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Journal of Petrology | Volume 43 | Number 11 | Pages 2097-2120 | 2002
© Oxford University Press 2002
Metamorphic, Thermal, and Tectonic Evolution of Central New England
1DEPARTMENT OF EARTH AND ENVIRONMENTAL SCIENCES, RENSSELAER POLYTECHNIC INSTITUTE, TROY, NY 12180, USA
2DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF SOUTH CAROLINA, COLUMBIA, SC 29208, USA
3DEPARTMENT OF GEOLOGY, AMHERST COLLEGE, AMHERST, MA 01002, USA
4SCIENCE DIVISION, JEFFERSON COMMUNITY COLLEGE, WATERTOWN, NY 13601, USA
Received April 19, 2001; Revised typescript accepted May 8, 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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A new, detailed tectonic model is presented for the Acadian orogenic belt of central New England (Vermont and New Hampshire) that accounts for a wide range of petrological and structural observations. Three belts are considered: the Eastern Vermont, Merrimack, and intervening Bronson Hill belts. Specific observations in eastern Vermont that are accounted for in the model include the following. P–T paths are clockwise with maximum pressures near the Athens, Chester, and Strafford domes of 8–11 kbar, but with maximum pressures decreasing to 3–5 kbar at the boundary with the Bronson Hill belt. Differential exhumation of the Vermont domes relative to the rocks in easternmost Vermont is required by the recorded differences in maximum pressure (5–6 kbar; 15–20 km) and the present-day geographical separation (7–10 km). Specific observations in New Hampshire that are explained include the following. P–T paths in the Merrimack belt are counter-clockwise with maximum pressures of 4–5 kbar and are related to high regional heat flow and heat transfer by early Acadian plutons. P–T paths in the Bronson Hill belt are intimately associated with structural position. An early contact metamorphism is evidenced in the Skitchewaug and Fall Mountain nappes near contacts with the early Acadian Bethlehem gneiss (
400–410 Ma). Peak metamorphic temperature rises upwards in the nappe sequence (an inverted metamorphic sequence) whereas peak pressures decrease. Near-simultaneous intrusion of the Bethlehem gneiss and Kinsman quartz monzonite is required to account for the low-P, high-T metamorphism observed in the Chesham Pond and Fall Mountain nappes. The dominant schistosity, which is related to isoclinal folding, postdates early contact metamorphism in the Fall Mountain and Skitchewaug nappes, and pre-dates peak metamorphism and isothermal loading in the Fall Mountain, Skitchewaug and Big Staurolite nappes. Reactivation of this fabric during thrusting is recorded in some rocks of the Big Staurolite nappe by rotated garnets that grew during near-isothermal loading. Only the sillimanite isograd crosses the Fall Mountain–Skitchewaug nappe boundary. Metamorphic breaks across the Skitchewaug–Big Staurolite nappe boundary, at the base of the Big Staurolite nappe, and at the margin of the Keene and Alstead domes require post-metamorphic thrusting when P–T conditions were in the greenschist facies. These observations can be explained by a relatively simple model involving in-sequence thrusting from east to west commencing in central New Hampshire at 400–410 Ma. Preservation of the low-grade belt along the Vermont–New Hampshire border requires that crustal thickening in Vermont was not caused by emplacement of New Hampshire nappes onto eastern Vermont and that the nappes of western New Hampshire had time to cool before final juxtaposition against the low-grade belt. Cooling ages constrain this final juxtaposition to have occurred in the Carboniferous, suggesting that the Acadian was a prolonged event spanning as much as 100 Myr. KEY WORDS: New England; Vermont; New Hampshire; Acadian; inverted metamorphism; P–T paths
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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The Acadian of central New England (Fig. 1) consists of three belts: the eastern Vermont, Bronson Hill, and the Merrimack belts. The traditional view (e.g. White & Jahns, 1950; Doll et al., 1961; Rosenfeld, 1968; Thompson & Norton, 1968; Thompson et al., 1968; Naylor, 1969, 1971; Robinson & Hall, 1980; Hatch et al., 1983; Robinson et al., 1991) of the tectonic evolution of this region is one in which two sedimentary basins, the Connecticut Valley trough of eastern Vermont and the Merrimack trough of central New Hampshire, coexisted beginning in the Silurian and extending through the late Early Devonian (Emsian). The Bronson Hill terrane, situated between the two basins, is believed to be the remnant of the Taconian arc responsible for the Ordovician Taconic orogeny (e.g. Tucker & Robinson, 1990). Both basins were deformed and metamorphosed during the Middle Devonian Acadian orogeny, which is believed to have been caused by collision of the Avalon terrane from the east. Large-scale, west-directed recumbent folding and thrusting are characteristic of all three terranes, and the pattern of isograds (Fig. 2) suggests a continuous metamorphic gradient across central New England.
