1999; Hepburn, 2004; Hibbard et al., 2006; Kuiper et al., 2013; Kay et al., 2017). Current synthesis suggests that the entire Putnam-Nashoba terrane represents an early Paleozoic arc remnant that correlates with the peri-Gondwanan Penobscot arc (Fig. 1; Hibbard et al., 2006, 2007a, 2007b; Johnson et al., 2012; Pollock et al., 2012; van Staal and Barr, 2012; Hepburn et al., 2014; Waldron et al., 2015; Kay et al., 2017). The geochemistry of the mafic to intermediate metamorphosed volcanic and plutonic rocks is interpreted to represent an arc to backarc tectonic setting (Wintsch et al., 1990; Wones and Goldsmith, 1991; Hepburn et al., 1995; Kay et al., 2009, 2011, 2017; Kay, 2012; Hepburn et al., 2014).

The entire Putnam-Nashoba terrane is interpreted as an upper amphibolite–facies, fault-bounded crustal wedge between greenschist-facies rocks in Massachusetts, and greenschist- to amphibolite-facies rocks in Connecticut (Fig. 2) (Zen et al., 1983; Wintsch et al., 1990, 2007, 2014). The terrane was originally conceived to be a westward-dipping and younging homoclinal sequence with the Tatnic Hill and Nashoba Formations at the stratigraphic top and the Quinebaug and Marlboro Formations at the base (Dixon, 1964; Bell and Alvord, 1976; Goldsmith, 1991b). Goldsmith (1991b, 1991c) questioned this scenario because the terrane was fault-bounded, contained only limited evidence for west-facing stratigraphy, had scant age control, and its stratigraphic top and base were unknown. An understanding of the age, origin, and tectonic setting has been problematic for some time due in large part to a lack of fossils and lack of reliably dated rocks (Hepburn and Munn, 1984; Wintsch et al., 1990; Goldsmith, 1991b; Hepburn et al., 1995, 2014; Hepburn, 2004; Hibbard et al., 2006; Kay et al., 2017). Upper amphibolite–facies metamorphism and tectonic thinning of lithostratigraphic belts into fault-bounded slices further complicated attempts to unravel the tectonic history. These complexities led Goldsmith (1991b) and Rast and Skehan (1993) to refer to the entire Putnam-Nashoba terrane as enigmatic because the age, tectonic setting, and metamorphic history were poorly constrained.

At the structural base of the Putnam-Nashoba terrane, the Quinebaug-Marlboro belt structurally overlies rocks of the Avalon terrane along the Honey Hill–Lake Char–Bloody Bluff fault system (Goldsmith, 1991b, 1991c; Wintsch et al., 1992; Figs. 1 and 2). The Avalon terrane contains a Neoproterozoic peri-Gondwanan marginal shelf sequence; Neoproterozoic arc-related plutonic rocks of the Rhode Island batholith; Ediacaran to Paleozoic sedimentary, volcanic, and plutonic rocks; and Carboniferous rocks of the Narragansett Basin (Zen et al., 1983; Zartman and Naylor, 1984; Bailey et al., 1989; Rast and Skehan, 1990; Goldsmith, 1991a; Hermes and Zartman, 1992; Skehan and Rast, 1995; Hibbard et al., 2006; Walsh et al., 2011a, 2011b; Thompson et al., 2012; Walsh, 2014). The Avalonian arc-related rocks range in age from ca. 615–590 Ma (Walsh et al., 2011a, 2011b; Thompson et al., 2012).

Rocks of the Tatnic Hill and Nashoba Formations structurally overlie the Quinebaug-Marlboro belt, and the boundary coincides with the Assabet River fault in Massachusetts and the Tatnic fault in Connecticut (Fig. 2). The Tatnic Hill and Nashoba Formations underlie rocks of the Merrimack belt along the Clinton-Newbury fault (Figs. 1 and 2). Silurian metasedimentary rocks in the adjacent Merrimack belt are considered a cover sequence consisting mostly of marine clastic rocks deposited above extended Ganderian crust (Tremblay and Pinet, 2005). Neoproterozoic Ganderian basement rocks occur outside the Putnam-Nashoba terrane in the Massabesic Gneiss Complex (Dorais et al., 2012) and the Lyme dome (Aleinikoff et al., 2007; Walsh et al., 2007) (Fig. 1), but the basement is not exposed in the Putnam-Nashoba terrane. Largely positive εNd(500) values of +4 to +7.5 for mafic rocks, +1.2 to −0.75 for intermediate to felsic rocks, and −6.0 to −0.8.0 for metasediments, combined with depleted mantle model ages of 1.6–1.8 Ga from the Marlboro and Nashoba Formations in Massachusetts, suggested to Kay et al. (2017) that the underlying crust is Ganderian basement.

The Quinebaug Formation in Connecticut and the Marlboro Formation in Massachusetts form the Quinebaug-Marlboro belt (Fig. 2). The two formations consist largely of mafic metavolcanic and volcaniclastic rocks and minor metasedimentary rocks, including layered and massive amphibolite, hornblende-biotite-quartz-plagioclase granofels, biotite-quartz-plagioclase schist, felsic gneiss, minor coticule, calc-silicate granofels, quartzite, and garnet-sillimanite schist (Dixon, 1964; Bell and Alvord, 1976; DiNitto et al., 1984; Wintsch et al., 1990; Goldsmith, 1991b; Hepburn et al., 1995).

In Connecticut, Dixon (1964, 1965) subdivided the Quinebaug Formation into three members: upper, Black Hill Member, and lower (Fig. 3). The upper and lower members are predominantly metavolcanic amphibolite-facies rocks, and the Black Hill Member consists mostly of calcareous metasedimentary rocks. Rodgers (1985) combined the upper and lower members of the Quinebaug Formation into a single unit but kept the Black Hill Member as a separate unit within the formation (Figs. 2 and 3). The age of the Quinebaug Formation is poorly constrained. In southeastern Connecticut, the Preston Gabbro intruded the Quinebaug Formation in the upper plate of the Lake Char–Honey Hill fault (Fig. 2) (Zartman and Naylor, 1984). A trondhjemite phase, considered to be cogenetic with the gabbro, yielded a concordant isotope dilution–thermal ionization mass spectrometry (ID-TIMS) U-Pb zircon age of 424 ± 5 Ma from two size fractions (Zartman and Naylor, 1984), providing a minimum age for the Quinebaug Formation.

The Marlboro Formation is considered mostly Cambrian in age based on the age of the Grafton Gneiss that intruded it, but younger ages have been reported. The Grafton Gneiss, a biotite granite gneiss within the Marlboro Formation, yielded a SHRIMP U-Pb zircon age of 515 ± 4 Ma, interpreted as a minimum age for part of the Marlboro Formation that was intruded by the gneiss (Figs. 2 and 3) (Walsh et al., 2011a). The Grafton Gneiss contains xenoliths of amphibolite, and its margin is migmatitic and exhibits lit-par-lit sheets of granite in amphibolite, dioritic gneiss, and calc-silicate gneiss of the Marlboro Formation. Acaster and Bickford (1999) dated a rock that they called the “Grafton Gneiss” and obtained a conventional U-Pb zircon age of 584 ± 8 Ma, but the sample location is not clear. Loan (2011) suggests that parts of the Marlboro Formation may range into the Early Ordovician based on preliminary laser ablation–inductively coupled plasma–mass spectrometry (LA-ICP-MS) detrital zircon analyses, which yield results that indicate the youngest grains are older than ca. 470–460 Ma. The LA-ICP-MS method may not be precise enough for some complex upper amphibolite–facies zircons because the beam may sample cores, overgrowths, and rims in 2D and 3D space, thus yielding mixed ages (e.g., Kröner et al., 2014; Schaltegger et al., 2015). Acaster and Bickford (1999) reported multi-grain TIMS U-Pb zircon ages of ca. 473–430 Ma from what they interpret as volcanic rocks in the Marlboro Formation. They report a lower intercept age of 445 ± 4 Ma and interpret a crystallization age between 473 and 445 Ma for the Sandy Pond amphibolite member and an upper intercept age of 425 ± 3 Ma for the Millham Reservoir “granulite” member (a granofels). The reliability of these ages is uncertain, however, because: (1) it is unclear whether the rocks are volcanic or intrusive; (2) it is unclear whether the zircons are igneous or metamorphic, and zircon images were not published; and (3) TIMS analyses of these complicated high-grade rocks may have yielded ages that are mixtures of igneous and metamorphic components. The Quinebaug-Marlboro belt is overlain by the largely metasedimentary Tatnic Hill and Nashoba Formations in Connecticut and Massachusetts, respectively (Figs. 2 and 3), and consists mostly of migmatitic garnet-sillimanite schist and paragneiss, rusty schist, calc-silicate, amphibolite, and local marble (Goldsmith, 1991b, and references therein). The formations contain minor mafic volcanic rocks that increase in abundance toward the basal contact with the structurally underlying Quinebaug and Marlboro Formations (Wintsch et al., 1990; Goldsmith, 1991b). The relative abundance of amphibolite toward the base led earlier workers to interpret the upper contact of the Quinebaug-Marlboro belt to be conformable with the overlying Tatnic Hill and Nashoba Formations (Dixon, 1964; Skehan and Abu-Moustafa, 1976). Subsequent work and depiction on the state geologic maps showed most of the boundary as a fault, with only local occurrences of conformable contacts (Hepburn and DiNitto, 1978; Zen et al., 1983; Rodgers, 1985; Wintsch et al., 1990; Kopera et al., 2006).

A study of metasedimentary rocks in the Tatnic Hill Formation provided the first U-Pb ages determined by SHRIMP analysis on detrital zircon and metamorphic monazite in the Putnam-Nashoba terrane (Wintsch et al., 2007) and indicated that deposition took place in the Silurian after ca. 425 Ma (age of the youngest detrital zircon population) and before 407 Ma (age of the oldest metamorphic zircon and monazite). Loan (2011) determined that the youngest detrital zircon grains in the Nashoba Formation and overlying Tadmuck Brook Schist range in age from ca. 477–461 Ma. In Massachusetts, a possible Silurian minimum age for the Nashoba Formation is provided by the crosscutting Sharpners Pond Diorite dated at 430 ± 5 Ma (multi-grain TIMS U-Pb analyses on zircon; Zartman and Naylor, 1984).

Evidence for prolonged and diachronous exhumation of the Putnam-Nashoba terrane comes from ⁴⁰Ar/³⁹Ar cooling ages of amphibole ranging from ca. 392–310 Ma (Wintsch et al., 1992; Attenoukon, 2009; Reynolds et al., 2010; Reynolds, 2012). These data show that rocks of the Putnam-Nashoba terrane were exhumed from north to south, with older ages (ca. 390–320 Ma) to the north in Massachusetts and younger ages (ca. 340–310 Ma) in Connecticut (Fig. 4). Devonian cooling ages are interpreted to reflect exhumation due to Acadian uplift, whereas Carboniferous cooling ages perhaps are related to uplift and tectonic transport over the adjacent Avalon terrane during Alleghanian tectonics (Wintsch et al., 1992, 1993, 2007, 2012, 2014; Attenoukon, 2009). Monazite growth at ca. 360 and 305 Ma (Buchanan et al., 2016, 2017), combined with biotite and muscovite ⁴⁰Ar/³⁹Ar cooling ages from ca. 300–267 Ma, show limited thermal resetting from major Alleghanian high-grade events to the south (e.g., Hepburn et al., 1987; Walsh et al., 2007; Wintsch et al., 2007, 2012; Attenoukon, 2009; Reynolds, 2012; Kuiper et al., 2013). Cooling of the Putnam-Nashoba terrane through greenschist-facies temperatures occurred from ca. 330–260 Ma over the same time interval that the adjacent Avalonian rocks structurally below and to the south were loaded and heated to peak metamorphic conditions as indicated by U-Pb ages from monazite and sphene from ca. 280–270 Ma (Wintsch and Aleinikoff, 1987; Wintsch et al., 1992; Attenoukon, 2009). Subsequent post-peak metamorphic, late Paleozoic to Mesozoic normal motion along the Clinton-Newbury and Lake Char–Bloody Bluff and related faults, led to the current configuration of fault-bounded lithotectonic terranes (e.g., Goldstein, 1982a, 1982b, 1994; Goldsmith, 1991c; Attenoukon, 2009; Wintsch et al., 2012; Walsh and Merschat, 2015).

