2 The Nashoba terrane in eastern Massachusetts (Fig. 1) is an early Paleozoic arc/back-arc complex that, based on Sm/ Nd isotope compositions and model ages, developed on an older crustal basement (e.g., Goldsmith 1991a; Hepburn et al. 1995; Kay et al. 2017; Walsh et al. 2021). It has been interpreted as the trailing edge of Ganderia in SE New England (e.g., Hibbard et al. 2006; van Staal et al. 2009; Walsh et al. 2011b, 2015, 2021; Kay et al. 2017). The Nashoba terrane consists of highly metamorphosed and multiply deformed early Paleozoic volcanic and volcanogenic sedimentary rocks that are intruded by abundant middle Paleozoic dioritic and granitic plutons (e.g., Wones and Goldsmith 1991; Zartman and Marvin 1991; Hepburn et al. 1995). It is separated from the Avalonian rocks to the southeast by the Burlington mylonite zone, which was later overprinted by the more brittle Bloody Bluff fault zone (Castle et al. 1976; Zen et al. 1983; Skehan et al. 1998; Kohut and Hepburn 2004), and from the Merrimack belt to the northwest by the Clinton-Newbury fault zone (Figs. 1, 2; Skehan 1968; Castle et al. 1976; Zen et al. 1983; Goldsmith 1991b). The stratified units generally dip moderately to steeply northwest (Bell and Alvord 1976; Goldsmith 1991a). In the southeast, the terrane consists largely of mafic metavolcanic and metasedimentary rocks of the Marlboro Formation, while to the northwest predominantly volcanogenic metasedimentary rocks of the Nashoba Formation dominate (Bell and Alvord 1976; Goldsmith 1991a; Hepburn et al. 1995). Other major units include a schistose unit (Tadmuck Brook Schist), an orthogneiss (Fish Brook Gneiss) and a paragneiss (Shawsheen Gneiss) (Fig. 2; Bell and Alvord 1976; Zen et al. 1983; Goldsmith 1991a; Hepburn et al. 1995). The metavolcanic and metasedimentary units in the southeast are interpreted to have been part of an arc or formed in a near-arc basin, while those in the northwest have been interpreted to represent largely contemporaneous to somewhat younger sedimentary deposits with minor volcanic input in a back-arc basin developed farther from an arc (e.g., Goldsmith 1991a; Hepburn et al. 1995; Kay et al. 2017).

3 The terrane is intruded by abundant Silurian to Carboniferous granitic and dioritic plutons. Some of these plutons have been dated using various methods of U–Pb zircon geochronology. These include the 430 ± 5 Ma Sharpner’s Pond Diorite (Zartman and Naylor 1984), two phases of the Andover Granite, which are 419 ± 0.5 Ma (Dabrowski 2013; Dabrowski et al. 2013) and 412 ± 2 Ma (Hepburn et al. 1995), the 420.5 ± 0.5 Ma Sudbury granite (Dabrowski 2013; Dabrowski et al. 2013), the 385 ± 10 Ma Straw Hollow diorite (Acaster and Bickford 1999) and the 349 ± 4 Ma Indian Head Hill Granite (Hepburn et al. 1995).

4 Previous age determinations on rocks associated with the stratified units in the Nashoba terrane are limited. They include a ~540 Ma mafic boudin in the Quinebaug Formation (a unit in the Putnam terrane to the south in Connecticut correlated with the Marlboro Formation), a 515 ± 4 Ma date on the Grafton Gneiss that cross-cuts part of the Marlboro Formation, and a 500.6 ± 4.2 Ma date on a feldspathic granofels in the Marlboro Formation structurally above the Grafton Gneiss (U–Pb zircon SHRIMP, Walsh et al. 2009, 2011a, b, 2021). The granofels was derived from a proximal volcanic source and its age is interpreted as the approximate age of deposition (Walsh et al. 2021). In addition to the ~540–500 Ma volcanic rocks of the Marlboro Formation and the Grafton Gneiss, the Fish Brook orthogneiss (Fig. 2) has also been dated as Cambrian (499 +6/-3 Ma, U–Pb zircon ID-TIMS, Hepburn et al. 1995).

5 Deformation and metamorphism in the Nashoba terrane are primarily a result of the late Silurian–Devonian Acadian orogeny that resulted from accretion of Avalonia and its subsequent westward convergence (e.g., Hepburn et al. 1995; van Staal et al. 2009, 2020; Kuiper et al. 2014; Walsh et al. 2015, 2021; Severson 2020). Post-Acadian zircon and monazite in the Nashoba terrane have been interpreted to be a result of the late Devonian Neoacadian and/or Pennsylvanian–Permian Alleghanian orogenic events and/or hydrothermal fluids (Wintsch et al. 2007; Stroud et al. 2009; Loan 2011; Buchanan et al. 2017; Walsh et al. 2021).

6 Late Ordovician to Early Devonian calcareous and pelitic turbiditic rocks form extensive cover successions on older Ganderian rocks throughout the northern Appalachians from southern New England to Newfoundland (e.g., Williams 1978; Tucker et al. 2001; Tremblay and Pinet 2005; Hibbard et al. 2006). In New England, these sedimentary rocks show increasing Laurentian affinity to the west where they define the Connecticut Valley-Gaspé trough (Fig. 1; e.g., Hibbard et al. 2006; Rankin et al. 2007). To the east (Fig. 1), two extensive depocenters of these rocks, the Fredericton trough and Central Maine basin, are defined by their location with respect to inliers of Cambrian–Ordovician basement rocks (e.g., Reusch and van Staal 2012; Ludman et al. 2018). However, the terminology that has been applied to these rocks across New England is complex and varied as it includes both structural and sedimentary terms (i.e., belt, synclinorium, basin, trough) as well as various stratigraphic names and correlations. Figure 3 summarizes the principal stratigraphic names by location for the relevant Late Ordovician to Early Devonian cover rocks in eastern and central New England and southern New Brunswick. Below we explain, from north to south, some of the definitions of basins used by others, and the rationale behind them. We then summarize the definitions used in this contribution, which are largely based on this previous work.

7 In eastern Maine and New Brunswick, the Fredericton trough lies between the Late Ordovician and older rocks of the St. Croix terrane to the SE and the Miramichi terrane to the NW (Fig. 1; e.g., Reusch and van Staal 2012; Ludman et al. 2017, 2018; Dokken et al. 2018). The broader Central Maine basin, including the along-strike Aroostock- Matapedia basin to the north, formed between the Miramichi terrane to the SE and older rocks of the Bronson Hill arc and related inliers to the NW (Fig. 1; Osberg et al. 1985; Tucker et al. 2001; Fyffe et al. 2009, 2011; Reusch and van Staal 2012; Bradley and O’Sullivan 2017; Ludman et al. 2017, 2018). Both the Fredericton trough and Central Maine basin contain late Ordovician (?) and Silurian sedimentary rocks. Deposition ended in the Fredericton trough by the late Silurian as it was deformed prior to intrusion by the 421.9 ± 2.4 Ma Pocomooshine gabbro-diorite (West et al. 1992; Ludman et al. 2018), while deposition continued into the Devonian in the Central Maine basin (Osberg 1988; Bradley et al. 2000; Tucker et al. 2001; Bradley and O’Sullivan 2017; Ludman et al. 2018). In south-central Maine and coastal southern Maine the Fredericton trough includes the Late Ordovician (?) and Silurian Ghent, Appleton Ridge and Bucksport formations (Figs. 1, 3; Hussey 1988; Berry and Osberg 1989; Hussey et al. 2010; West and Condit 2016; Cartwright et al. 2019) that lie southeast of Miramichi terrane-correlative Ordovician and older rocks in the Falmouth-Brunswick Group and Casco Bay Group (West et al. 2004, 2006; Hussey et al. 2010; Reusch and van Staal 2012; West and Hussey 2016, Hussey and West 2018).