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On closer inspection, the metamorphic histories of the three belts contrast markedly (Figs 1 and 2). Rocks from eastern Vermont were metamorphosed to the staurolite–kyanite zone (Fig. 2) and experienced dominantly clockwise P–T paths (Fig. 1; Armstrong et al., 1992; Menard & Spear, 1994; Armstrong & Tracy, 2000) reaching pressures of 8–11 kbar on the flanks of the Chester and Athens domes (Kohn & Spear, 1990; Ratcliffe et al., 1992; Kohn & Valley, 1994; Menard & Spear, 1994). In contrast, the Merrimack belt is characterized by low-pressure, high-temperature metamorphism (Fig. 2) and dominantly counter-clockwise P–T paths (Fig. 1 inset). Metamorphic parageneses in the Bronson Hill belt are intimately related to a series of nappes in which rocks of higher metamorphic grade are found in higher structural levels. This inverted metamorphic sequence in the Bronson Hill belt has been recognized for several decades (Chapman, 1953; Thompson et al., 1968), and recent findings (e.g. Spear, 1992, 1993; Kohn et al., 1997) reveal that juxtaposition of higher-grade rocks upon lower-grade rocks could not have occurred during the peak of metamorphism, but must have occurred following substantial cooling of the high-grade rocks (Kohn et al., 1997). Differences in the P–T evolution of the three Acadian belts continue through their cooling histories as evidenced by thermochronology studies (e.g. Harrison et al., 1989; Spear & Harrison, 1989).
The purpose of this paper is to present a synthesis of the metamorphic history of rocks from central New England and to show the metamorphism relates to the tectonic assembly of the terrane. Key to the interpretation of the tectonic assembly of this region is the relationship between the P–T evolution and the fabric development in each structural level. The P–T histories have been deduced from the metamorphic recrystallization using a variety of methods including thermobarometry, garnet zoning analysis (e.g. Spear & Selverstone, 1983), pseudomorph textures, and comparison of inferred reactions with petrogenetic grids. The tectonic fabrics are related to major deformation events that involve isoclinal folding and transport. Therefore, the relative timing of a specific part of a P–T path to a fabric with a known tectonic significance reveals the depth and thermal conditions of the crust when the deformation occurred. For example, several nappes in western New Hampshire experienced P–T paths that include a segment of isothermal or near-isothermal loading. The loading is interpreted to have occurred in response to emplacement of higher-level nappes. Therefore, the relationship between the dominant fabric in the lower nappe and the metamorphic recrystallization that records the change in pressure reveals when the lower nappe was deformed (the fabric-producing event) relative to the emplacement of the higher-level nappes (the loading event).