Zircon, monazite, and titanite were dated by U-Pb geochronology. Ten samples were collected to constrain the primary ages of the protoliths, the age of crosscutting plutonic rocks, the age of geologic structures observed in outcrops, and the timing of metamorphic mineral growth. The dated minerals may contain multiple age domains, necessitating use of high spatial resolution spot analysis (i.e, Kröner et al., 2014; Schaltegger et al., 2015). For this purpose, we dated our samples using the U.S. Geological Survey (USGS)/Stanford sensitive high-resolution ion microprobe-reverse geometry (SHRIMP-RG) at Stanford University, Palo Alto, California, USA. Location of spot analysis on mineral grains was guided by reflected and transmitted light imaging on a petrographic microscope and cathodoluminescence (CL) or backscattered electron (BSE) imaging on a scanning electron microscope. SHRIMP geochronology results from this study are summarized in Table 1, and data (including sample locations) are provided in Tables 2 and 3. Details of the sampled outcrop locations are included in Tables 2 and 3. See the Appendix for a complete description of the geochronology procedures.

A roadcut in Plainfield, Connecticut, USA, exposes layered amphibolite gneisses of the Quinebaug Formation. Boudins of amphibolite gneiss both contain, and are wrapped by, foliated tonalitic leucosome (Fig. 5). We sampled two rocks: one from the inside (mafic core) and one from the outside (leucosome rim) of a single mafic boudin (Fig. 5). The sampled rocks show coarse-grained leucrocratic partial melts surrounding medium-grained dioritic to amphibolitic gneiss in boudins. We collected a sample from the core of the boudin that might yield igneous or metamorphic zircon, and a sample of the exterior leucocratic portion that would record the time of subsequent partial melting.

Zircon grains from the mafic boudin interior (sample 629A2) generally are colorless, prismatic with length-to-width ratios (l/w) of ~3–6, and display fine concentric oscillatory zoning (Fig. 6). Most SHRIMP analyses were located on near-vertical oscillatory zones (when viewed on the horizontal mount) toward the outer portions of the grains, to avoid possible inherited material. A few grains show (in CL) light-gray rims that in places invade the core region, and a few grains have black outermost rims. Oscillatory-zoned cores contain between ~100 and 500 ppm U and have Th/U of 0.4–0.8, characteristic of igneous origin (Hoskin and Schaltegger, 2003). Light-gray rims have low U contents (~9–190 ppm), whereas black outermost rims have high U contents (~1100–1500 ppm). Both types of rims have low Th/U of 0.02–0.13, typical of metamorphic origin (Table 2; Fig. 7A) (Hoskin and Schaltegger, 2003).

Although we interpret the oscillatory-zoned areas as igneous in origin, U-Pb data suggest that these zircon grains formed during at least three melt-forming episodes. We are unable to distinguish these populations on an a priori basis of external morphology or zoning pattern; instead, we use the age data for determining different populations. The four oldest analyses of cores result in a weighted average of ²⁰⁶Pb/²³⁸U ages of 573 ± 7 Ma, whereas 17 (of 18) analyses of cores yielded a weighted average of ²⁰⁶Pb/²³⁸U ages of 540 ± 6 Ma, and five other grains yielded ²⁰⁶Pb/²³⁸U ages between ca. 500 and 460 Ma (Fig. 7B). In addition, two light-gray (in CL) rims are ca. 450 Ma, and two outermost black rims are 337 ± 3 and 315 ± 3 Ma, respectively (Fig. 7B).

Zircon grains from the tonalitic leucosome boudin exterior (sample 629A1) adjacent to, and outside of, the sampled mafic boudin interior either are pale to medium brown, prismatic (l/w mostly of 4–6) with concentric oscillatory zoning, or less elongate or equant (l/w of 3–4 or 1–2, respectively), and display sector or irregular zoning (Fig. 8). Cores in elongate grains contain ~100–300 ppm U, whereas rims have ~300–1000 ppm U (one dark unzoned grain has ~2450 ppm U). Values for Th/U in cores are mostly between ~0.5 and 1.0, whereas Th/U values in rims and metamorphic grains are lower (~0.1–0.3) (Table 2; Fig. 7C).

Thirteen oscillatory-zoned prismatic grains yield somewhat scattered U-Pb data and a weighted average of ²⁰⁶Pb/²³⁸U ages of 539 ± 7 Ma (Fig. 7D); one grain is older at ca. 591 Ma. Eight analyses from five grains (Fig. 8) yield a weighted average of ²⁰⁶Pb/²³⁸U ages of 325 ± 2 Ma (Fig. 7D).

Titanite was obtained from the boudin interior (sample 629A2). These grains are pale brown and anhedral. In BSE images, most grains have dark-gray cores, intermediate gray mantles, and light-gray rims (Fig. 9). In BSE, the grayscale zoning is related to U contents, i.e., dark-gray cores contain ~8–180 ppm U (mostly between 10 and 25 ppm), mantles contain ~175–700 ppm U, and light-gray rims contain ~100–1350 ppm U (Fig. 7E). Cores have the highest proportion of common Pb (up to nearly 40%, Table 2). Because of high common Pb contents of unknown Pb isotopic composition, a Tera-Wasserburg plot (Fig. 7F) is constructed using U-Pb isotopic data that are uncorrected for common Pb (Table 2). The highest U analyses (with the highest precision, as shown by the smallest error ellipses) are from the light-gray rims; these are the least discordant and least scattered data. A regression through six analyses (five rims and one core) yields a lower intercept age of 301 ± 5 Ma (red error ellipses in Fig. 7F). Isotopic data for most cores and mantles are less precise and scattered; no ages were calculated for these analyses.

In summary, based on SHRIMP U-Pb age data from prismatic oscillatory zircon from the interior of the boudin and from leucosome surrounding the boudin, we conclude that mafic rocks of the Quinebaug Formation were partially melted at ca. 540 Ma. Zircon dated at ca. 591 Ma from the leucosome and 573 Ma from within the mafic boudin could be either: (1) original igneous zircon from the protolith mafic rock, or (2) inherited xenocrystic zircon included within the protolith mafic rock. Because mafic rocks rarely contain primary zircon, it is unlikely that the 591–573 Ma zircon crystallized in the igneous protolith. However, we have no evidence to distinguish between these two possible origins. Regardless, the age of these older grains serves as a constraint on the age of the protolith. We conclude that the protolith mafic igneous rock crystallized at 573 Ma or younger, but before the high-grade metamorphism that culminated in partial melting at ca. 540 Ma. Additional metamorphism and deformation occurred at ca. 325 Ma (age of zircon rims) and continued (probably episodically) until ca. 300 Ma (age of titanite rims).

The sampled rock was collected from an outcrop on the south side of Black Hill in the town of Plainfield, Connecticut (Plainfield 7.5′ quadrangle; Dixon, 1965). Dixon (1964, 1965) interpreted the calc-silicate granofels Black Hill Member of the Quinebaug Formation as a stratigraphic unit within the middle of the formation. The sampled rock (616H1, Fig. 10F) is a light-gray to medium-gray micaceous quartz granofels with thin dark-gray horizons of muscovite-biotite schist cut by quartz veins.

Zircon extracted from sample 616H1 was subdivided into two morphologic populations: (1) a small percentage of the zircon grains are colorless, euhedral, elongate, and show fine concentric, oscillatory zoning (Fig. 10A), and (2) most grains are rounded, frosted, and display a wide variety of colors and CL zoning patterns (Fig. 10B). Sixty euhedral grains (type 1) and 54 rounded grains (type 2) were dated; all analyses are concordant (Table 3; Figs. 10C and 10D) and were included in calculation of relative probability age distribution curves (Fig. 10E).

Relative probability plots for both populations (Fig. 10E) show similar age populations including Ediacaran–Cambrian (625–525 Ma) and early Paleozoic (450–400 Ma) ages. Also noteworthy is the paucity of Grenville-age (1.35–0.99 Ga) grains. The youngest sub-population of Type 1 (euhedral) grains is ca. 392 Ma (n = 3), whereas the youngest sub-population of Type 2 (rounded) grains is ca. 371 Ma (n = 7) (Fig. 10E). We considered the possibility that the youngest grains might have originated in felsic dikelets within the sampled outcrop, and as such, would postdate the time of deposition. However, we dismiss this possibility because (1) the only crosscutting features are quartz veins with trace amounts of feldspar (Table 3), and no young crosscutting igneous material was observed in outcrop; and (2) the rock is a low-medium–grade phyllite and granofels. Thus, we conclude that the youngest grains are detrital and provide a weighted mean age constraint for the time of deposition of the sedimentary protolith. One metamorphic rim from a euhedral grain is 339 ± 3 Ma (Table 3).

Because of the occurrence of a significant number of detrital zircon grains younger than 500 Ma, the granofels cannot be a member of the Ediacaran–Cambrian Quinebaug Formation. Instead, ages of the youngest detrital zircon sub-populations from both morphologic groups indicate a Devonian age of deposition after ca. 371 Ma for the granofels at Black Hill. The age of ca. 340 Ma from one metamorphic rim constrains the minimum age of deposition. Thus, the depositional age ranges from Late Devonian to Middle Mississippian.

Mappable feldspathic granofels occurs in the western part of the Marlboro Formation, structurally above the Grafton Gneiss, in the Grafton (Walsh et al., 2011a) and Worcester South 7.5′ quadrangles (Walsh and Merschat, 2015) (Fig. 2). The sample selected for analysis is a gray biotite-quartz-feldspar granofels that locally occurs as banded gneiss interlayered with amphibolite.

Zircon grains from sample GR-098A are colorless, prismatic (l/w of 3–5), and display fine concentric oscillatory zoning. Many grains have narrow, unzoned light-gray inner rims, and dark-gray to black outer rims (Fig. 11). Cores contain ~50–500 ppm U, whereas rims have higher U contents of ~640–1380 ppm. Cores have Th/U of ~0.6–0.8, typical of igneous zircon; rims have Th/U of 0.04–0.11, typical of some metamorphic zircon (Table 2; Fig. 12A). SHRIMP analyses of 13 grains result in a weighted average of ²⁰⁶Pb/²³⁸U ages of 500.6 ± 4.2 Ma (Fig. 12B). This age is interpreted as the time of igneous crystallization of the protolith; however, the rock may have a detrital component, and thus the age may represent the time of deposition of a proximal source rock because the zircons are not rounded, or deposition of a volcanic protolith. One grain with irregular zoning yielded an age of ca. 460 Ma. Five grains have older ²⁰⁶Pb/²³⁸U ages of ca. 532–522 Ma, and probably are xenocrystic in origin. Rims and unzoned grains, with Th/U of 0.04–0.11 have ages in the range of 410–380 Ma, interpreted as the time(s) of metamorphic growth. The clastic nature of the rock and the positive Zr-Hf anomalies (Fig. 19G) suggest that the ca. 500 Ma age is the depositional age of the rock. The low mean square of weighted deviates (MSWD) age suggests a single population of igneous zircon with some older inherited grains. The prismatic zircon morphology supports the interpretation that the age may closely represent the depositional age of the rock derived from a proximal volcanic source. An epiclastic origin is most likely.

To bracket the timing of deposition and subsequent complex metamorphic and igneous events in the adjacent and overlying formations, we sampled three igneous rocks (samples WO-125A, B, and D) and one metasedimentary rock (sample WO-125C) in Oxford, Massachusetts, from a 220-m-long roadcut in the Nashoba Formation on Interstate 395 (I-395) (Figs. 13–15). In addition, we sampled a paragneiss from the Tatnic Hill Formation in northeastern Connecticut (Fig. 2), for comparison with detrital zircon U-Pb ages in the Nashoba Formation.

Numerous photographs and structural measurements from the I-395 roadcut accompany the geologic map of the area (Walsh and Merschat, 2015). The sampled rocks on I-395 include tonalite (WO-125A), granitic pegmatite (WO-125B), paragneiss (WO-125C), and granodiorite (WO-125D) (Figs. 13–15). The roadcut is located ~1 km northwest of the Assabet River fault (Fig. 1), ~0.5 km structurally above the underlying Marlboro Formation.