8 In southwestern Maine and southeastern most New Hampshire, Late Ordovician and Silurian rocks that are approximately along strike, and can be correlated with those of the Fredericton trough and eastern part of the Central Maine basin form the Merrimack basin or Merrimack trough (Lyons et al. 1997; Fig. 1) and were originally included in the Merrimack Group (Fig. 3; Eliot, Kittery and Berwick formations; e.g., Hitchcock 1877; Katz 1917; Billings 1956; Lyons et al. 1997; Fargo and Bothner 1995; Hussey et al. 2010; see Hussey et al. 2010 for details). Hussey et al. (2010), however, redefined the Merrimack Group in SW Maine and New Hampshire to include only sedimentary rocks interpreted as Late Ordovician–early Silurian (Eliot and Kittery formations) and correlated them with those in the Fredericton trough. They correlated sedimentary rocks interpreted to be late Silurian–Early Devonian (Berwick Formation) with the Central Maine basin. Hussey et al. (1999, 2010) suggested that their redefined Late Ordovician–early Silurian Merrimack Group, in addition to the Fredericton trough deposits, including the Bucksport and associated formations of south-central Maine, represent a potentially continuous single basin and named it the Merribuckfred basin (Fig. 3). However, it is now known that the Fredericton trough also includes late Silurian deposits (Dokken et al. 2018; Ludman et al. 2018). Therefore, the Merribuckfred basin now refers more to a location than an age range and may include multiple basins deposited through time. We will use Merribuckfred basin as a general term for the Late Ordovician–Silurian basins east of the Miramichi terrane and its equivalents to the south, to differentiate them from the Central Maine basin on the west. We will use Fredericton trough and Merrimack basin when referring to deposits in specific locations (Figs. 1, 3).

9 The Central Maine basin continues southwest from eastern Maine and New Brunswick into southeastern New Hampshire (Fig. 1; Osberg et al. 1985; Zen et al. 1983; Lyons et al. 1997; Tucker et al. 2001; Hibbard et al. 2006; Wintsch et al. 2007; Hussey et al. 2010; Reusch and van Staal 2012). However, Ordovician and older rocks of the Miramichi terrane, Casco Bay Group or Falmouth-Brunswick Group are not present south of southwestern Maine, making the boundary between rocks in the Central Maine and Merribuckfred basins difficult to distinguish. In SE New Hampshire, Dorais et al. (2012) and Reusch and van Staal (2012) interpret these basins to be separated by the Ganderian Neoproterozoic Massabesic Gneiss Complex (Fig. 1). However, Hussey et al. (2010, 2016) separate these basins east of the Massabesic Gneiss Complex along the Nonesuch River fault in southwestern Maine and its continuation, the Calef fault, in New Hampshire (Fig. 1).

10 In Massachusetts, there is no clearly identified boundary between rocks correlative with those in the Merribuckfred or Central Maine basins. Therefore, Zen et al. (1983) and Robinson and Goldsmith (1991) included all the late Ordovician (?), Silurian and early Devonian metasedimentary rocks between Ordovician and older rocks in the Nashoba terrane to the east and the Bronson Hill anticlinorium to the west in the Merrimack belt (Figs. 1, 2). However, on some maps, the Central Maine basin continues south into Massachusetts between the Merrimack belt and the Bronson Hill anticlinorium (e.g., Hibbard et al. 2006; Rankin et al. 2007; Wintsch et al. 2007, 2014; Reusch and van Staal 2012; Kopera and Walsh 2014, Walsh et al. 2021). Here, we will follow the usage of Zen et al. (1983) and Robinson and Goldsmith (1991) and use the inclusive term Merrimack belt to include the Worcester, Oakdale, Paxton, Eliot, Kittery and Berwick formations (Figs. 2, 3) in Massachusetts and SE New Hampshire, but recognize that the Merrimack belt likely includes rocks that can be correlated with those in both the Merribuckfred and Central Maine basins (Tucker et al. 2001; Wintsch et al. 2007; Hussey et al. 2010). The Tower Hill Formation occurs along the Clinton-Newbury fault zone on the east side of the Merrimack belt in east-central Massachusetts (Fig. 2; Zen et al. 1983; Robinson and Goldsmith 1991). However, its placement within the Merrimack belt has been debated (see Robinson and Goldsmith 1991) and it will be considered separately below.

11 The Merrimack belt, excluding the Tower Hill Formation, generally consists of a thick succession of metamorphosed calcareous and pelitic turbiditic units containing sandstone, siltstone, phyllite and impure quartzite (e.g., Grew 1973; Peck 1976; Robinson 1981; Zen et al. 1983; Robinson and Goldsmith 1991; Hussey and Bothner 1993; Lyons et al. 1997; Hussey et al. 2010). Because of (1) the lack of fossils, (2) the presence of thick successions of similar rock types in various areas, (3) a paucity of clear sedimentary structures to give facing directions, (4) varying degrees of metamorphism across and along strike, and (5) structural complexities due to several generations of folds, the internal stratigraphic order and along-strike correlations of units of the Merrimack belt has long been debated (e.g., Billings 1956, Robinson and Goldsmith 1991, Hussey et al. 2010). Robin-son and Goldsmith (1991) presented a thorough discussion of the stratigraphic arguments for rocks in the Merrimack belt in central Massachusetts and Hussey et al. (2010) for SE New Hampshire. These are not repeated or further discussed here.

12 The Merrimack belt in Massachusetts is divided into six stratigraphic-tectonic subbelts by Robinson and Goldsmith (1991). Our samples are from the eastern three: the Wachusett Mountain, Nashua and Rockingham subbelts (Fig. 2). The Rockingham subbelt continues into SE New Hampshire and contains the Eliot, Kittery and Berwick formations of the original Merrimack Group (Hitchcock 1877; Katz 1917; Billings 1956; Lyons et al. 1997; Fargo and Bothner 1995; Hussey et al. 2010). The Tower Hill Formation is within Robinson and Goldsmith’s (1991) Nashua subbelt, but occurs along the Clinton-Newbury fault zone adjacent to the contact with the Nashoba terrane and will be discussed separately.

13 The metamorphic grade of the sedimentary rocks of the Merrimack belt in central Massachusetts varies from lower greenschist facies in the southeast to upper amphibolite facies in the northwest (Thompson and Norton 1968; Zen et al. 1983; Goldsmith 1991b; Robinson and Goldsmith 1991; Lyons et al. 1997). The Merrimack belt has been intruded by a series of Silurian to Devonian granitic to dioritic plutons (Gore 1976; Zen et al. 1983; Zartman and Naylor 1984; Robinson and Goldsmith 1991; Wones and Goldsmith 1991; Zartman and Marvin 1991; Bothner et al. 1993, 2009; Far-go and Bothner 1995; Lyons et al. 1997; Watts et al. 2000; Wintsch et al. 2007; Walsh et al. 2013a, b, 2015, 2021; Charnock 2015) and is deformed by at least four generations of folds, including large recumbent folds (Peck 1975, 1976; Hepburn 1976; Robinson 1978, 1981; Goldstein 1992, 1994; Goldsmith 1991b; Hussey and Bothner 1993; Kopera and Walsh 2014; Kuiper et al. 2014). Deformation and metamorphism in the Merrimack belt resulted from Acadian, Neoacadian and/or Alleghanian events (Zen et al. 1983; Goldstein 1994; Attenoukon et al. 2006; Attenoukon 2009; Kuiper et al. 2014; Kopera and Walsh 2014).

20 Sample locations and LA-ICPMS data are presented in Appendix Tables A1 and A2, respectively. Rock unit and sample descriptions, and data for detrital zircon based on the criteria outlined above are presented in this section (Fig. 4). Data from metamorphic zircon are discussed separately below.