The roadcut on I-395 consists largely of migmatitic gray schist to paragneiss typical of the Nashoba Formation in Massachusetts. The rock is medium- to coarse-grained, well-foliated, migmatitic, biotite-plagioclase-quartz-muscovite gneiss to muscovite-K feldspar-plagioclase-quartz-biotite schist. The migmatitic rock contains pink and white to white, white-weathering, coarse-grained, unfoliated, steeply dipping, tabular, muscovite granitic pegmatite dikes and mylonitic to weakly foliated syn-tectonic granitic pegmatite sills. The leucocratic rocks mainly occur parallel to the first and second generation foliations in the outcrop (S₁ and S₂ foliations, respectively) (Fig. 14D) consistent with formation during upper amphibolite–facies metamorphism (Walsh et al., 2011a). Sillimanite + alkali feldspar assemblages are pre- to syn-tectonic relative to the second-generation (D₂) deformation. Mylonitic foliation subparallel to the S₂ foliation in the Nashoba Formation locally postdates K-feldspar porphyroblast growth, and many feldspars exhibit grain-size reduction in mylonite zones (Walsh et al., 2011a). In Massachusetts, Abu-Moustafa and Skehan (1976) estimated peak metamorphic conditions of 650°C to 625°C at 6 kilobars (kb), and Buchanan et al. (2015) estimated 600°C to 700°C at 4.5–6.0 kb. In Connecticut, conditions reached 700°C at 5.7 kb (Moecher and Wintsch, 1994).

Most of the large conspicuous granitoid layers at the I-395 roadcut (Fig. 13 and 15) postdate the S₁ foliation, although some are sub-parallel to S₁. Most granitic pegmatites are either folded by the D₂ deformation or occur within the plane of the S₂ foliation (Fig. 14D). The largest zone of pegmatite occurs in the low strain zone in the hinge of the large south-verging F₂ antiform (Fig. 13), implying the magma was mobile and actively migrating during D₂ deformation. Late-D₂ to post-D₂ mylonitic shear zones cut the largest granitic pegmatites. D₃ folding and cleavage formation is associated with late crosscutting shear bands, faults, and minor pegmatite dikes (Fig. 13). Thin tabular pegmatite dikes locally crosscut the dominant foliation. The distribution of the granitic pegmatites and the D₂ to D₃ shear zones are consistent with formation during episodic melt weakening (for example, Handy et al., 2001; Weinberg and Mark, 2008). Based on a detailed structural analysis of this outcrop, Goldstein (1982a, 1982b) reported top-down-to-the-north motion during the latest D₃ shearing and attributed it to regional motion on the Bloody Bluff–Lake Char fault and development of the Douglas Woods anticline. Small sigmoidally deformed D₃ pegmatite dikes observed at the roadcut during our study agree with D₃ normal motion reported by Goldstein (1982a, 1982b).

Biotite granodiorite was collected from the south-central part of the I-395 exposure (Fig. 13) by using a ladder (Fig. 15). The sampled layer contains a rare F₁ fold observed at this outcrop, and for this reason, it was specifically targeted to constrain the timing of early deformation in the Nashoba Formation (Fig. 15).

Zircon from the granodiorite (sample WO-125D) is subhedral to euhedral, colorless to pale tan, and has l/w of 2–6. In CL, many grains have dark-gray to black oscillatory zoned cores, and light-gray, weakly zoned rims (Fig. 16A). Cores contain ~230–1010 ppm U, whereas light rims and metamorphic grains contain ~130–320 ppm U. Cores have Th/U mostly in the range of 0.3–0.6, whereas rims have Th/U mostly of 0.07–0.2 (Table 2; Fig. 16B inset). The combination of CL zoning patterns and Th/U suggest that the cores are of igneous origin and that the rims formed due to metamorphism. SHRIMP U-Pb analysis of 15 cores yield a weighted average ²⁰⁶Pb/²³⁸U age of 409 ± 3 Ma (Fig. 16B). This age is interpreted as the time of igneous crystallization of the granodiorite. One core is slightly older at 422 ± 3 Ma (Table 2). Rims yielded a weighted average ²⁰⁶Pb/²³⁸U age of 392 ± 5 Ma (Fig. 16B). However, the age data are scattered (MSWD = 4.1), suggesting the possibility of multiple times of sub-solidus growth. Using the Unmix routine of Isoplot for two age components, these ages are 400 ± 4 and 385 ± 3 Ma. One rim is slightly younger at 374 ± 4 Ma.

Gray equigranular tonalitic granofels to gneiss layers locally occur within the Nashoba Formation. The I-395 roadcut exposes two areas of tonalite, which occur as layers parallel to the dominant foliation (Figs. 13 and 14A). When the tonalite was first mapped and sampled, it was not clear from field observations whether the rock had an igneous or sedimentary origin.

Zircon from the tonalite (sample WO-125A) is euhedral, colorless to pale brown, and has l/w of 2–5. In CL, cores display fine, concentric oscillatory zoning; rims are light gray and unzoned (Fig. 16C). Although both cores and rims have similar U contents (~150–350 ppm U), they have very different Th/U (Table 2; Fig. 16D inset). Cores have unusually high Th (~180–625 ppm) resulting in Th/U of 1.0–2.2, much higher than typical Th/U values of 0.3–0.8 for igneous zircon. Rims have Th/U of 0.01–0.18, characteristic of metamorphic sub-solidus formation. SHRIMP U-Pb analysis of 13 cores yield a weighted average of ²⁰⁶Pb/²³⁸U ages of 322 ± 3 Ma (Fig. 16D). This age is interpreted as the time of igneous crystallization of the tonalite. Seven analyses of rims yield an age of 324 ± 6 Ma and indicate that the tonalite was metamorphosed shortly after crystallization, implying syn-tectonic emplacement of the tonalite magma under high-temperature conditions.

Granitic pegmatite was collected from the north end of the I-395 exposure (Figs. 13 and 14B). The sampled granitic pegmatite occurs as a 10–20-cm-thick layer parallel to the dominant foliation (S₂) within the gray migmatitic paragneiss (Figs. 13 and 14B) and was collected to constrain when the rock experienced partial melting and emplacement of granitic leucosome.

Zircon from the pegmatite (sample WO-125B) is subhedral to euhedral, dark brown, and has l/w of 1–3. In CL, the grains display either broad, concentric zoning or are unzoned (Fig. 16E). Ten (of 13) analyzed grains have high U contents (1170–2240 ppm U), two grains have very high U contents (4330 and 8140 ppm). All high-U grains have very low Th/U of 0.01–0.03 (Table 2). Only one dated grain has relatively low U of nearly 400 ppm and relatively high Th/U of 0.22. Excluding the very high-U grain, 12 SHRIMP analyses yield a weighted average of ²⁰⁶Pb/²³⁸U ages of 331 ± 10 Ma (Fig. 16F). The grain with very high U (8144 ppm) has a ²⁰⁶Pb/²³⁸U age of 379 ± 16 Ma (Table 2). This old age probably is caused by the matrix effects characteristic of very high-U zircon (Williams and Hergt, 2000; White and Ireland, 2012) and is regarded as geologically meaningless.

Two samples of paragneiss from the Nashoba and Tatnic Hill Formations were analyzed to constrain the times of deposition of these units. The Nashoba Formation paragneiss (sample WO-125C) was collected from the very top edge of the roadcut (Fig. 13) where thick crosscutting leucocratic layers were easily identified and purposefully avoided during collection of the paragneiss (Fig. 14C). Very thin, mm-scale leucosomes occur throughout the sampled rock (Fig. 14C) and are unavoidable at the outcrop. For comparison with the data from the Nashoba paragneiss, a sample of similar paragneiss (sample 040814-1) was collected from the Tatnic Hill Formation in northeastern Connecticut (Fig. 2). The rock closely resembles the major rock type sampled in the Nashoba Formation to the north in Massachusetts.

Zircon grains from Nashoba sample WO-125C and Tatnic Hill sample 040814-1 have rounded or partially resorbed cores (Figs. 17A and 17C); oscillatory zoning in the cores is truncated, indicative of mechanical abrasion and detrital origin. Grains from the Nashoba Formation are overgrown by narrow to broad, light-gray (in CL) rims, whereas grains from the Tatnic Hill Formation have light- to medium-gray rims (Figs. 17A and 17C). Monazite from WO-125C mostly has irregular patchwork-zoned or spotted cores, typical of metamorphic origin (grains 5 and 8, Fig. 17B). Concentric oscillatory zoning is preserved in one dated monazite grain (grain 9, Fig. 17B). A plot of nested Concordia U-Pb data for the Nashoba and Tatnic Hill detrital zircons shows that most of the grains are concordant or nearly concordant (Fig. 17D). U-Pb data for monazite result in two ages: 12 analyses of cores yield a weighted average of ²⁰⁶Pb/²³⁸U of 391 ± 4 Ma, whereas three analyses of rims are 374 ± 6 Ma (Fig. 17E).

Age distributions of detrital zircon from the Nashoba and Tatnic Hill Formations are similar (Figs. 17F and 17G). Both samples yield major peaks in the range of ca. 650–500 Ma and minor peaks in the range of ca. 1300–800 Ma. Younger ages of ca. 437 and 428 Ma and a major peak at ca. 390 Ma in both samples result from analysis of metamorphic zircon rims. The 390 Ma metamorphic zircon age agrees with the older monazite age of 391 ± 4 Ma. Cambrian peaks are likely related to sources in the adjacent and underlying Quinebaug-Marlboro belt, and the detrital zircon ages overlap with the ca. 540–500 Ma ages identified from igneous rocks in this study. Although several grains in each sample have Archean and Paleoproterozoic ²⁰⁷Pb/²⁰⁶Pb ages (Table 3), these ages are not shown on Figure 17F because the data mostly are >10% discordant. The ages for the relative probability curve (Fig. 17F) are limited to 1400 Ma to better visualize Neoproterozoic and Paleozoic age peaks, and Figure 17G shows all concordant ²⁰⁷Pb/²⁰⁶Pb ages and all ²⁰⁶Pb/²³⁸U ages <1.0 Ga. In summary, detrital and metamorphic ages bracket the time of deposition to be between ca. 486 and 437 Ma for the Nashoba sample and 502 and 428 Ma for the Tatnic Hill sample, or ca. 485–435 Ma overall (Fig. 17F).

In southeastern Connecticut, the Preston Gabbro intruded the Quinebaug Formation and experienced cataclastic deformation in the hanging wall of the Lake Char–Honey Hill fault (Fig. 2; Zartman and Naylor, 1984). Dixon and Goldsmith (1983) report that the gabbro core grades into diorite to quartz diorite in a border phase around the pluton, and that intrusion of the Preston Gabbro took place after the host Quinebaug Formation had experienced significant metamorphism and deformation. A trondhjemite phase, considered to be cogenetic with the gabbro, yielded a multi-grain TIMS U-Pb zircon age of 424 ± 5 Ma (Zartman and Naylor, 1984). The trondhjemite occurs as dikes (Sclar, 1958), and one of the dikes cross cuts the contact between the diorite and the Quinebaug Formation at the locality that yielded the 424 Ma age (Dixon and Goldsmith, 1983). The diorite crosscuts a foliation in the adjacent Quinebaug Formation demonstrating that some deformation in the adjacent Quinebaug Formation pre-dates emplacement of the pluton. We collected a sample of the dioritic phase of the pluton for comparison with the published trondhjemite age. The dioritic phase occurs as dikes 15–50 cm wide with sharp contacts cutting the host gabbro.

Zircon from the dioritic phase of the Preston Gabbro is euhedral, colorless to pale tan and has l/w of 1–4. In CL; all grains have cores with fine concentric oscillatory zoning; most grains have thin, discontinuous white overgrowths that were not dated (Fig. 18A). Most zircon grains contain between ~200 and 300 ppm U and have Th/U of ~0.5 and 1.1. Two analyses made on slightly darker (in CL) areas have higher U (~500–580 ppm) and Th/U (~1.5) (Table 2; Fig. 18B inset). SHRIMP analyses of 12 cores yield a Concordia age of 418 ± 3 Ma, interpreted as the time of crystallization of the diorite. This age is within uncertainty of the TIMS age of 424 ± 5 Ma from the trondhjemite (Zartman and Naylor, 1984).

Twenty-eight samples from the Putnam-Nashoba terrane were analyzed for whole-rock geochemistry. Samples for major- and trace-element analysis, plus isotope geochemistry on most rocks include the following: 16 amphibolite and three felsic granofels from the Quinebaug and Marlboro Formations, five amphibolites from the Tatnic Hill and Nashoba Formations, one sample of the Grafton Gneiss, and three dated granitoids from the Nashoba Formation. Table 4 presents major- and trace-element data, Table 5 presents sample locations, and Table 6 presents isotope data. The locations of the samples are shown on Figure 2, and the Supplemental Material¹; the latter includes an enlarged figure of the sample locations and a Google Earth file of the data in Table 5.

Disturbances to the bulk chemistry before or during metamorphism cannot be ruled out. Most of the geochemistry samples have values for loss on ignition (LOI) lower than ~2.7 wt% (Table 4), but several samples were likely pervasively altered based on their mineral assemblages, in particular, amphibolites with higher values of loss on ignition (LOI) and those rich in chlorite and albitic plagioclase (e.g., MA-10, MA-2, and CT-8). Because alteration processes are related to enhanced mobility of the major oxides (SiO₂, Fe₂O₃, MgO, CaO, Na₂O, and K₂O) and open-system behavior of the low-field-strength elements (LFSEs: Cs, Rb, Ba, and Sr) (Wood et al., 1976; Winchester et al., 1995), these oxides and elements are for the most part unreliable as indicators of the original magmatic abundances, and are only used for preliminary characterization of the felsic rocks (Fig. 19) and comparison with previously published data (Fig. 20). Most of the rare-earth elements (REEs) (particularly middle- and heavy-REE), the high-field strength elements (HFSEs: Hf, Nb, Ta, Ti, Y, and Zr), Th (e.g., Pearce, 1983; Gao et al., 2007), and some of the transition metals (Cr, Ni, Sc, and V) are thought to be relatively immobile except during the most intense stages of alteration (MacLean and Krandiotis, 1987; Knoper and Condie, 1988; Staudigel et al., 1996), and for this reason, we rely mainly on these elements. In this section, multi-element mantle-normalized and chondrite-normalized diagrams and abundance plots of relatively immobile elements are primarily used to characterize the samples. Also, we use the three radiogenic isotope systems to characterize the source and evolution of the amphibolitic rocks, highlighting the results for the Sm-Nd and U-Th-Pb systems.

Our limited data set for the felsic rocks (Table 4; Fig. 19) is meant to classify the chemistry of the three dated granitoids and three samples of granofels (one of which was dated) but is insufficient to adequately address the potentially complex origin of these rocks, and plots of alkalis may be impacted by alteration (Figs. 19A–19D). The geochemistry shows that the three dated rocks at the I-395 roadcut are ferroan, peraluminous, calc-alkaline granitoids with syn-subduction to postcollisional signatures (Fig. 19). For those samples with SiO₂ between 55% and 70%, a discrimination diagram for arc versus slab failure (Fig. 19H) shows only two samples plot in the fields defined by Whalen and Hildebrand (2019). One granofels sample (GR-098A) plots in the arc field, and the tonalite (WO-125A) plots in the slab failure field (Fig. 19H). The remaining samples have SiO₂ values between 71.0% and 74.5%. Chemically, the dated granofels sample (GR-098A) is a calc-alkaline dacite, and a second nearby sample of the granofels (GR-098B) is rhyolitic (Figs. 19 and 20). All three granofels samples show negative Nb and Ta anomalies typical of volcanic arc signatures (Fig. 19G). Two granofels plot in the arc field (Fig. 19H), and all three have slight positive Zr-Hf anomalies (Fig. 19G), the latter of which may also occur in clastic rocks due to concentrated detrital zircons.

Our effort to classify and characterize the geochemistry of the protoliths focused mostly on the amphibolites (21 of 28 samples, Fig. 20A). The sampled exposures all contain variable amounts of migmatitic rock with different degrees of partial melt, leucosome, melanosome, granitoid sheets, felsic granofels, and paragneiss. In places, amphibolite occurs only as boudins in deformed and migmatitic rock. Amphibolitic rocks with visible amounts of partial melt or quartz-feldspar layers were largely avoided in this study. For comparison purposes, an IUGS plot of total alkalis versus silica (Na₂O + K₂O versus SiO₂) from Le Bas et al. (1986) shows that most of the amphibolites from our study and two previous studies (Wintsch et al., 1990; Kay et al., 2017) plot as basaltic and overlap into adjacent fields (Fig. 20A). Granofels and quartz-feldspar gneiss, including the Grafton Gneiss and the Fish Brook Gneiss, plot as intermediate to acidic rocks (Fig. 20A).

The largely basaltic rocks in the Quinebaug-Marlboro belt and from a small sampling of rocks from the overlying Nashoba and Tatnic Hill Formations form two overlapping compositional clusters on the basis of major- and trace-element contents. Values of Mg# [Mg# = 100 × MgO/(MgO + FeO)] illustrate the multiplicity of the compositions (Mg # ~25–57) (Table 4; Fig. 20B). If Mg and Fe contents were not significantly disturbed as a result of alteration and metamorphism, the wide range and generally low Mg# values can be taken as a record of magma fractionation in the amphibolites, but the rocks experienced high-grade metamorphism, so the trace-element diagrams are more reliable. Low values of Mg# (<50), as well as relatively low contents of Ni (most <200 ppm), Co (<60 ppm), and Cr (most <300 ppm) are consistent with mafic magmas that have undergone olivine-pyroxene fractionation (Table 4).

Figure 20C shows the Nb/Y-Zr/Ti diagram of Pearce (1996) for rock classification. Figure 20D is a classification figure that shows only our amphibolites divided according to the presence of a Nb-Ta anomaly compared to Th and Ce, a key feature of volcanic arc basalt. Most amphibolites are basic to intermediate (Zr/Ti ~0.01). One suite of amphibolites has Nb/Y ~0.1–0.2 and plots in the sub-alkalic basalt field and adjacent to or on the boundary with basaltic andesite and andesite (Fig. 20C); the second has Nb/Y ~0.5–2.8 and mostly plots as alkali basalt or straddles the boundary between basalt and alkali basalt (Fig. 20C). Most classification diagrams, including those using various combinations of Zr, TiO₂, and P₂O₅, indicate the dominance of tholeiitic basalts compared to alkali basalts for the Quinebaug and Marlboro Formations. The classification plot in Figure 20D shows that the occurrence of a negative Nb-Ta anomaly cannot be used to distinguish basaltic amphibolites from the different units.

Mantle-normalized trace-element abundance patterns (Fig. 21) (Sun and McDonough, 1989) of the basaltic amphibolites show moderate enrichment in Th (Quinebaug Formation) and K (Quinebaug, Nashoba, and Marlboro Formations), prominent spikes for Pb in all units, and the diagnostic negative Nb-Ta anomaly typical of volcanic arc magmas. A few samples, however, show relative enrichments in Nb and Ta compared to Th and Ce (normally indicative of enriched mantle sources; Pearce, 1996) (Fig. 21). Samples displaying negative Nb-Ta anomalies are also compositionally the most diverse. The variety of compositions, together with the prominent Pb spikes and development of negative Nb-Ta anomalies, could all reflect contamination by crustal rocks. Many of these samples also display moderate depletions of P, Zr, and Ti. Contents of Sr, Eu, and Zr show minor anomalies. A key observation is that the basaltic amphibolites are characterized by significant trace-element compositional range, even within single formations (e.g., Quinebaug and Nashoba) (Figs. 21A–21E).

Figures 22A–22E show rare-earth element concentrations for the basaltic amphibolites normalized to chondritic values. Both Figures 21 and 22 show the basaltic rocks plotted according to host formation (Figs. 21A–21D) and according to the occurrence of Nb-Ta anomalies. The rare-earth element patterns of the amphibolites plot between characteristic compositions of N-MORB and OIB. Few samples display a positive or a negative Eu anomaly. For the most part, the REE patterns cannot be used to distinguish among the different units. Moreover, the basaltic rocks do not differ in their chondrite-normalized patterns according to the presence of the Nb-Ta anomaly (Fig. 22E). The data show an overlap with basalts from the Ellsworth terrane (Schulz et al., 2008) (Fig. 22E), and samples from the Castine Volcanics (their TB-2) show the most similarity. Our data do show more elevated light rare earth elements, and a slightly negative slope, probably as a function of a greater crustal component compared to the Ellsworth terrane basalts. La/Ce depletion (concave down pattern) seen in Ellsworth terrane TB-1 samples is not present in our data (Fig. 22E). A comparison with data from the modern Mariana region shows that the REE data in this study are similar to the arc-rifting zone where the Mariana arc and the Mariana backarc basin converge (Pearce et al., 2005) suggesting a link to possible incipient backarc basin formation in our data.

Diagrams involving Nb/Yb, Th/Yb, and Ti/Yb (Pearce, 1996) and La-Y-Nb (Figs. 23A–23D) (Cabanis and Lecolle, 1989) are also effective in showing a negative Nb-Ta anomaly in basaltic rocks (Fig. 23A and 23D). The plot of Nb/Yb versus Th/Yb shows that the basaltic rocks have a wide compositional range and extend from typical mantle compositions (plotting in the MORB-OIB array) to the basalts formed at convergent margins or may have compositions consistent with magma-crust interaction (Fig. 23A). Examination of basaltic rocks according to the presence of the Nb-Ta anomaly (Fig. 23B) shows important differences: a group of rocks (most amphibolites from the Marlboro) that have no anomaly plot between N-MORB (normal mid-ocean ridge basalt) and E-MORB to OIB (enriched mantle). Basalts without negative Nb-Ta anomalies plot in the field of continental tholeiites. Another suite displays a negative Nb-Ta (most are from the Quinebaug) depart from the mantle array toward higher Th/Yb ratios, a feature characteristic of rocks associated with crust interaction.

The La-Y-Nb diagram (Figs. 23C and 23D) shows that the amphibolites from the Quinebaug and Marlboro Formations cannot be geochemically distinguished. The compositions include a group displaying a negative Nb-Ta anomaly and plot mostly as calc-alkali and continental basalts (Quinebaug Formation), and a few samples plot as backarc basin basalts (Quinebaug and Tatnic Hill Formations). A second compositional suite, lacking a negative Nb-Ta anomaly, overlaps the fields of continental basalt, plots near the boundary with backarc basalts and on the field boundaries for E-MORB and alkaline intercontinental rifts (Figs. 23C and 23D) (Marlboro, Quinebaug, and Nashoba Formations). The spatial overlap of the basaltic rocks (and overall compositional diversity) is also reflected in the isotopic data.

The Nd-Pb-Sr isotopic compositions of the basaltic rocks are listed in Table 6. Whole-rock Pb isotopic analyses of metabasaltic amphibolite yielded a wide range of compositions (Table 6). Present-day ratios of ²⁰⁶Pb/²⁰⁴Pb vary from 18.30 to 19.50, ²⁰⁷Pb/²⁰⁴Pb from 15.56 to 15.70, and ²⁰⁸Pb/²⁰⁴Pb from 37.90 to 39.60. Age-corrected Pb isotopic compositions of metabasalts were estimated using present-day U, Th, and Pb contents (Fig. 24). In the ²⁰⁶Pb/²⁰⁴Pb versus ²⁰⁷Pb/²⁰⁴Pb plot (Fig. 24A), no linear array is observed for the metabasalts as a group, and the metabasaltic groups with and without the Nb-Ta anomalies overlap (Fig. 24C). As was the case with the Nd (and Sr) isotopic compositions, no simple correlation between the Pb isotope ratios, their host formation, or their major- and trace-element grouping (with or without Nb-Ta anomaly) is evident. Most metabasalts plot below the model curve for average crust (μ = 9.74, Stacey and Kramers, 1975) (Figs. 24A–24D), which is nearly equivalent to the orogenic curve from the plumbotectonics model (Doe and Zartman, 1979, 1982). The metabasalts plot as an overlapping cluster shifted to higher values of ²⁰⁷Pb/²⁰⁴Pb relative to the Northern Hemisphere Reference Line, which defines the 1.77 Ga trend for MORB (Hart, 1984) (Figs. 24A–24D). Most of the data points plot to the right of the 4.53 Ga Geochron (Fig. 24A). The amphibolites plot mostly below the composition of average crust (evolution curve of Stacey and Kramers, 1975), and a few samples plot above the crustal evolution curve, at higher values of ²⁰⁷Pb/²⁰⁴Pb, suggesting a contribution from more crustal sources. Values of ²⁰⁸Pb/²⁰⁴Pb are not enriched relative to a given value of ²⁰⁶Pb/²⁰⁴Pb (Figs. 24C and 24D), consistent with the lack of influence from older lower continental crust that was depleted in U relative to Th. Figures 24A–24D also show that the Pb isotope values of the metabasaltic rocks overlap the field of the Ellsworth Schist in Maine characterized by Schulz et al. (2008) and thus provide a good match with the only other complete data set of Pb isotopes for these rocks.