21 The Marlboro Formation occurs in the southeastern Nashoba terrane (Fig. 2). It is composed largely of hornblende-plagioclase amphibolite, but it also contains felsic granofels, gneiss and metasedimentary rocks that commonly include rusty weathering sillimanite-bearing schist (Bell and Alvord 1976; DiNitto et al. 1984; Goldsmith 1991a; Kopera et al. 2006). The amphibolites include rocks with basaltic and basaltic-andesite compositions with inferred arc, MORB and alkalic signatures. Whole-rock major and trace-element geochemistry and Sm–Nd isotopes indicate that they formed in a primitive volcanic arc/back-arc setting (DiNitto et al. 1984, Goldsmith 1991a, Kay et al. 2017) with minimal crustal contamination (Kay et al. 2009, 2017; Walsh et al. 2015, 2021). This arc has been interpreted to be the source for the majority of the metasedimentary rocks in the Nashoba terrane (Hepburn and Munn 1984, Goldsmith 1991a, Kay et al. 2017).

22 Zircon in metasedimentary rocks of the Marlboro Formation is rare, and commonly metamict. In order to obtain enough grains for a statistical representation, the Marlboro Formation was sampled in four different locations (Fig. 2; Loan 2011). Of the approximately 150 kg of rock processed, a total of only 9 zircon grains yielded concordant data. The data were combined as one sample, MLMRC (Table A2). Sample MLMR1 is from a rusty weathering, black to dark-grey garnet-muscovite-biotite-quartz (± sillimanite) schist interlayered with a biotite-quartz-hornblende-plagioclase amphibolite. Garnet and sillimanite are present in thin layers within the schist and the unit is locally mylonitic. The sample was collected from the pelitic layers only and yielded four concordant analyses. Sample MLMR2 is from a silvery to dark-grey, rusty weathering, fine-grained garnet-biotitemuscovite-quartz schist. It resulted in one concordant analysis. Sample MLMR5 showed two distinct layers in hand specimen: (1) a muscovite-rich schistose layer and (2) a coarser-grained quartz-rich layer with larger, black, quartz crystals. Black quartz forms due to radiation damage (Klein and Hurlbut 1993). The radiation damage extended to the zircons, resulting in rusty, amorphous, highly metamict grains that gave no concordant data. Sample MLMR6 is from a rusty, heavily weathered garnet-bearing schist that yielded four concordant analyses.

23 The youngest detrital zircon analysis from the combined Marlboro Formation samples yielded a ²⁰⁶Pb/²³⁸U age of 470 ± 46 Ma. Sample MLMR6 contained the oldest grain found in the Nashoba terrane with a core of ~3.36 Ga and a rim of ~2.58 Ga. The relative probability plot of the Marlboro Formation shows a peak in age at ~532 Ma and a minor peak at ~642 Ma (Fig. 4a).

24 The Shawsheen Gneiss (Fig. 2) is separated from the Nashoba Formation by the Fish Brook Gneiss and the Ass-abet River fault zone (Zen et al. 1983; Goldsmith 1991a). It is generally quartz-plagioclase-muscovite-biotite schist to paragneiss, which is interpreted to have been derived from the detritus of volcanic rocks of intermediate to mafic composition (Olszewski 1980). Sample MLSG1 is a medium- grained garnet-plagioclase-muscovite-biotite-quartz (± sillimanite) schist to gneiss collected from multiple locations around an industrial park to ensure that a statistical representation of detrital zircon was analyzed.

25 The Shawsheen Gneiss was previously dated by Olszewski (1980) using U–Pb analyses of zircon on multiple-grain fractions. The resulting discordia chord yielded 2042 ± 52 Ma upper and 517 ± 16 Ma lower intercept ages, but because these are based on only three zircon fractions (Olszewski 1980), the interpretation may not be accurate. Sample MLSG1 yielded 100 concordant detrital zircon analyses. The weighted average ²⁰⁶Pb/²³⁸U age of the three youngest detrital zircon analyses is 470 ± 22 Ma (MSWD = 0.26). The main age population in the sample is ~562 Ma, and a minor population exists at ~646 Ma (Fig. 4b). In addition, there are small age clusters at ~2.1–2.0 Ga, ~1.8–1.3 Ga and ~1.1 Ga, and a single grain at ~2.4 Ga.

26 The Nashoba Formation occupies about one-third of the Nashoba terrane and is composed largely of biotite-feldspar gneiss and biotite ± sillimanite schist with subordinate calcsilicate gneiss, impure feldspathic quartzite and pelitic schist (Bell and Alvord 1976; Zen et al. 1983; Hepburn and Munn 1984; Goldsmith 1991a). Hornblende amphibolite dominates the Boxford Member (Bell and Alvord 1976) along the eastern margin of the formation. The Nashoba Formation is metamorphosed in the sillimanite and sillimanite - K-feldspar zones. Samples of the gneiss, schist and calc-silicate gneiss were analyzed.

27 The Nashoba Formation gneiss (sample MLNB1; Fig. 2) is from a garnet-biotite-muscovite-plagioclase-quartz- (± sillimanite) gneiss. The unit contained interbedded layers of mylonitic gneiss with a few ~4 cm thick strongly sheared quartz veins. Areas with large quartz inclusions were not used. The sample yielded 84 concordant detrital zircon analyses. The three youngest grains yielded a weighted average ²⁰⁶Pb/²³⁸U age of 462 ± 20 Ma (MSWD = 0.061). The major age population is ~544 Ma (Fig. 4c). In addition, there are a few ~2.8–2.6 Ga, ~2.4 Ga, ~2.1 Ga, ~1.6 Ga, ~1.4–1.1 Ga zircons.

28 The Nashoba Formation schist (sample MLNS1; Fig. 2) is a garnet-muscovite-biotite-quartz- (± sillimanite) schist. Of the abundant schistose units in the Nashoba terrane this sample was selected because of its minimal quartz vein inclusions. However, the sample yielded only 4 concordant detrital zircon analyses, the other 118 concordant analyses being interpreted as metamorphic (Fig. 4d). The youngest detrital zircon in the Nashoba Formation schist yielded a ²⁰⁶Pb/²³⁸U age of 477 ± 32 Ma. Two remaining detrital grains yielded ²⁰⁶Pb/²³⁸U ages of 538 ± 20 Ma, 611 ± 94 Ma, and one a ²⁰⁷Pb/²⁰⁶Pb age of 1694 ± 32 Ma.

29 The Nashoba Formation calc-silicate gneiss outcrop (sample MLBM1; Fig. 2) includes rock types ranging from rusty garnet-bearing schist to calc-silicate gneiss and amphibolite. The calc-silicate gneiss sampled contains diopside, actinolite, phlogopitic biotite, and abundant titanite. It likely formed from original more dolomitic layers (Hepburn and Munn 1984). The unit is deformed by a shear zone separating the more calcareous layers of the outcrop from the schistose layers. Samples were taken away from the shear zone and included a mixture of calc-silicate rocks with and without the micaceous lenses. No detrital zircon was found and all zircon analyzed is interpreted to be of metamorphic or carbonate fluid/hydrothermal origin as discussed below.

30 The Tadmuck Brook Schist is a rusty-weathering micaceous schist with thin quartzose layers that lies along the northwestern boundary of the Nashoba terrane (Fig. 2) adjacent to the Clinton-Newbury fault zone. The metamorphic grade increases from greenschist facies in the NW to upper amphibolite facies in the SE within the ~1 km thick unit (Jerden and Hepburn 1996; Jerden 1997). It has been uncertain whether the Tadmuck Brook Schist is part of the Nashoba terrane or of the Merrimack belt (Goldsmith 1991a). Sample (MLTMBC; Fig. 2) is from a sillimanite-bearing sulfidic mica schist in the high-grade (southeastern) member and contains thin quartz-rich layers. The Tadmuck Brook Schist has unique mineralogy as it is the only unit that contains scheelite, a fluorescent tungsten-bearing mineral that is commonly associated with high-temperature hydrothermal veins and contact metamorphism in skarns (Klein and Dutrow 2007). The unit also contains abundant titanite. This sample yielded 47 concordant detrital zircon analyses. Due to the small grain size, all the zircons in this sample were ablated using a 30×30 μm square raster, which led to increased uncertainty. The youngest three detrital zircon analyses yielded a weighted average of ²⁰⁶Pb/²³⁸U ages of 462 ± 47 Ma (MSWD = 0.049). The sample had a major ~537 Ma age peak, and scattered ~2.2–1.9 Ga, ~1.7 Ga and ~1.4–1.1 Ga ages (Fig. 4e).