The ε_Nd values for the metabasaltic rocks overlap and have positive values from ~0.4–7.3 (Fig. 25), and each of the geologic units contains basalts with a wide diversity of isotopic compositions. The ε_Nd signatures of the metabasalts with the most strongly positive values have isotopic compositions typical of lower Paleozoic mantle (Lower to Middle Ordovician) associated with ocean floor basalts (N-MORB-type mantle, ε_Nd ~9 and OIB, ε_Nd ~3–6 at 460 Ma; Swinden et al., 1997; Hofmann, 2007).

Figure 25 also shows that the ε_Nd values of the metabasaltic rocks overlap the field of Avalonian (mafic and felsic rocks) characterized by Murphy et al. (2009) and are significantly more juvenile (or mantle-like) than Ganderia (mafic and felsic rocks) as determined by Hibbard et al. (2007a). Comparison with other ε_Nd values highlights the overlap between Ganderian and Avalonian signatures, thus making it difficult to assign terrane affinity on these values alone (Fig. 25A). The wide spread in ε_Nd values of the Quinebaug-Marlboro belt amphibolites defines a broad trend that may owe its origin to magma mixing or source mixing of juvenile basalt (at ca. 0.5 Ga) and basement rocks with considerably longer residence in the crust (ca. 1.5–1.3 Ga). The range of ε_Nd data of the amphibolites suggests that the crustal component is Avalonian, although dispersion of ε_Nd toward lower values does not allow for such unique identification. Overlap with Ganderian mafic and felsic rocks (Hibbard et al., 2007a) suggests that crustal components with lower ε_Nd values and older T_DM ages (1.5 Ga) could have been involved as components in addition to the predominant Avalonian source.

Initial ⁸⁷Sr/⁸⁶Sr ratios of metabasaltic rocks range from 0.70334 to 0.70714 (Table 6). The isotopic ratios are scattered, and the wide range of compositions, particularly for samples that have ⁸⁷Sr/⁸⁶Sr >0.705, likely reflects redistribution of Rb and Sr during alteration, metamorphism, and/or from a contribution of radiogenic Sr from an external source (crustal contamination). The considerable scatter in the ⁸⁷Sr/⁸⁶Sr isotopic compositions of the metabasaltic rocks is unlikely to reflect uncertainties in the ages of the samples.

Some samples in this study show significant negative Nb-Ta anomalies, similar to arc-like rocks, and some show no negative Nb-Ta anomalies and are more similar to rocks from sources from enriched mantle such as E-MORB and OIB. In a few cases (two samples) the compositions point to backarc settings. An important result is that volcanic arc compositions dominate in the Quinebaug Formation (consistent with the negative Nb-Ta anomaly) and a more varied affinity from volcanic arc to some type of enriched mantle characterizes the Marlboro Formation. Only two samples indicate a backarc basin setting (Tatnic Hill and Quinebaug Formations), and they have negative Nb-Ta anomalies in contrast to the majority of samples that mostly scatter and plot in the fields of calc-alkaline basalts or continental tholeiites in the La-Y-Nb diagram (Fig. 23C).

Amphibolites from the Marlboro Formation in this study do not show a definitive affinity to backarc basin basalts in the La-Y-Nb diagram (Fig. 23C) (Kay et al., 2017). Several samples, however, plot near the boundary for backarc basins (Fig. 23D). Such backarc basin basalts are considered to be generated from melting sources ranging from normal to enriched-type MORB in areas underlain by a subducted slab (e.g., Saunders and Tarney, 1979). Our Marlboro Formation amphibolites also have widely varying V/Ti ratios (~20–50) that include a range from backarc basalts to alkaline basalts, which is another feature that differs from the mostly homogeneous group analyzed by Kay et al. (2017). Amphibolites from the Quinebaug Formation are even more dispersed than our Marlboro Formation samples and include V/Ti ratios (~10–100) typical of island-arc to alkaline basalts.

Marlboro Formation samples in our study show a clear affinity to E-MORB to OIB compositions (Nb/Yb-Th/Yb) and do not depart from the mantle array. This composition contrasts with samples by Kay et al. (2017); these samples deviate from the array toward rocks that may have been affected by deep crustal recycling. All rocks in our study show a Pb spike indicating an enriched mantle source, possibly as a result of metasomatic processes or one that interacted with continental crust. Moreover, the Marlboro and Quinebaug Formation amphibolites that differ from the mantle array (Fig. 23) have a negative Nb-Ta anomaly, which is once more consistent with the influence of crustal rocks. Despite the above differences, in general, our geochemical data for the Quinebaug, Marlboro, Tatnic Hill, and Nashoba Formations are similar to, and in general agreement with earlier interpretations on the Marlboro Formation (DiNitto et al., 1984; Hepburn et al., 1995; Oakes-Coyne et al., 1996; and Kay et al., 2011, 2017).

Rocks in the Ellsworth terrane in coastal Maine offer the best regional correlations, although some of our data suggest more of a possible backarc environment similar to Kay et al. (2017) than to a proto-oceanic rift environment as proposed for the rocks in Maine. The close match in age and chemistry supports the correlation with the Ellsworth terrane and suggests the southern New England rocks were not lost to subduction as interpreted by Kay et al. (2017).

Isotopically, the data for the Quinebaug and Marlboro Formations overlap, but the Marlboro is more homogeneous. Data show limited involvement of old crustal sources and an enriched mantle source. Isotopically, the rocks are more like Avalonia than Ganderia; however, the isotope data cannot uniquely distinguish Ganderian versus Avalonian basement either because the full range for these crustal blocks is not yet well established in New England or because the basement underlying these rocks is “isotopically indistinguishable” (Keppie et al., 2012, p. 319). Considerable isotopic overlap exists between rocks identified as Ganderia and Avalonia (Fig. 25), and additional work is needed to fully characterize these two peri-Gondwanan terranes in southeastern New England and Atlantic Canada (Kerr et al., 1995; Samson et al., 2000; Barr et al. 2003, 2010, 2019; Murphy et al., 2009; van Staal and Barr, 2012). Our data compiled with Avalonian data (Fig. 25) show that the distinction is not clear and question the interpretation that the two terranes have significantly different underlying crust (e.g., Aleinikoff et al., 2007; van Staal and Barr, 2012) and agree with the assessment that the terranes cannot be distinguished by isotopes alone (Murphy et al., 2009; Keppie et al., 2012). Perhaps a better measure is the observation that Avalonia generally contains very little Cambrian-aged arc-related components (van Staal and Barr, 2012).

Lithological similarities between formations in the Putnam-Nashoba terrane in Connecticut and Massachusetts historically led to local correlations of rocks in the two states (Hansen, 1956; Zen et al., 1983; Zartman and Naylor, 1984; Rodgers, 1985; Goldsmith, 1991c). Local lithologic correlations between the Quinebaug and Marlboro Formations (Zen et al., 1983; Zartman and Naylor, 1984; Rodgers, 1985; Goldsmith, 1991b) are supported by our geochronology. Available SHRIMP data and preliminary LA-ICP-MS data (Loan, 2011) do show, however, that the Quinebaug may be somewhat older than the Marlboro, and the entire belt may young to the north. Additional high-precision geochronology is needed to test this hypothesis. Primary igneous ages range from ca. 540–500 Ma with inheritance from ca. 591–573 Ma (samples 629A1, 629A2, and GR-098A, Table 1, Figs. 7 and 12). The belt is intruded by the Grafton Gneiss at ca. 515 Ma and the Fish Brook Gneiss at ca. 500 Ma. Inherited zircon grains in the Marlboro Formation granofels show an age cluster at ca. 525 Ma. A limited sampling of concordant zircon ages obtained by LA-ICP-MS shows a similar peak at ca. 525 Ma, with the youngest grains at ca. 470–460 Ma suggesting deposition into the Middle Ordovician (Loan, 2011).

The local long-standing lithologic correlation between the Tatnic Hill and the Nashoba Formations is corroborated by this study as both formations contain comparable detrital zircon signatures and depositional ages (Fig. 17). A comparison between the detrital zircon samples in this study with previously analyzed samples from the Tatnic Hill Formation (Wintsch et al., 2007) yields ambiguous results (Fig. 26). Our two new SHRIMP-analyzed samples do not have the same provenance as the Merrimack belt samples of Wintsch et al. (2007) (Fig. 26). Our new data suggest a depositional age of ca. 485–437 Ma. The oldest metamorphic rims occur at ca. 437 Ma, with a well-defined metamorphic peak related to dated metamorphic rims at ca. 400 Ma. Thus, we conclude that deposition of the Tatnic Hill and Nashoba Formations occurred between ca. 485 and 435 Ma. These findings support the conclusion that the Merrimack belt is younger than, and not directly correlative with, the Tatnic Hill and Nashoba Formations in the Putnam-Nashoba terrane. A sample with 40 concordant analyses from the Tatnic Hill Formation in Wintsch et al. (2007); sample T2 yielded only four Silurian zircons with ages ranging from 437 to 420 Ma, and more analyses are needed to test the validity of the sample.

The occurrence of the Quinebaug-Marlboro belt at the structural base of the Putnam-Nashoba terrane has long been interpreted as evidence that it is the stratigraphic base of the exposed rocks in the terrane (Bell and Alvord, 1976). Recognition of significant faults raised questions about the stratigraphic continuity of the Quinebaug-Marlboro belt to the overlying Tatnic Hill and Nashoba Formations (Bell and Alvord, 1976; Hepburn and Munn, 1984; Zartman and Naylor, 1984; Wintsch et al., 1990; Goldsmith, 1991b; Stroud et al., 2009; Loan, 2011). The boundary between the upper and lower parts of the Putnam-Nashoba terrane locally coincides with the Assabet River and Tatnic faults (Fig. 2). The Assabet River fault divides the Nashoba and Marlboro Formations for much of the contact (Hepburn and DiNitto, 1978; Zen et al., 1983; Goldsmith, 1991b; Hepburn et al., 1995) and was originally based on truncations of the Shawsheen and Fish Brook gneisses (Bell and Alvord, 1976; Goldsmith, 1991a). Skehan and Abu-Moustafa (1976), however, interpreted the contact between the Nashoba and Marlboro Formations as stratigraphic. Goldsmith (1991b), in support of Bell and Alvord (1976), concluded that the simplest interpretation was that the faults had disrupted the stratigraphy, but the sequence was still homoclinal. Loan (2011) concluded that no significant difference existed in the age populations of detrital zircons across the Assabet River fault, and thus it did not represent a terrane boundary. In Connecticut, the Tatnic fault locally coincides with the contact between the Tatnic Hill and Quinebaug Formations, but not along its entire length (Dixon, 1965; Dixon and Lundgren, 1968; Wintsch, 1979; Rodgers, 1985). Rodgers (1985) and Dixon (1965) showed it as a fault along much of its length, but a depositional contact in some places.

Our geochronological data support the conventional interpretation that the rocks of Quinebaug-Marlboro belt are older than the overlying Tatnic Hill and Nashoba formations. Detrital zircons in the Nashoba and Tatnic Hill Formations are consistent with provenance from the underlying Quinebaug-Marlboro belt. The similarities in the geochemistry of metabasaltic rocks of all four formations is permissive for the interpretation that the units were derived from the same mantle at different times (Murphy and Dostal, 2007), were at one time somewhat contiguous, and that the two internal faults, the Assabet River and Tatnic (Fig. 2), are not major terrane-bounding features. Our data agree with previously recognized trends for the similar geochemistry across the terrane (DiNitto et al., 1984; Hepburn et al., 1995, 2014; Oakes-Coyne et al., 1996; Kay et al., 2009, 2011, 2017; Kay, 2012). Considering that the dated Carboniferous granitoids in this study are near the Assabet River fault and are locally highly strained and mylonitic suggests that, rather than being significant early faults, these two features experienced late Paleozoic motion during Alleghanian deformation. This finding is at odds with the interpretation that monazite ages (obtained by electron microprobe) from within the Assabet River fault zone date the faulting between 393 and 390 and 384–363 Ma (Stroud et al., 2009). These authors also obtained younger monazite ages between 360 and 305 Ma, and they attributed these to fluid-related hydrothermal activity near faults. The local correlations for the granofels sample from Black Hill are less clear. The young detrital zircon ages and the comparative lack of upper amphibolite–facies metamorphism in the upper greenschist–facies rocks suggest that the sampled rock is either in fault contact with, or is unconformably on top of, the older parts of the Quinebaug Formation, and is not an internal member of the formation. It shares all the dominant north-south–trending structures found in the Quinebaug Formation, and the contact relationships, although poorly exposed, show the boundary experienced Paleozoic ductile deformation.