31 Youngest zircon grains of selected samples were re-dated using CA-TIMS methods (Table 1) with the purpose to better constrain the maximum depositional age of the formations. While three or four grains per sample were selected for analysis, not all survived sample preparation and chemical abrasion. Only one grain remained from the Nashoba terrane, sample MLNB1 from the Nashoba Formation gneiss. Its ²⁰⁶Pb/²³⁸U CA-TIMS age is 465.9 ± 0.7 Ma, which is within error of the weighted average of ²⁰⁶Pb/²³⁸U LA-ICPMS ages of 462 ± 20 Ma for the three youngest grains from this sample (Table A2; see above).

32 Nashoba terrane samples from the five formations other than the calc-silicate gneiss contained 159 zircons interpreted as metamorphic per the criteria listed above (Fig. 5). They range from ~454 Ma to 324 Ma, with the results for the individual formations as follows: Marlboro Formation, ~443–324 Ma (weighted average of ²⁰⁶Pb/²³⁸U ages of 396 ± 20 Ma, n = 10, MSWD = 1.6); Shawsheen Gneiss, 422 ± 19 Ma (n = 5, MSWD = 0.27); Nashoba Formation gneiss ~454–353 Ma (or weighted average age of 408 ± 13 Ma, n = 22, MSWD = 5.6); Nashoba Formation schist ~433–327 Ma (or weighted average age of 377.2 ± 4.4 Ma, n = 118, MSWD = 1.9); and Tadmuck Brook Schist 433 ± 33 Ma (n = 4, MSWD = 0.031) (cf. Table A2).

33 The calc-silicate gneiss (MLMB1) yielded 61 ²⁰⁶Pb/²³⁸U zircon ages of ~434–310 Ma (or weighted average age of 354.4 ± 5.9 Ma, n = 61, MSWD = 2.2; Fig. 5f). Fifty-three of these zircon analyses gave Th/U ratios <0.20, with only one Th/U ratio below 0.10, and 9 analyses gave ratios of 0.20– 0.37. These Th/U ratios are higher than the typical <0.1 ratio for metamorphic zircon (Rubatto 2017; Yakymchuk et al. 2018). The calc-silicate gneiss is within the stratified succession of the Nashoba terrane, is gradational with the rest of the Nashoba Formation and is similarly metamorphosed under upper amphibolite facies conditions. There is no field evidence to suggest that it is much younger than the other Nashoba terrane rocks. Furthermore, while Th/U ratios >0.1 are not typical for metamorphic zircon, they are not uncommon (Rubatto 2017; Yakymchuk et al. 2018). Yakymchuk et al. (2018) conclude that “igneous zircon rarely has Th/U < 0.1 and metamorphic zircon can have values ranging from <0.01 to >10.” For these reasons, the ages of zircons in the calc-silicate gneiss are interpreted as metamorphic, likely having formed in the presence of fluids.

34 Thus, we interpret metamorphic zircon in the Nashoba terrane to have grown between ~433 Ma and ~356 Ma, based on the weighted averages of ²⁰⁶Pb/²³⁸U ages. The Marlboro Formation, Shawsheen Gneiss, Nashoba Formation Gneiss and Tadmuck Brook Schist generally record older metamorphic zircon ages (weighted average age of 408.0 ± 8.8 Ma, n = 41, MSWD = 3.5) while zircon growth continued until somewhat later in the Nashoba Formation schist and especially the calc-silicate gneiss (weighted combined average age of 368.4 ± 4.0 Ma, n = 180, MSWD = 3.8).

35 The six formations analyzed from the Merrimack belt are presented below in a general NE to SW order (Fig. 2). The Eliot Formation consists of thinly layered grey to green phyllite, calcareous feldspathic metasiltstone and quartzite, quartz-mica schist, and calc-silicate gneiss (Lyons et al. 1997; Hussey et al. 2010, 2016). It has been interpreted as Silurian–Ordovician based on correlation with parts of the Casco Bay Group in Maine (Hussey et al. 2010). Sample KSELI from SE New Hampshire (Fig. 2) consists of generally thin-bedded alternations of tan-grey weathering, turbiditic calcareous quartzite and silvery dark phyllite (Fargo and Bothner 1995). Quartz veins were avoided during sample collection and removed prior to crushing. The sample yielded 119 concordant analyses. The weighted average of the ²⁰⁶Pb/²³⁸U ages of the four youngest detrital zircon grains is 409 ± 19 Ma (MSWD = 0.046). The major population peaks at ~446 Ma (Fig. 6a). Two grains are ~2.9 Ga and ~2.7 Ga and other ages spread between ~1.9 Ga and ~512 Ma.

36 The Kittery Formation consists of feldspathic and calcareous, ankeritic metaturbidite successions and minor fine-grained metasiltstone and feldspathic quartzite (Robinson and Goldsmith 1991; Bothner and Hussey 1999; Hussey et al. 1984, 2010, 2016). The fine-grained, thinly bedded to laminated rocks and thicker quartzitic beds contain cross lamination and other primary sedimentary structures (Hussey et al. 1984, 2010, 2016). Several paleocurrent indicators in the Kittery Formation in New Hampshire suggest a source to the southeast (Rickerich 1983, 1984; Hussey et al. 1984). Sample KSKTI (Fig. 2) was taken from an outcrop in SE New Hampshire, immediately across the Piscataqua River from the type locality in Maine (Katz 1917; Bothner and Hussey 1999) and is a tan weathering, grey to purplish, massive, fine-grained, feldspathic quartzite and quartzitic phyllite. It yielded 101 concordant analyses. The weighted average of ²⁰⁶Pb/²³⁸U ages of the four youngest detrital zircon grains is 413 ± 13 Ma (MSWD = 0.089). The main age group is ~461 Ma, and older populations are ~2.8 Ga, ~1.8– 1.0 Ga, and ~769 and 543 Ma (Fig, 6b).

37 The Berwick Formation is composed primarily of thin to thick tabular and lenticular beds of calcareous metasiltstone, biotite-rich metasiltstone, purplish biotite-quartz-feldspar granofels, and fine-grained metasandstone (Robinson and Goldsmith 1991; Fargo and Bothner 1995; Hussey et al. 2016). Some layers contain actinolite were metamorphosed to the greenschist facies, while at higher metamorphic grade, diopside, hornblende, and plagioclase are common (Robinson and Goldsmith 1991). Most of the Berwick Formation is in the garnet zone, although portions reach the sillimanite zone (Zen et al. 1983; Robinson and Goldsmith 1991; Hussey et al. 2016). Sample KSBWI, taken from NE Massachusetts (Fig. 2) consists of a fine-grained, light-grey to purplish-grey calcareous metasiltstone with beds a few cm to 1 m thick and interbeds of a dark-grey micaceous phyllite. It yielded 133 concordant analyses. The five youngest zircon grains provided a weighted average of ²⁰⁶Pb/²³⁸U ages of 409 ± 12 Ma (MSWD = 0.27). Zircon yielded a main age peak at 443 Ma, a few analyses at ~2.8–2.7 Ga and ~2.3 Ga and small populations between ~1.8 Ga and ~1.0 Ga and at ~634 Ma (Fig. 6d).

38 The Oakdale Formation is a thick unit composed of inter-layered brownish-grey to light-grey ankeritic metasiltstone, green-grey to purplish-grey impure quartzite, muscovite schist, and greenish-grey, grey, and dark brown calcareous phyllite (Grew 1973; Hepburn 1976; Peck 1976; Zen et al. 1983; Robinson and Goldsmith 1991). Sample KSOKI (Fig. 2) is from the type area in east-central Massachusetts (Emerson 1917; Hepburn 1976; Robinson and Goldsmith 1991) and consists of a grey weathering, light-grey quartz-rich phyllite and micaceous phyllite, interbedded with tan to grey-green weathering beds of metasiltstone and impure quartzite. It yielded 108 concordant zircon analyses. The youngest detrital zircon age population of the four youngest grains is 431 ± 12 Ma (weighted average ²⁰⁶Pb/²³⁸U age, MSWD = 0.69) and the main age peak is ~451 Ma (Fig. 6c). One analysis is ~2.7 Ga, and other ages show a continuous spread between ~2.0 Ga and the main age population.