Several possible correlations are proposed. One alternative is that the sampled rock is equivalent to quartzite in the Scotland Schist (Fig. 2) (Rodgers, 1985), which, in eastern Connecticut, has been interpreted as: (1) conformably overlying the late Silurian to Early Devonian Hebron Formation in the adjacent Merrimack belt (Dixon and Lundgren, 1968; Rodgers, 1985; Wintsch et al., 2007), or (2) as an internal member of the Silurian sequence, and part of the Oakdale Formation (Pease, 1989). A second alternative correlation concerns the Harvard Conglomerate in Massachusetts (Fig. 2), which rests nonconformably above the Merrimack belt rocks just west of the Clinton-Newbury fault (Robinson and Goldsmith, 1991; Goldstein, 1994). Charnock (2015) reports a preliminary maximum depositional age by LA-ICP-MS on detrital zircon of ca. 414 Ma with Silurian and Proterozoic sources from the underlying Merrimack belt. Zircon data from sample 616H1 show a younger depositional age between ca. 371–340 Ma (Fig. 10E), and the sample also yielded detrital zircon ages that could conceivably come from dated igneous rocks in the nearby Merrimack belt—for example, the Canterbury Gneiss (414 ± 3 Ma; Wintsch et al., 2007), Ayer Granodiorite (407 ± 3 Ma; Walsh et al., 2013b), Chelmsford Granite (375 ± 3 Ma; Walsh et al., 2013b), and Eastford Gneiss (379 ± 4 Ma; Wintsch et al., 2007). A third correlation might be the Wamsutta Formation near the base of the Devonian to Pennsylvanian Narragansett Basin in Massachusetts and Rhode Island (Fig. 2) where rhyolite yielded a U-Pb zircon TIMS age of 373 ± 3 Ma and Thompson and Hermes (2003) established a link with rhyolite in the Fisset Brook Formation (373 ± 4 Ma, U-Pb zircon, Barr et al., 1995) in the Maritimes Basin in Canada. The issue of the complete distribution, age, and history of the apparently young rocks of the granofels at Black Hill remains unresolved at this time and awaits further study. It is clear, however, that the sampled rock in this study (616H1) is not part of the Quinebaug Formation (Fig. 4).

Lithological similarities between formations in the Putnam-Nashoba terrane historically led to regional correlations in the northern Appalachians. The rocks in Massachusetts have been correlated not only with those in Connecticut, but also with rocks in coastal New Hampshire, Maine, and Atlantic Canada (Fig. 1B) (Goldsmith, 1991b; Hepburn et al., 1995, 2014; Robinson et al., 1998; Kay et al., 2017). Current regional correlations are summarized in Table 7.

The best regional correlation, based on age, trace and Pb isotope geochemistry, is with the lower greenschist-facies metabasalts of the fault-bounded Ellsworth terrane in Maine (Figs. 1B, 22, 24, and 25; Schulz et al., 2008). Ages of ca. 509–502 Ma in the Ellsworth Schist and Castine Volcanics overlap with the ca. 500 Ma age from granofels in the Marlboro Formation. Trace-element geochemistry overlaps with Castine type TB-2 basalts (Fig. 22E) and Pb isotope data matches volcanic rocks in the Ellsworth Schist (Figs. 24A and 24C). Independently, Waldron et al. (2015) used a palinspastic reconstruction of Late Paleozoic faults in Maritime Canada and concluded that the Ellsworth and Putnam-Nashoba terranes were correlative.

Despite the lack of reliable geochronology, lithic similarities led workers in coastal New Hampshire and Maine to suggest a correlation between the metasedimentary rocks in the Nashoba Formation with the Rye Complex, and the rocks in the Liberty-Orrington belt (Fig. 1B) (Goldsmith, 1991b; Bothner and Hussey, 1999). Provisional detrital zircon data by LA-ICP-MS represent the first modern ages for the Rye Complex and suggest Cambrian to Ordovician deposition with the youngest zircons dated at ca. 530 Ma (Kane et al., 2014; Bothner et al., 2014). These workers report a principal peak at ca. 627 Ma with secondary peaks at 1224 and 1506 Ma. Two of the peaks occur in our Tatnic Hill Formation sample (ca. 620 and ca. 1220 Ma), but only one peak occurs in our Nashoba Formation sample (ca. 620 Ma) (Fig. 26). The detrital zircon sample from the Rye Complex lacks the detrital peak at ca. 540 Ma and the distinctive metamorphic peak found in our samples at 400 Ma.

Robinson et al. (1998) and Hibbard et al. (2006) correlated the Rye Complex with the Liberty-Orrington belt and St. Croix terrane in Maine (Fig. 1B), all of which have been interpreted as either part of the Ganderian passive margin (Hibbard et al., 2006, 2007a, 2007b), or as an extension of the Miramichi arc complex (Fig. 1B) (West et al., 2004; Schulz et al., 2008). According to West and Condit (2016) and Hussey et al. (2010), both the Falmouth-Brunswick sequence and Casco Bay Group and in the Liberty-Orrington belt in Maine consist of volcanic and sedimentary rocks from an Ordovician backarc tectonic environment deposited on Ganderian crust. Our detrital zircon data and the preliminary data of Loan (2011) suggest deposition of the largely metasedimentary Nashoba and Tatnic Hill Formations occurred between ca. 485 and 435 Ma, or as young as Silurian (Wintsch et al., 2007).

In southern New Brunswick (Fig. 1B), multiple Ganderian terranes occur in low-grade fault-bounded lithostratigraphic sequences with well-constrained Ediacaran to Ordovician ages, including, from northwest to southeast, the St. Croix, Annidale, New River, and Brookville terranes (Fig. 1B, i.e., Fyffe et al., 2011). In Maine and New Brunswick, the St. Croix terrane (Fig. 1B) lies structurally below and to the north and west of the Ellsworth terrane (Stewart et al., 1993, 1995; Schulz et al., 2008; Reusch et al., 2018). The New River terrane (Fig. 1B) lies south and east of the Ellsworth terrane across a series of sub-vertical faults (King and Barr, 2004; Schulz et al., 2008; Fyffe et al., 2011). The Annidale terrane lies north of the New River terrane, north of the Falls Brook–Taylor Brook fault, and along strike to the northeast of the St. Croix terrane (Fyffe et al., 2011). All these terranes contain well-defined sedimentary sequences and volcanic rocks with preserved volcanic textures that are entirely absent in the Putnam-Nashoba terrane due to overprinting of upper amphibolite–facies metamorphism and deformation. Nonetheless, permissible correlations can be made based on lithology, geochemistry, and geochronology (Table 7).

In the St. Croix terrane, a fragmental tuff in the Penobscot Formation overlaps in age with dated granofels in the Marlboro Formation at ca. 500 Ma. The existence of MORB-like metavolcanic rocks and felsic volcanic rocks dated at 503 ± 5 Ma (Osberg et al., 1995; Tucker et al., 2001) in the adjacent St. Croix terrane supports a Ganderian affinity for both the Ellsworth and St. Croix terranes, although faults prevented direct correlation between the two terranes (Schulz et al., 2008; Reusch et al., 2018).

Correlations with Newfoundland are speculative, and the rocks there are part of the Penobscot-Victoria arc system (Zagorevski et al., 2010; Fyffe et al., 2011). The arc-related rocks of the Victoria Lake Supergroup are considered to be the older part of the Penobscot arc in Newfoundland where they are interpreted to have formed the leading (paleowestern) edge of Ganderia (Rogers et al., 2006; van Staal et al., 2009; Zagorevski et al., 2010), instead of the trailing edge as seen in southern New England.

Our new ca. 540 Ma igneous ages from the Quinebaug Formation have not previously been recognized in the New England part of Ganderia. The closest potentially correlative rocks occur in southeast New Brunswick in the New River and Brookville terranes (Fig. 1B; Barr et al., 2003, 2010; Fyffe et al., 2011). In the latest Ediacaran to early Cambrian, between ca. 548 Ma and 528 Ma, the Ganderian margin of New Brunswick experienced subduction-related calc-alkaline arc volcanism (Fyffe et al., 2011). These ca. 550–530 Ma arc rocks apparently pre-date the development of the Penobscot arc-backarc system in the Maritime provinces of Canada (ca. 515–485 Ma; Johnson et al., 2012; van Staal and Barr, 2012; van Staal et al., 2012). Older ca. 555–550 Ma rocks in the New River terrane (of McLeod et al., 2003) are not yet known in the Quinebaug-Marlboro belt. In the Brookville terrane, potentially correlative arc rocks have largely calc-alkaline, continental arc geochemistry (Fyffe et al., 2011, and references therein). Correlations with remote island locations in the Gulf of Maine are also possible based on the work by Barr et al. (2010), and range in age from ca. 547–535 Ma, are correlated with the New River terrane, and show subduction-related signatures typical of the Ganderian margin.

Overall, lithological, geochemical, and geochronological similarities support a correlation between the Quinebaug-Marlboro belt with rocks exposed along strike across the Gulf of Maine to coastal Maine and New Brunswick (Fig. 1B). The data suggest that a consistent terrane affinity exists for these rocks as they share similar peri-Gondwanan characteristics of an ancient Penobscot or pre-Penobscot arc system, in places matching both Ganderian and Avalonian signatures. As depicted by Murphy et al. (2009), Ganderia may be the leading edge of Avalonia. Our detrital zircon samples do not show typical Gondwanan ages of ca. 2.2–2.0 Ga found in Avalonia yet they do show peaks from ca. 620–570 Ma which are common in both Ganderia and Avalonia (Pollock et al., 2009; Barr et al., 2019).

Future work will test whether these relict arc fragments can be separated by isotope geochemistry and age. Although it is possible that the overlying, largely metasedimentary rocks of the Tatnic Hill and Nashoba Formations may correlate and be coeval with rocks of the St. Croix terrane or Liberty-Orrington belt, the identification of primary volcanic rocks in the southern New England section remains a challenge for future work.

The entire Putnam-Nashoba terrane is an upper amphibolite–facies, fault bounded, downward thinning tectonic wedge sandwiched between greenschist to amphibolite-facies rocks of the Avalon terrane and Merrimack belt (Goldsmith, 1991b, 1991c; Zen et al., 1983). Figures 27 and 28 summarize the prolonged Paleozoic orogenic history of the Quinebaug-Marlboro belt and the larger Putnam-Nashoba terrane.

The rocks in the Quinebaug-Marlboro belt developed as part of a Ganderian arc complex in the very latest Ediacaran to early Cambrian (van Staal et al., 2012; Kay et al., 2017). This arc developed off the greater Gondwanan-Amazonia margin at ~50° south paleolatitude due to subduction of Iapetus Ocean crust (Fig. 28; Keppie et al., 2012; van Staal et al., 2012). The oldest reliable ca. 540 Ma ages presented in this study document partial melting of a mafic host rock to form tonalitic leucosome in the Quinebaug Formation (Fig. 27) and suggest that these rocks either pre-dated the Penobscot arc system (515–485 Ma, Johnson et al., 2012; van Staal and Barr, 2012; van Staal et al., 2012), or were an early phase of a complex arc system. This cryptic event is documented in New England for the first time, and its regional extent is currently unknown. Similar rocks occur in New Brunswick in the New River–Annidale, and Brookville terranes (Fig. 1B; Barr et al., 2003, 2010; Fyffe et al., 2011; Kay et al., 2017) where latest Ediacaran to early Cambrian rocks, between ca. 548 Ma and 528 Ma, are related to subduction and calc-alkaline arc magmas along the Ganderian margin (Fyffe et al., 2011).

In the middle to late Cambrian, arc to backarc volcanic rocks formed the core of the Quinebaug-Marlboro belt. The felsic Grafton Gneiss intruded mafic volcanic rocks at ca. 515 Ma, followed by deposition of the Marlboro Formation granofels at ca. 500 Ma and intrusion of the protolith of Fish Brook Gneiss at ca. 499 Ma (Hepburn et al., 1995). Circa 509–505 Ma, Ganderian crust rifted away from greater Gondwana-Amazonia during the opening of the Rheic Ocean, and Ganderia began its journey across Iapetus at this time (van Staal et al., 2012, 2020; Eusden et al., 2013). Avalonia and Ganderia may have been adjacent or along strike of each other (Murphy et al., 2009; van Staal et al., 2020) and may have rifted away from Gondwana at about the same time; the two crustal fragments were likely separated by an Acadian seaway (Schulz et al., 2008; van Staal et al., 2009, 2012). The rocks in the Ellsworth terrane in coastal Maine (Schulz et al., 2008) offer the best regional correlations with the Quinebaug-Marlboro belt. Our geochemical data, however, and that reported by Kay et al. (2017), suggest more of a possible backarc environment in southern New England than solely a proto-oceanic rift environment as proposed for the rocks in Maine (Schulz et al., 2008). The arc-rifting zone where the Mariana arc and the Mariana backarc basin converge may, in part, offer a possible modern analog (Pearce et al., 2005).

By the Late Ordovician in New England, the leading edge of Ganderia and marginal peri-Laurentian volcanic arcs had collided with Laurentia during the Taconic orogeny (Stanley and Ratcliffe, 1985; Karabinos et al., 1998, 2017; Ratcliffe et al., 1998; Aleinikoff et al., 2007, Macdonald et al., 2014, 2017; Valley et al., 2020). Variously named peri-Gondwanan arcs including the Tetagouche, Popelogan, Victoria, and Bronson Hill arcs approached Laurentia as the Iapetus Ocean closed (Aleinikoff et al., 2007; van Staal et al., 2016; Tremblay and Pinet, 2016; Valley et al., 2020). Events tied to the Taconic orogeny ca. 460–445 Ma involved the arrival of the Ganderian arc rocks along the Laurentian margin (e.g., Ratcliffe et al., 1998; van Staal and Barr, 2012; Karabinos et al., 2017; Macdonald et al., 2017; Valley et al., 2020). To the east, at the trailing edge of Ganderia in southern New England, arc to backarc signatures in the limited volcanic rocks found in the Nashoba and Tatnic Hill Formations support a waning arc to backarc environment prior to this time (this study; Kay et al., 2017). A backarc environment was also prevalent in Avalonia at about this time (Hamilton and Murphy, 2004; Jutras et al., 2020). Limited Ordovician detritus is indicated by the gap in 485–435 Ma detrital zircon populations and suggests that the source for the Nashoba and Tatnic Hill sedimentary rocks came from the older underlying Quinebaug and Marlboro Formations or Ganderian basement (Fig. 28), and not from Ordovician arcs now found farther to the west (e.g., Aleinikoff et al., 2007; Wintsch et al., 2007). The detrital zircon gap (Fig. 27) also supports the interpretation that this was the main period of deposition of protoliths of the metasedimentary units. Similar detrital zircon populations (peaks at 620–520 Ma) are found in the Silurian–Lower Devonian Arisaig Group in the Avalon terrane of Nova Scotia and are interpreted to have been deposited near the trailing edge of Avalonia (Murphy et al., 2004). The Arisaig Group also contains low ε_Nd isotope geochemistry values (−4.8 to −9.3 at 430 Ma; Murphy et al., 2004) that are similar to the Putnam-Nashoba terrane rocks (Fig. 25), raising the possibility that the Quinebaug-Marlboro belt and correlative rocks were providing detritus to nearby crustal fragments at this time.

By the late Silurian, the sedimentary rocks above the Putnam-Nashoba terrane in the overlying Merrimack belt were deposited, with the possibility that deposition of some of the rocks mapped as Tatnic Hill Formation in Connecticut may also span into the Silurian (Wintsch et al., 2007). Silurian metasedimentary rocks of the Merrimack belt are considered a Ganderian cover sequence (Fig. 28) consisting mostly of marine clastic rocks deposited above extended crust (Tremblay and Pinet, 2005, 2016). Farther west, rocks of the Connecticut Valley–Gaspé trough (CVGT, Fig. 28) span this same time period and basin formation may be related to late Silurian lithospheric delamination and crustal extension (Tremblay and Pinet, 2005, 2016; Rankin et al., 2007; McWilliams et al., 2010; Perrot et al., 2018). Regionally this event is considered the Salinic disturbance (Boucot, 1962), or Salinic orogeny (i.e., van Staal et al., 2009), possibly related to collision of Laurentia with the leading edge of Ganderia (van Staal et al., 2008, 2009). On the eastern margin of Ganderia, subduction of the Rheic Ocean, or an arm of the Rheic called the Acadian seaway, led to formation of the widespread Coastal Volcanic belt in Maine (Fig. 28) and its less voluminous extension into southern New England (van Staal and Barr, 2012; Pinan-Llamas and Hepburn, 2013; Wilson et al., 2017). The Silurian to Early Devonian calc-alkaline rocks are related to subduction prior to closure and collision in the Salinic or early part of the Acadian orogeny (Wones and Goldsmith, 1991; Watts et al., 2000; van Staal et al., 2009; Hussey et al., 2010; Barr et al., 2011; Pinan-Llamas and Hepburn, 2013; Hepburn et al., 2014). In Massachusetts, the Newbury Volcanic Complex (Zen et al., 1983) at the northern end of the Putnam-Nashoba terrane (Fig. 2) is a likely extension of the Coastal Volcanic belt (Pinan-Llamas and Hepburn, 2013). Igneous rocks that were emplaced into the Putnam-Nashoba terrane at this time include Preston Gabbro (418 ± 3 and 424 ± 5 Ma; this study; Zartman and Naylor, 1984), Millham Reservoir granofels (425 ± 3 Ma; Millham Reservoir “granulite” member of the Marlboro Formation of Acaster and Bickford, 1999), Sharpners Pond Diorite (430 ± 5 Ma; Zartman and Naylor, 1984), and Andover Granite (ca. 420 and ca. 419 Ma; Dabrowski et al., 2013; Dabrowski, 2014). Associated with this period of plutonism is the first distinct period of widespread Buchan-style metamorphism in the Putnam-Nashoba terrane, which is the earliest phase of metamorphism in the terrane (Stroud et al., 2009). Stroud et al. (2009) reported electron microprobe (EPMA) monazite ages from this metamorphism at ca. 435–400 Ma (Fig. 27), with an average age of 423 Ma (Hepburn et al., 2014). The oldest metamorphic ages from the two detrital zircon samples in our new study fall within this same time span in the range of ca. 435–400 Ma and corroborate the EPMA ages. Metamorphic monazite in the Fish Brook Gneiss yielded a TIMS age of 425 ± 3 Ma (Hepburn et al., 1995), further documenting this event.

Trace-element geochemical data suggest that late Silurian to Early Devonian arc-related calc-alkaline plutons in the Putnam-Nashoba terrane and adjacent overriding Merrimack belt developed in different settings as indicated by mid- to high-K, enrichment of LILE, LREE, Ba, and Sr, and negative Nb and Ta anomalies in the latter (Hepburn et al., 1995; Watts et al., 2000). The slightly younger plutonic rocks in the overlying Merrimack belt show an overlap in age with rocks in the underlying Putnam-Nashoba terrane rocks. Igneous rocks that were emplaced into the Merrimack belt at this time include Ayer Granodiorite (407 ± 4 Ma; Walsh et al., 2013b), Ayer Granodiorite at Eddy Pond (424 ± 3 Ma; Walsh et al., 2013a), Canterbury Gneiss (414 ± 3 Ma; Wintsch et al., 2007), Exeter diorite (407.4 ± 0.5 Ma; Hussey et al., 2016), and Newburyport quartz diorite (418 ± 1 Ma; Fargo and Bothner, 1995).

The Early Devonian Acadian orogeny is a widespread tectono-thermal event in the northern Appalachians. This even has been attributed to the collision between Avalonia and peri-Laurentian or peri-Gondwanan crustal fragments and the closing of the Acadian seaway (Fig. 28) (Bradley, 1983; Robinson et al., 1998; Bradley et al., 2000; van Staal and Barr, 2012; Tremblay and Pinet, 2016), although the timing of its impact on adjacent Avalonia has been lacking and debated (Wintsch et al., 2014). The Silurian to Devonian Acadian foreland basins of the Connecticut Valley–Gaspé trough and Central Maine basin are the largest metasedimentary basins in the northern Appalachians and show prominent peaks in magmatic activity at 407 and 379 Ma (Bradley et al., 2015; Bradley and O’Sullivan, 2016). The timing of peak metamorphism in the Putnam-Nashoba terrane has long been considered middle Paleozoic (Goldsmith, 1991c; Wintsch et al., 1992, 1993; Hepburn et al., 1995; Stroud et al., 2009). These findings are reflected in our data for the Putnam-Nashoba terrane where peak metamorphic ages range from ca. 400–380 Ma (Fig. 28) during the first regional, widespread upper amphibolite–facies metamorphic conditions and first major period of deformation (D₁) producing isoclinal folds of biotite granodiorite sills and leucosomes in the I-395 roadcut, dated at 409 ± 3 Ma. Metamorphism at this time is attributed to the Stroud et al. (2009) “M2” event, which they dated at ca. 390 Ma (Fig. 27). In this study and mapping in southern Massachusetts (Walsh et al., 2011a; Walsh and Merschat, 2015), we observed leucosomes associated with migmatization of the Nashoba Formation that formed during both the D1 and younger D2 events. Regionally, ages from migmatites range from Devonian (Hepburn et al., 1995; Stroud et al., 2009) to Carboniferous (Acaster and Bickford, 1999; Stroud et al., 2009), which reflects the onset of the development of the Putnam-Nashoba terrane high-grade tectonic wedge (Fig. 28).

By the Middle Devonian, the docking of the Meguma terrane, at least in Nova Scotia, to the modified Laurentian craton was under way and spanned the period from ca. 400–350 Ma during an event called the Neo-Acadian orogeny, postdating the classical Acadian orogeny (Robinson et al., 1998; van Staal et al., 2009). This orogenic activity is also informally called the “Famennian event” and spanned the period from ca. 375–345 Ma (Hibbard et al., 2010). Available data suggest that magmatic activity at this time is not as widespread as earlier events and may have been related to slab break-off of the Avalonian lithospheric mantle below Ganderia (Fig. 28; Aleinikoff et al., 2007). The Indian Head Hill pluton (349 ± 4 Ma, TIMS U-Pb zircon and titanite; Hepburn et al., 1995) intruded the Nashoba Formation, and the event is also recorded in TIMS U-Pb zircon ages of 365–333 Ma (Acaster and Bickford, 1999). In the Marlboro Formation, SHRIMP U-Pb data from metamorphic zircons from amphibolite indicate significant Carboniferous thermal events evidenced by metamorphic cores at 356 ± 4 Ma and rims at 339 ± 17 Ma (Walsh et al., 2011a). Several samples in our study show metamorphic zircons within this age bracket (Table 1).

Farther west in southern Connecticut in the Bronson Hill arc, the Hidden Lake gneiss intruded the core of the Killingworth dome during this interval (339 ± 3 Ma, SHRIMP; Aleinikoff et al., 2007). Similar Neo-Acadian zircon ages, at ca. 360–350 Ma are reported to the west and north where this event is more widespread in central New England in the Bronson Hill arc and Central Maine basin including the following:

• Hardwick Tonalite (e.g., 360 ± 1 and 361 ± 2 Ma, TIMS; Robinson and Tucker, 1992; Robinson et al., 1998)

• Wachusett tonalite of the Fitchburg Complex (359 ± 1 Ma, TIMS; Robinson et al., 1998)

• Fitzwilliam Granite (354 ± 1 Ma, TIMS; Robinson and Tucker, 1996; Robinson et al., 1998)

• Walker Mountain orthogneiss (356 ± 4 Ma, SHRIMP; Massey et al., 2017)

• Pegmatites in the Walker Mountain orthogneiss (353 ± 3 Ma, SHRIMP; Massey et al., 2017).