39 The Worcester Formation, also called Worcester Phyllite, is a turbidite succession composed of grey and dark-grey carbonaceous slate, well foliated micaceous phyllite or schist containing andalusite, including chiastolite, inter-bedded with layers a few centimetres to one metre thick of impure quartzite, calc-silicate or metasiltstone (Grew 1973; Hepburn 1976; Peck 1976; Robinson and Goldsmith 1991). Sample KSWSI (Fig. 2) was taken largely from turbiditic sandy layers in a grey-weathering, light-grey, fine-grained turbidite with fissile phyllite interlayered with metasandstone and siltstone layers that range from a few cm to ~0.5 m thick. The sample yielded 138 concordant analyses. The weighted average of ²⁰⁶Pb/²³⁸U ages of the four youngest grains is 436 ± 9 Ma (MSWD = 0.89). The main age group is ~470 Ma. One grain is ~2.2 Ga, and other populations are between ~2.0 Ga and ~950 Ma, and ~800–470 Ma (Fig. 6e).

40 The Paxton Formation contains several sub-lithologies (Robinson and Goldsmith 1991). In the area of our study it is composed predominantly of grey-weathering, slabby, quartz-plagioclase-biotite granofels and well-layered purplish biotite granofels with calcic plagioclase (Robinson et al. 1982; Robinson and Goldsmith 1991). Sample KSPXIII from east-central Massachusetts (Fig. 2), is composed of a greenschist grade (garnet zone or lower at the locality sampled) laminated, slabby, quartz-rich, purple-grey granofels to schist. It yielded 113 concordant analyses. The youngest detrital zircon age population for the Paxton Formation is 426 ± 6 Ma (MSWD = 0.56) based on the weighted average of the ²⁰⁶Pb/²³⁸U ages of the ten youngest grains. The main peak is ~463 Ma with a smaller peak at ~431 Ma. Two analyses are ~2.6 Ga and most other analyses are between ~1.8–0.9 Ga with a few grains between ~700 and ~528 Ma (Fig. 6f).

41 Youngest zircon grains of selected samples from the Merrimack belt were re-dated using CA-TIMS methods (Table 1; Fig. 7), but as with the Nashoba terrane samples above, not all survived the processing. All CA-TIMS data for the Merrimack belt were within error of, or older (some much older, see below) than the LA-ICPMS data (Tables 1, A2). The youngest CA-TIMS ²⁰⁶Pb/²³⁸U ages from five grains from the Kittery, Paxton and Oakdale formations ranged from 434 Ma to 425 Ma.

43 Based on the LA-ICPMS results, the youngest detrital zircon population of all samples of the Nashoba terrane combined is 466 ± 13 Ma (MSWD = 0.12, n = 11). Given this, and taking the errors into account, the maximum ages of the Nashoba terrane metasedimentary rocks could range from Early to Late Ordovician. If the higher precision but single CA-TIMS age of 465.9 ± 0.7 Ma (Table 1; Fig. 7) on grain mr02a06 of the Nashoba Formation gneiss (LA-ICPMS age of 480 ± 13 Ma) is representative, portions of the Nashoba Formation could have a maximum depositional age of early Middle Ordovician. Similar results were reported by Walsh et al. (2021) for two samples from the Nashoba Formation in southern Massachusetts and its correlative in Connecticut, the Tatnic Hill Formation, where detrital zircons gave an age of deposition between the ~485 Ma age of the youngest detrital zircon, and the ~435 Ma age of metamorphism. Because the 515 Ma granitic Grafton Gneiss cuts some of the mafic metaigneous rocks of the Marlboro Formation, and a ~500 Ma feldspathic granofels structurally overlies these (Walsh et al. 2009, 2011a, b, 2021), at least part of the Marlboro Formation is late Cambrian or older. The scarcity of detrital zircon in the metasedimentary rocks of the Marlboro Formation may indicate that they are primarily derived from the local, largely mafic, volcanic rocks of the Marlboro Formation. Thus, at least parts of the Marlboro Formation, and likely portions of the other Nashoba terrane units, are Cambrian, while other units, or parts of them, are Ordovician or younger.

44 The Assabet River fault zone (Fig. 2) is a major fault zone within the Nashoba terrane, separating the Nashoba Formation and Tadmuck Brook Schist to the northwest from the Shawsheen Gneiss and Marlboro Formation to the southeast (Bell and Alvord 1976; Hepburn and DiNitto 1978; Zen et al. 1983; Goldsmith 1991a, b; Kopera et al. 2006). To determine the significance of this fault zone, samples from the northwestern and southeastern Nashoba terrane were plotted separately in Figs. 8a and b. Samples from both sides show very similar patterns with a significant ~560–540 Ma detrital zircon age population and minor Mesoproterozoic and Paleoproterozoic populations. Archean grains are rare in both. The similarity of these data indicates that there is no significant difference in detrital age populations across the Assabet River fault zone and that it thus does not represent a terrane boundary. The combined Nashoba terrane data are plotted in Fig. 8c. The similarity of the detrital zircon age distribution in the Tadmuck Brook schist (Fig. 4e) to that of the other units of the Nashoba terrane, and its difference from the distributions in the Merrimack belt (Figs. 6, 8f) indicates that it is part of the Nashoba terrane.

45 The Nashoba terrane has been interpreted as a primitive arc/back-arc complex (Hepburn et al. 1995; Kay et al. 2009, 2011, 2017; Walsh et al. 2011b, 2015, 2021). Archean and Proterozoic detrital zircon and Mesoproterozoic depleted mantle Sm/Nd model ages (Kay et al. 2017) suggest that this complex formed on older crustal material. Paleoproterozoic and Mesoproterozoic detrital zircon in the metasedimentary rock samples support this, and similar youngest zircon age populations and provenance in all four units of the Nashoba terrane indicate that they all formed as part of the same early Paleozoic arc/back-arc complex.

46 Both Avalonia and at least part of Ganderia experienced active ~650–590 Ma magmatism (e.g., Murphy et al. 2004; Fyffe et al. 2009; Satkoski et al. 2010; Thompson et al. 1996, 2007, 2010; Pollock et al. 2009; van Staal et al. 2009, 2020; Barr et al. 2012, 2019). For Ganderia, this abundant arc magmatism continued until the Cambrian, whereas in Avalonia it ceased or greatly decreased by ~565 Ma as Avalonia became a stable platform (Barr et al. 2003a; Murphy et al. 2004; Samson et al. 2005; Rogers et al. 2006; Hibbard et al. 2007; van Staal et al. 2009, 2020; van Staal and Barr 2012). Detrital zircon age populations from the metasedimentary units of the Nashoba terrane are generally consistent with those from the Ediacaran–early Silurian Ganderian rocks from Newfoundland, New Brunswick and east-central Maine (Fig. 9b; e.g., White and Barr 1996; Barr et al. 2003b, 2019; Valverde-Vaquero et al. 2006b, Pollock et al. 2007, 2009; Schulz et al. 2008; Fyffe et al. 2009; Ludman et al. 2018). The large ~540 Ma detrital zircon age population in all units, and the smaller ~640 Ma age population with minor Mesoproterozoic and Paleoproterozoic populations suggest a Ganderian affinity (Barr et al. 2014; Rogers et al. 2006; Fyffe et al. 2009; Pollock et al. 2007, 2009). It is less consistent with Avalonia, which is characterized by a large ~640 Ma population and a smaller ~540 Ma population (Fig. 9c; cf. Keppie et al. 1998; Thompson and Bowring 2000; Barr et al. 2003b, 2012; Murphy et al. 2004; Pollock et al. 2009; Satkoski et al. 2010; Thompson et al. 2014). Strongly negative ε_{Nd (500)} values and Paleoproterozoic model ages for the metasedimentary rocks in the Nashoba terrane (Kay et al. 2017) are also characteristic for Ganderian sedimentary rocks (Samson et al. 2000; Schofield and D’Lemos 2000; Rogers et al. 2006; Pollock et al. 2012) and not for typical Avalonian crust (Nance and Murphy 1996; Barr et al. 1998; Hibbard et al. 2007 Thompson et al. 2012). We therefore interpret the Nashoba terrane as having Ganderian affinity.

47 Ganderia is interpreted to have formed along the Iapetus ocean-facing Amazonian margin of Gondwana in the Neoproterozoic and rifted away in the middle Cambrian, opening the Rheic ocean behind it (Figs. 10a, b) (e.g., van Staal et al. 1996, 2009, 2012, 2020; Fyffe et al. 2009, 2011; Schulz et al. 2008; Pollock et al. 2012; Barr et al. 2014; van Staal and Barr 2012; Willner et al. 2014). Prior to rifting a ~550–528 Ma extensive arc complex developed on the Ganderian margin (Fyffe et al. 2011; van Staal and Barr 2012). It formed the basement to the ~515–485 Ma Penobscot arc/back-arc complex that developed on the leading edge of Ganderia as it moved into the Iapetus ocean (Fig. 10b) (Valverde-Vaquero et al. 2006a; Rogers et al. 2006; Murphy et al. 2006; van Staal et al. 2009; Fyffe et al. 2009, 2011; Zagorevski et al. 2007a, b, 2010; van Staal and Barr 2012). At ~486–478 Ma, the Penobscot arc and back-arc fragments consolidated outboard of Laurentia during the Penobscot orogeny (Fig. 10c) (Zagorevski et al. 2007b, 2010; van Staal et al. 2009; Johnson et al. 2009, 2012; van Staal and Barr 2012). Penobscot arc/ back-arc rocks are found, south of Newfoundland, in the Brookville, New River and Annidale belts in New Brunswick, the Bras d’Or belt in Cape Breton Island, and the Ells-worth Formation and St. Croix terrane in eastern Maine and share many characteristics with the Nashoba terrane (White et al. 1994; Barr et al. 1998; Schulz et al. 2008; Johnson et al. 2009, 2012; Fyffe et al. 2011; van Staal and Barr 2012).

48 Shortly after the Penobscot orogeny, a second expansive arc/back-arc complex (Popelogan-Victoria cycle) formed from ~478–453 Ma on the remnants of the now consolidated Penobscot arc/back-arc deposits (Fig. 10d) (e.g., van Staal 1994; van Staal et al. 1998, 2016; Zagorevski et al. 2010; van Staal and Barr 2012; Wilson et al. 2017) and opened up the expansive Tetagouche and Exploits back-arc basins in Atlantic Canada – New England and in Newfoundland, respectively. These basins divided the Ganderian terrane into a separate active leading margin and a passive trailing margin (where the Nashoba terrane is located). The leading margin collided with Laurentia at about 455–450 Ma (Taconic phase 3 orogeny of van Staal et al. 2009). During the next ~20–25 myr the Tetagouche-Exploits basin gradually closed by west-directed subduction under Laurentia (Salinic orogenic cycle) (Figs. 10e, f) until the trailing margin of Ganderia collided with Laurentia in the terminal 426–420 Ma Salinic orogeny (Fig. 10g) (e.g., van Staal et al. 1998, 2003, 2008, 2012, 2016; Valverde-Vaquero et al. 2006a; Zagorevski et al. 2008; Fyffe et al. 2011; Zagorevski and van Staal 2011; van Staal and Barr 2012).

49 In the Nashoba terrane, igneous rocks of the ~540–500 Ma Marlboro Formation, the 515 Ma Grafton Gneiss and the 499 Ma Fish Brook Gneiss (Hepburn et al. 1995; Acaster and Bickford 1999; Walsh et al. 2011a, 2015, 2021) originated during the time of formation of the Cambrian Penobscot arc/back-arc complex (Figs. 10a, b) or its Ediacaran–Cambrian basement. An abundance of detrital zircon of that age in metasedimentary units of the Nashoba terrane reflects internal sources within the Penobscot arc/back-arc complex and its basement. Older zircon grains in the Nashoba terrane were eroded into the sedimentary basins from either Gondwanan (Amazonian) sources prior to the Cambrian rifting and drifting of Ganderia, or originated from reworked basement rocks beneath the Nashoba terrane. Such Proterozoic basement is exposed beneath Penobscot-cycle Ganderian rocks in the Brookville and New River terranes and on Grand Manan Island in New Brunswick (Johnson et al. 2009, 2012; Fyffe et al. 2009, 2011; Pollock et al. 2012; van Staal and Barr 2012; Fyffe 2014; Barr et al. 2019), the Bras d‘Or terrane in Cape Breton Island (Barr et al. 1998; van Staal and Barr 2012) and the Islesboro block in coastal Maine (Reusch et al. 2018, 2021). The combined LA-ICPMS youngest detrital zircon age population of all samples from the Nashoba terrane (466 ± 13 Ma) and the single CA-TIMS age (465.9 ± 0.7 Ma) suggest that Ordovician metasedimentary rocks also exist in the terrane and may have formed during the Popelogan tectonic cycle (Fig. 10d).

50 The Nashoba terrane contains abundant metamorphic zircon grains. These show a significant spread of weighted average ages between ~433 Ma and ~356 Ma (Fig. 5) and primarily indicate zircon growth during the Acadian orogeny. Younger zircon ages, while showing scatter and large uncertainties, likely reflect growth and/or recrystallization that extended into the Neoacadian and Alleghanian orogenies. Metamorphic zircon formed prior to ~425 Ma may have originated during the earlier Salinic orogeny. Our ages are consistent with published ages of metamorphism in the Nashoba terrane using various methods (Hepburn et al. 1995; Wintsch et al. 2007; Stroud et al. 2009; Walsh et al. 2013a, 2015, 2021; Buchanan et al. 2016a, b; Severson 2020). Because the precision of our ages is not higher than those reported previously, and because of consistency between the various data sets, we will not discuss the ages of metamorphism further.

51 The youngest LA-ICPMS detrital zircon age population in the Tower Hill Formation is 513 ± 15 Ma (3 grains), considerably older than any of the six formations sampled from the Merrimack belt (Figs. 6, 8f). A single grain from this unit (not included in the grouping at 513 Ma) has a LA-ICPMS ²⁰⁶Pb/²³⁸U age of 463 ± 9 Ma, which is within error of the youngest age population of the Nashoba terrane. However, the ~513 Ma population is older than any youngest populations in the metasedimentary rocks of the Nashoba terrane and other aspects of the zircon population in the Tower Hill Formation, such as the 630 Ma main age peak, are different from any of the other samples (Figs. 4, 8c). Of particular interest is the ~950–750 Ma age population that is not characteristic for known Avalonian or Ganderian rocks.

52 The 513 ± 15 Ma youngest age population of the Tower Hill Formation is similar to the maximum ages of deposition for Cambrian rocks found in several Ganderian terranes in New England and New Brunswick. These include the Ellsworth Formation of the Ellsworth terrane, the Calais Formation of the St. Croix terrane, and the Baskahegan Lake Formation of the Miramichi terrane (Fyffe et al. 2009). However, these formations do not show a large peak at ~630 Ma. Basement rocks (Martinon Fm.) below the Ganderian Brookville terrane and on Grand Manan Island in New Brunswick (Flagg Cove and Long Pond Bay Formations, Fyffe et al. 2009; Fyffe 2014; Barr et al. 2019) do show a Neoproterozoic peak of approximately this age. Rocks with both Cambrian youngest detrital zircon age populations and ~630 Ma peaks occur in some samples in the Moretown, Albee and Dead River formations of the Moretown terrane in western Massachusetts, Vermont, New Hampshire and western Maine (Karabinos et al. 2017) and the Ganderian Kedears Hill greywacke of North Islesboro, Maine (Reusch et al. 2018, 2021). However, because none of their detrital zircon signatures is similar to the Tower Hill formation, no specific correlations can be made. The detrital zircon ages do indicate that the Tower Hill Formation is not a part of the Merrimack belt stratigraphy and that the boundary between the Nashoba terrane and Merrimack belt is complex and likely incorporates exotic fault-bounded slices.