Exhumation of the Putnam-Nashoba terrane began following the Neo-Acadian event. Wintsch et al. (2014) describe the onset of exhumation as the result of tectonic wedging of Avalonia between Ganderian cover and its basement at the earliest stages of the Alleghanian orogeny, at ca. 340 Ma. Extrusion and channel flow of a high-grade, low viscosity layer in the middle crust may represent a viable mechanism (Beaumont et al., 2001; Buchanan et al., 2016; Severson et al., 2017; Severson, 2020). The protolith of the granofels at Black Hill was deposited, perhaps in a successor basin, sometime in the interval between ca. 371–340 Ma, and shows evidence of first metamorphism at ca. 340 Ma. The deep burial of the Putnam-Nashoba terrane led to intrusion of the Carboniferous granitoids in the Nashoba Formation, perhaps as the result of either decompression melting or crustal thickening at ca. 330–320 Ma (Wintsch et al., 2014). Importantly, the granofels at Black Hill lacks the high-grade metamorphic mineral assemblages and Carboniferous granitoids found in the adjacent rocks, thus the rocks at Black Hill represent a younger metasedimentary sequence that may be entirely fault-bounded.

Evidence for Carboniferous exhumation included ⁴⁰Ar/³⁹Ar cooling ages from amphibole ranging from 354 to 325 Ma (Hepburn et al., 1987). Regionally, ⁴⁰Ar/³⁹Ar amphibole cooling ages indicate diachronous exhumation from north to south (Fig. 3). More recent ⁴⁰Ar/³⁹Ar data show uplift and cooling through the amphibole closure temperature (~500°C) by ca. 350–330 Ma in the Mississippian and cooling through the muscovite closure temperature (~350°C) by ca. 320 Ma in the Pennsylvanian (Wintsch et al., 1992; Attenoukon, 2009; Reynolds, 2012) (Fig. 3). Both closure temperatures confirm minimal high-temperature thermal resetting of the Putnam-Nashoba terrane during the Alleghanian (Goldsmith, 1991c; Wintsch et al. 1992, 1993, 1998; Hepburn, 2004). However, the age of this resetting varies along the length of the terrane (Fig. 3) with older ages preserved to the north. The distribution of younger ⁴⁰Ar/³⁹Ar ages due to Alleghanian thermal resetting in southern New England closely approximates the timing of intrusion of pegmatite, tonalite, and migmatization at ca. 330–325 Ma recognized in our study (Fig. 27). The location of the Putnam-Nashoba terrane, as the lowest structural level and deepest of all the slices of the Ganderian cover rocks, affords an ideal place to document this late Paleozoic thermal overprint. By ca. 300 Ma, the terrane experienced the local potentially prograde growth of titanite, retrogression of high grade assemblages, faulting, and continued exhumation due to the Alleghanian orogeny.

This study establishes an Ediacaran to Cambrian age for the protoliths of the Quinebaug-Marlboro belt, which developed in an arc to backarc complex along the Ganderian margin. Ages from the Quinebaug-Marlboro belt span a period from ca. 540–500 Ma. The existence of older ages in the south and younger ages in the north of the Quinebaug-Marlboro belt suggests that the entire belt preserves a cryptic, internal stratigraphy that is truncated by subsequent faulting. Rocks of the overlying Nashoba and Tatnic Hill Formations were deposited from ca. 485–435 Ma. Provenance of detrital zircon samples from the Nashoba and Tatnic Hill Formations is distinctly peri-Gondwanan, showing peaks between ca. 500 and 650 Ma similar to ages from Ganderian or Avalonian arcs (Fyffe et al., 2009; Pollock et al., 2009, 2012).

Trace-element data show that some of the mafic rocks have significant negative Nb-Ta anomalies similar to arc-like rocks, whereas other mafic rocks lack negative Nb-Ta anomalies and thus are more similar to rocks from E-MORB to OIB or backarc settings.

Geochemical and geochronological data support a regional correlation of the Quinebaug-Marlboro belt with the Ellsworth terrane in Maine and extend the temporal and spatial range of the Penobscot arc and its precursor from Connecticut through Maine, to New Brunswick and Newfoundland. The older ca. 540 Ma rocks of the Quinebaug Formation predate the traditional age of the Penobscot arc (ca. 515–485 Ma; Johnson et al., 2012; van Staal and Barr, 2012) but match a period of subduction-related calc-alkaline arc magmas in New Brunswick (Fyffe et al., 2011).

SHRIMP analyses of detrital zircon from granofels at Black Hill yield ages indicating deposition between ca. 371–340 Ma, showing that the protolith was a Devonian (or younger) sediment that is not a part of the Quinebaug Formation.

All sampled rocks show evidence of late Silurian to Carboniferous metamorphism. Metamorphic zircon, zircon overgrowths, and titanite ages range from 410 to 305 Ma with peaks at ca. 400, 350, and 325 Ma corresponding to Acadian, Neo-Acadian, and early Alleghanian orogenesis. SHRIMP data show that a Silurian metamorphic event may be limited to contact thermal effects from calc-alkaline plutons such as the Preston Gabbro (dated here at 418 ± 3 Ma). Silurian plutons place an upper limit on the age of the sedimentary and volcanic rocks in the Putnam-Nashoba terrane and the Quinebaug-Marlboro belt at ca. 420 Ma. The protoliths of the entire terrane span a maximum of ~120 million years ranging in age from ca. 540–420 Ma.

Exhumation of the Putnam-Nashoba terrane began following the Neo-Acadian event due to escape of a hot, high-grade tectonic wedge, or orogenic channel, in the earliest stages of the Alleghanian orogeny, at ca. 340 Ma. The continued relatively deep crustal position of the terrane led to intrusion of the Carboniferous granitoids in the Nashoba Formation, perhaps as the result of either decompression melting, slab break-off, or crustal thickening at ca. 330–320 Ma.

Deciphering the protoliths of high-grade metamorphic rocks is a challenge that can be overcome through the combined characterization of mapping, structures, comprehensive geochemistry, and precise U-Th-Pb geochronology involving high-resolution mineral imagery—all of which are essential tools for understanding the nature of complex enigmatic terranes.

Zircon, monazite, and titanite were extracted from rock samples collected at outcrops in Massachusetts and Connecticut using standard mineral separation techniques including crushing, pulverizing, Wilfley Table, magnetic separator, and heavy liquids. For igneous rocks, individual grains of each mineral were handpicked; detrital zircons from metasedimentary samples were sprinkled onto double-stick tape. All mineral grains were mounted in epoxy, ground to nearly half-thickness using 2500-grit wet-dry sandpaper, and polished using 6 μm and 1 μm diamond suspension. All grains were photographed in reflected and transmitted light on a petrographic microscope. Using the USGS JEOL 5800LV scanning electron microscope, compositional zoning of zircon was imaged in cathodoluminescence (CL), whereas zoning in monazite and titanite is shown by backscattered electron (BSE) imaging.

Zircon, monazite, and titanite were analyzed on the USGS/ Stanford sensitive high-resolution ion microprobe-reverse geometry (SHRIMP-RG) in Palo Alto, California. The analytical procedures for SHRIMP U-Pb dating follow methods described in Williams (1998). The primary oxygen ion beam, operated at ~4–6 nA, excavated an area of ~25–30 μm in diameter to a depth of ~1 μm. The magnet was cycled through the mass stations five times per analysis for igneous or metamorphic zircon, metamorphic monazite, and metamorphic titanite and four times for detrital zircons. Elemental fractionation was corrected by analyzing mineral standards of known age every fourth analysis, including: (1) zircon standard R33 (419 ± 1 Ma; Black et al., 2004), (2) monazite standard 44069 (425 ± 1 Ma; Aleinikoff et al., 2006), and titanite standard BLR-1 (1047 Ma; Aleinikoff et al., 2007). Calculated concentrations of Pb and U are believed to be accurate to ~10%–20%.

U-Pb isotopic data from zircon, monazite, and titanite were reduced using the Squid 1 or 2 (Ludwig, 2002, 2009) and plotted using Isoplot/Ex version 3.0 (Ludwig, 2003). Common Pb in zircon and monazite is corrected using model Pb isotopic compositions from Stacey and Kramers (1975). Corrected data for zircon are plotted on ²³⁸U/²⁰⁶Pb-²⁰⁷Pb/²⁰⁶Pb Tera-Wasserburg Concordia diagrams to visually assess the degree of discordancy of each data point. Ages for igneous rocks are determined by calculating the weighted average of ²⁰⁶Pb/²³⁸U. Nearly all analyses of zircon and monazite have very low common Pb contents (common ²⁰⁶Pb is mostly less than 0.2% of total ²⁰⁶Pb; Table 1) so that the individual age of each spot analysis is quite insensitive to the common Pb correction. Weighted average calculations of ²⁰⁶Pb/²³⁸U ages incorporate both the 2σ external spot-to-spot error of the standard and the 2σ error of the mean of the standard (Ludwig, 2003). All errors for weighted average ages are reported at the 95% confidence limit. For detrital zircons, ²⁰⁷Pb/²⁰⁶Pb age data that are greater than 10% discordant are excluded from relative probability plots. Although analysis of at least 100 detrital zircon grains is typically considered adequate for characterization of the main provenance age groups (Gehrels, 2014), in our study, because of limited instrument availability and financial constraints, only ~60 analyses were done for two (of three) samples. Thus, relative probability curves for samples 040814-1 and WO-125C are less statistically robust than the curve generated from data for sample 616H1. To construct the plots, we use ²⁰⁶Pb/²³⁸U ages for grains younger than ca. 1000 Ma and ²⁰⁷Pb/²⁰⁶Pb ages for grains older than ca. 1000 Ma (Black et al., 2004).

Geochemical samples were split and analyzed by major (X-ray fluorescence [XRF]) and trace (ICPMS) elements, mostly at Activation Labs, Ancaster, Ontario, Canada (Table 4). Major and trace elements for samples WO-125A and WO-125B were analyzed at SGS Canada, Inc., Toronto, Ontario, Canada, and samples GR-098A and GR-098B were analyzed at Acme Labs, Vancouver, British Columbia, Canada (Table 4). Representative samples were analyzed for Nd, Sr, and Pb isotopic compositions using a multicollector, automated Spectromat/Finnigan-MAT 262 mass spectrometer at the U.S. Geological Survey, Reston, Virginia, USA. Analytical techniques for isotope geochemistry are given in Ayuso and Till (2014). Long-term reproducibility of the Nd isotopic work was monitored using the La Jolla standard: average value of ¹⁴³Nd/¹⁴⁴Nd = 0.511845 ± 5 (n = 22) and JNd-1 standard: average value of ¹⁴³Nd/¹⁴⁴Nd = 0.512111 ± 5 (n = 13). For the Sr isotopic analyses, National Institute of Standards and Technology (NIST) Standard Reference Material (SRM) 987 yielded an average value of ⁸⁷Sr/⁸⁶Sr = 0.710245 ± 5 (n = 25). Lead isotope ratios of whole rocks were corrected for mass fractionation by ~0.12% amu-1 according to replicate measurements of National Bureau of Standards (NBS) 981 (n = 29). Total blanks during the course of this study were <40 pg for Nd, <50 pg for Sr, and 70 pg for Pb; they are insignificant relative to the Nd, Sr, and Pb abundances. Depleted mantle model ages were calculated as in the model of DePaolo (1981). Table 6 shows the results of our isotopic analyses.

We thank Joe Wooden (U.S. Geological Survey [USGS] in Menlo Park, California) for help with SHRIMP analyses. Renee Pillers of the USGS in Denver, Colorado, helped with mineral separations. John Jackson of the USGS in Reston, Virginia, assisted with geochemical sample collection and preparation. Joe Kopera of the University of Massachusetts at Amherst assisted with geochemical sample collection and discussions. Beneficial conversations with Chris Hepburn, Yvette Kuiper, and Wes Buchanan aided this study. Special thanks go to Klaus Schulz, Arthur Merschat, and Nancy Stamm of the USGS for their suggestions and critical reviews. We sincerely thank Brendan Murphy, Yvette Kuiper, and Robert Miller for their constructive, helpful reviews. This work was supported by the USGS National Cooperative Geologic Mapping Program and by National Science Foundation grants EAR-9418203 awarded to Wintsch and EAR-0510293 awarded to Wintsch and Mike Dorais. Any use of trade, firm, or product names is for descriptive purposes only and does not imply endorsement by the U.S. Government.