53 The LA-ICPMS youngest detrital zircon populations for individual formations of the Merrimack belt in southeastern New Hampshire and Massachusetts range between 436 ± 9 Ma (Worcester Formation) and 409 ± 12 Ma (Berwick Formation). The CA-TIMS ²⁰⁶Pb/²³⁸U ages for the five youngest detrital zircon grains from the Merrimack belt samples are between 434 and 425 Ma. Three CA-TIMS ages from the Merrimack belt are within error of LA-ICPMS analyses that were used as part of the youngest populations (Tables 1, A2). These include two grains from the Oakdale Formation (mr16b23, LA-ICPMS: 409 ± 17 Ma, CA-TIMS: 425.53 ± 0.64 Ma and mr16a24 LA-ICPMS: 436 ± 8 Ma, CA-TIMS: 433.70 ± 0.41 Ma respectively) and grain au26a47 of the Paxton Formation (LA-ICPMS: 412 ± 18 Ma, CA-TIMS: 429.16 ± 0.82 Ma). The remainder of the CA-TIMS ages on grains from the youngest groups of zircons are older than the LA-ICPMS ages, up to ~30 myr for a grain from the Worcester Formation with a CA-TIMS age of 486 Ma (Table 1). This may reflect either unaccounted for Pb-loss in the LA-ICPMS analyses or the presence of older cores that survived chemical abrasion better than overgrowths prior to analysis. These older CA-TIMS ages are not interpreted as reflecting the ages of the youngest zircon populations.

54 The CA-TIMS ages for the three youngest detrital zircons in the Oakdale and Paxton formations (434–426 Ma) are consistent with the 429 ± 5 Ma LA-ICPMS youngest detrital zircon age population (MSWD = 0.82, n = 18) for the three formations in the central Massachusetts portion of the belt (Worcester, Oakdale, and Paxton formations). The Worcester Formation is intruded by a 424 ± 3 Ma phase of the Ayer granodiorite at Eddy Pond, near Auburn, Massachusetts (U–Pb secondary ion mass spectrometry (SIMS) zircon data; Walsh et al. 2013a; Walsh and Merschat 2015), and the Oakdale Formation is intruded by the 420.1 ± 0.1 Ma Clinton facies of the Ayer granodiorite (U–Pb CA-TIMS zircon data; Charnock 2015). Therefore, the ages of these two formations are constrained between ~434 Ma and ~420 Ma (late Llandovery to Pridoli).

55 The combined youngest detrital zircon age population for the Kittery, Eliot and Berwick formations in the northeastern part of the sample area is 411 ± 8 Ma (MSWD = 0.14, n = 13). The LA-ICPMS weighted average of ²⁰⁶Pb/²³⁸U ages of the four youngest detrital zircon grains from the Eliot Formation is 409 ± 19 (MSWD = 0.046), of the four youngest grains from the Kittery Formation is 413 ± 13 Ma (MSWD = 0.089), and of the five youngest detrital zircon grains from the Berwick Formation is 409 ± 12 Ma (MSWD = 0.27). Two grains from the Kittery Formation youngest population yielded CA-TIMS ages of 432 Ma and 434 Ma (Table 1). The previously folded and regionally metamorphosed Kittery and Eliot formations are intruded in SE New Hampshire by the 418 ± 1 Ma Newburyport quartz diorite and the 407.4 ± 0.5 Ma Exeter diorite (U–Pb ID-TIMS and CA-TIMS zircon data; Bothner et al. 1993, 2009; Hussey and Bothner 1993; Fargo and Bothner 1995; Wintsch et al. 2007). The Berwick Formation was intruded by the 407 ± 4 Ma Devens Long Pond facies of the Aver Granite in Westford, Massachusetts (U–Pb SHRIMP, Walsh et al. 2013b, 2021). This indicates that at least the Kittery and Eliot formations must be older than 418 Ma and that the LA-ICPMS youngest average detrital zircon age for these formations is inconsistently young.

56 Based on the LA-ICPMS and CA-TIMS data and the dated intrusive rocks, the Merrimack belt units in Massachusetts and SE NH are interpreted to have been deposited between ~434 Ma and ~420 Ma. However, since the Berwick and Paxton formations are not intruded by any pre-407 Ma dated igneous rocks in the area their minimum age limit is not well constrained, and they could be somewhat younger.

57 In Massachusetts and SE New Hampshire, the Merrimack belt rocks have ages of deposition between ~434 and ~420 Ma, large Middle to Late Ordovician zircon populations, smaller ~500–800 Ma populations, significant Mesoproterozoic populations, and scattered Archean grains. The Mesoproterozoic populations are much more prominent in the Kittery, Berwick and Paxton formations than in the Eliot, Oakdale, and Worcester formations (Figs. 6, 8d, e). While both Laurentia and Gondwana are potential viable sources for Mesoproterozoic zircon (e.g., Murphy et al. 2004; Fyffe et al. 2009; Pollock et al. 2009; Barr et al. 2019; Figs. 9a, b), Laurentia is a better source for ~1.2–1.0 Ga zircon (e.g., Cawood and Nemchin 2001; Hibbard et al. 2007; Waldron et al. 2014, 2018; Dokken et al. 2018; Severson 2020; Kuiper and Hepburn 2021). Furthermore, ~2.2–2.0 Ga zircons are largely absent in Laurentia, while Gondwana provides some zircon of this age, and Neoproterozoic zircon is characteristic for Gondwana, but not for Laurentia (Figs. 9a, b; Pollock et al. 2007; Waldron et al. 2009, 2014; Severson 2020). In the Merrimack belt samples the detrital zircon signatures of the Kittery, Berwick and Paxton formations show a Mesoproterozoic population, including a prominent 1.2–1.0 Ga population, no ~2.2–2.0 Ga grains and minor Neoproterozoic zircon, indicating evidence of Laurentian input. In contrast, the Eliot, Oakdale, and Worcester formations have a much smaller Mesoproterozoic population, three ~2.2–2.0 Ga grains and abundant Neoproterozoic zircon (Figs. 6, 8d, e), indicating Gondwanan but not Laurentian input.

58 In south-central Maine, Cartwright et al. (2019) concluded that detrital zircon in the older rocks of the Fredericton trough (Llandovery Appleton Ridge Formation and Ghent Phyllite; Figs. 1, 3) have a Gondwanan source and no evidence of Laurentian sediment input, similar to the Llandovery Digdeguash Formation of the Kingsclear Group, southeast of the Fredericton fault (northern extent of the Norumbega fault system; Fig. 1) in southern New Brunswick (Dokken et al. 2018). However, the overlying Wenlock–Ludlow Flume Ridge Formation in the eastern Fredericton trough in New Brunswick and the Llandovery Hayes Brook and Wenlock–Ludlow Burtts Corner formations northwest of the Fredericton fault contain detrital zircon with Mesoproterozoic Laurentian input (Dokken et al. 2018). Likewise, rocks deposited in the Central Maine basin (Vassalboro Group) contain mixed Gondwanan and Laurentian detrital zircon signatures with Mesoproterozoic Laurentian input to its eastern structural margin in central Maine (Cartwright et al. 2019). Thus, rocks that contain a Laurentian detrital zircon signature were either deposited in a more westerly location, i.e., in the Central Maine basin or western portion of the Fredericton trough, or else are younger deposits in the eastern Fredericton trough.

59 In the Merrimack belt in Massachusetts and SE New Hampshire the Worcester, Oakdale and Eliot formations, with minor Laurentian Mesoproterozoic detrital zircon likely correlate with the early Silurian deposits in the Fredericton trough in Maine and New Brunswick. The mixed detrital zircon source of the Kittery Formation with the presence of Laurentian detrital Mesoproterozoic input is consistent with the younger Silurian rocks in the Fredericton trough in New Brunswick (Dokken 2018) and the interpretation of Hussey et al. (2010) and Wintsch et al. (2007) that it represents the youngest deposit in the Merrimack Group in SE Maine and New Hampshire. The Paxton and Berwick formations (Figs. 6d, f) also have mixed Gondwanan and Laurentian detrital zircon signatures. Hussey et al. (2010, 2016; cf. Wintsch et al. 2007) place the Berwick Formation stratigraphically above the Kittery and Eliot formations in SE Maine and New Hampshire and interpret it as part of the Central Maine basin. The similar lithologies of the Paxton and Berwick formations, their similar detrital zircon signatures including Laurentian 1.2–1.0 detrital zircons, and the fact that neither is known to be crosscut by late Silurian–Early Devonian plutons support their potential correlation (Robinson and Goldsmith 1991). Thus, the Paxton and Berwick formations are interpreted to be the youngest rocks of those sampled in the Merrimack belt in SE New Hampshire and Massachusetts, consistent with interpretations by Hussey et al. (2010, 2016; cf. Wintsch et al. 2007) and further suggesting that the Paxton Formation was likely deposited in the Central Maine basin or its equivalent in Massachusetts.

60 Dokken et al. (2018) and Cartwright et al. (2019) propose that Laurentian detrital material only reached the eastern Fredericton trough by the middle Silurian and that a barrier existed between Laurentian margin and Gondwanan sources prior to this time (see also Ludman et al. 2018). A similar barrier may have existed during the deposition of the Worcester, Oakdale and Eliot formations, which was not present during deposition of the Kittery, Paxton and Berwick formations.

61 Our detrital zircon age populations for the Kittery and Berwick formations are generally similar to those based on the SHRIMP U–Pb analysis of Wintsch et al. (2007; Figs. 9d, e). Although minor differences exist in the sizes of the populations, these may be a result of differences in the number of analyses. Wintsch et al. (2007) correlated the Berwick Formation in Maine with the Hebron Formation in Connecticut. The Hebron Formation has a zircon age distribution (Fig. 9f) similar to the Berwick Formation, although the Hebron Formation contains a relatively larger ~1.2–1.0 Ga Laurentian population. This may be a result of local provenance differences, e.g., a higher local input from a Grenville source. Alternatively, it may indicate that the Hebron Formation, where sampled, is somewhat younger than the Berwick Formation, resulting in an increased Laurentian signal. The similarity of the detrital zircon distributions in the Berwick, Paxton and Hebron formations to those of samples from the Central Maine basin (Fig. 9g; Bradley and O’Sullivan 2017) indicate that they all were likely deposited in the Central Maine basin or its equivalent in Massachusetts and Connecticut.

62 The Merrimack belt in Massachusetts and SE New Hampshire, the Merribuckfred basin and eastern portion of the Central Maine basin in Maine and New Brunswick all show a mix of Gondwanan and Laurentian sources. Here we briefly summarize the spatial and temporal differences in sources for the Merribuckfred and Central Maine basins, from north to south, so as to better identify the sources for Merrimack belt. Ludman et al. (2018, 2019) interpret the rocks in the Fredericton trough in eastern Maine to have originated largely from Gondwanan sources, either from emergent basement highlands underlain by Ganderian rocks such as in the Miramichi terrane, or by Wenlock–Ludlow time, from arc sources to the east as Avalonia approached and docked with the trailing edge of Ganderia. Such arc sources are present in the Coastal Volcanic belt in eastern Maine and the Kingston belt in New Brunswick (Fig. 1; Gates and Moench 1981; Seaman et al. 1999, 2019; Barr et al. 2002; Van Wagoner et al. 2002; Piñán-Llamas and Hepburn 2013; Ludman et al. 2017, 2018). In southern coastal Maine, Hussey et al. (2010) interpreted the Llandovery Bucksport and Appleton Ridge formations in the Fredericton trough to have Gondwanan sources and to have been deposited between a western Ganderian Ordovician arc (present in the Casco Bay and Falmouth-Brunswick Groups; Fig. 1) and the eastern trailing margin of Ganderia. Similarly, late Ordovician (?) to early Silurian deposits in the eastern Fredericton trough in New Brunswick (Dokken et al. 2018) and in south- central Maine (Cartwright et al. 2019) show only Gondwanan sources. However, younger deposits in the Kingsclear Group in the eastern Fredericton trough in southern New Brunswick and those in the western Fredericton trough (Fig. 3) include Laurentian detrital zircon material (Dokken et al. 2018).

63 The majority of the zircon in the Merrimack belt in Massachusetts and SE New Hampshire (Figs. 6, 8f) has Gondwanan affinity and most formed during the ~478–453 Ma Popelogan-Victoria volcanic arc/back-arc cycle that included the opening of the broad Tetagouche-Exploits back-arc basin (Fig. 10d; e.g., Valverde-Vaquero et al. 2006a; Fyffe et al. 2009; Johnson et al. 2009; Zagorevski et al. 2007a, 2010; van Staal and Barr 2012; Wilson et al. 2015; van Staal et al. 2016). Input of zircons into the Merrimack belt during the existence of the Tetagouche-Exploits back-arc basin likely came from erosion of emergent local highlands underlain by Ganderian Ordovician rocks, such as in the Miramichi and St. Croix terranes in eastern Maine and New Brunswick (Fig. 10e; Fyffe 1995; Fyffe et al. 2009, 2011; Ludman et al. 2017, 2018) and in the Casco Bay and Falmouth-Brunswick Groups in south-central and coastal southern Maine (West et al. 2003, 2004, 2008; Gerbi and West 2007; Hussey 2010, West and Hussey 2016, Hussey and West 2018). Input of Laurentian-derived zircons to the eastern portions of the trailing margin of Ganderia during this time was probably minor as the Tetagouche-Exploits basin and perhaps these highlands formed a barrier to Laurentian zircon transport (Fig. 10e), but was more prominent to the west in the Central Maine basin (Bradley and O’Sullivan 2017; Cartwright et al. 2019).

64 The boundary between the Merribuckfred and Central Maine basins in New England has been interpreted as the line of closure of the Tetagouche-Exploits basin (the Dog Bay line in Newfoundland) during the Salinic orogeny (Reusch and van Staal 2012). During the closure of this basin, it is likely that the younger Silurian deposits in the Merrimack belt (Kittery, Berwick and Paxton formations) received increasing amounts of Laurentian detritus (Fig. 10f). By the middle Silurian, when the Tetagouche-Exploits basin was closing, the oceanic crust between the approaching Avalonia and Ganderia was subducting below the trailing margin of Ganderia (Fig. 10f). The Silurian volcanic arc formed at this time is represented in Massachusets by the Newbury Volcanic Complex (Shride 1976a, b; Zen et al. 1983) along the northeastern boundary of the Nashoba terrane (Fig. 2). Avalonia collided by ~425 Ma (Fig. 10g; e.g., Fyffe et al. 1999; Bradley et al. 2000; van Staal et al. 2009; van Staal and Barr 2012; Piñán-Llamas and Hepburn 2013; Wilson et al. 2017). Paleocurrent directions in the Kittery Formation in the Merrimack belt suggest that some of the youngest detritus was sourced east of the belt (Rickerich 1983, 1984; Hussey et al. 1984). Thus, middle to late Silurian detrital zircon in the Merrimack belt could have been derived from the Salinic orogen and Laurentian margin to the west and/or from arc volcanism and orogenesis related to the approach and collision of Avalonia to the east (Figs. 10f, g). As the Acadian orogenic front advanced westward, earlier deposits to the east were uplifted and likely provided reworked detritus for younger, more westerly, deposits in the Central Maine basin (Bradley et al. 2000; Hussey et al. 2010; Bradley and O’Sullivan 2017; Ludman et al. 2018).