The paper focuses on a transect at the approximate latitude of Fall Mountain, New Hampshire (see Fig. 3) starting with the lowest Acadian structural levels of eastern Vermont and working eastwards and structurally upwards. Along each part of the transect, the P–T history, fabric development, and their relationship will be stressed. A summary of mineral assemblages, compositions, and peak P–T conditions for all samples referenced in this study is presented in Table 1. Particular emphasis is placed on the evolution of western New Hampshire (the Bronson Hill terrane), because it forms the transition zone between the fundamentally distinct eastern and western belts. Important variations also occur along strike to the north in western New Hampshire, and these will be summarized where appropriate. Finally, the thermal, baric, and structural evolution of the region based on these data will be integrated into a synthesis that constrains the timing of the juxtaposition of the distinct tectonic slices.
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EASTERN VERMONT P–T EVOLUTION |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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The eastern Vermont belt includes the Connecticut Valley synclinorium, a sequence of Silurian to Devonian metasedimentary and metaigneous rocks, and a sequence of Cambrian to Ordovician rocks that flank a series of major north–south-trending domes (Athens, Chester, Strafford), in which are exposed higher-grade and, in the case of the Athens and Chester domes, Proterozoic age rocks (e.g. White & Jahns, 1950; Doll et al., 1961; Rosenfeld, 1968) (Fig. 3).
Two major periods of deformation affected the rocks of the eastern Vermont belt during the Acadian, an earlier nappe stage and a later dome stage (e.g. Rosenfeld, 1968). The dominant fabric in the Silurian–Devonian rocks is related to WSW-directed transport and nappe emplacement (White & Jahns, 1950; Rosenfeld, 1968; Menard & Spear, 1994) and early garnet growth in the domes and vicinity is synchronous with this fabric (Rosenfeld, 1968; Woodland, 1977; Menard & Spear, 1994; Armstrong et al., 1997). Additionally, Rosenfeld (1968) described late top-to-the-east shearing, recorded in rotated garnets and thus synchronous with late garnet growth. In contrast, garnet growth east of the domes near the garnet isograd entirely postdates the dominant fabric (Menard & Spear, 1994; sample TM549). Late chlorite overgrowths on the fabric and replacement of garnet are common across the region.
Peak metamorphic conditions during the Acadian orogeny reached staurolite–kyanite grade in the deepest exposed Silurian–Devonian rocks of the domes and estimated P–T conditions are in the range of 500–650°C and 8–11 kbar (Kohn & Spear, 1990; Menard & Spear, 1994) (Fig. 4a). Similar peak P–T conditions are calculated from pre-Silurian rocks using garnet rim + matrix compositions (Fig. 4b) (Kohn & Spear, 1990; Armstrong et al., 1992; Vance & Holland, 1993; Kohn & Valley, 1994; Armstrong & Tracy, 2000). The similarity of peak P–T conditions, an Acadian Sm/Nd age for garnet rim growth from Gassetts, Vermont (378 Ma; Vance & Holland, 1993) and late Acadian hornblende cooling ages (355–379; Laird et al., 1984; Spear & Harrison, 1989) indicate that all of the peak metamorphic mineral assemblages observed throughout the eastern Vermont belt were produced in the Acadian. In contrast to the relatively high pressures recorded in the vicinity of the domes, the peak P–T conditions at the garnet isograd east of the domes are 450–500°C and 4–5 kbar (Menard & Spear, 1994) (Fig. 4: sample TM549).
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Published P–T paths for Silurian–Devonian rocks of eastern Vermont are dominantly clockwise (e.g. Fig. 4a; see also Menard & Spear, 1994), although near-isothermal loading is also evident over parts of some paths. P–T paths of pre-Silurian rocks (Fig. 4b) display heating with loading (e.g. Kohn & Valley, 1994), and heating with unloading (e.g. Vance & Holland, 1993). However, the significance of the core P–T conditions of the Gassetts schist examined by Vance & Holland (1993) is not clear, because of the possibility that this may represent pre-Acadian recrystallization (e.g. Rosenfeld, 1968; Karabinos, 1984). Of particular interest is the absence of sillimanite in regionally metamorphosed rocks from this part of Vermont, indicating that the P–T paths did not enter the sillimanite field.
The differences in P–T conditions between the domes and the garnet isograd to the east reveal significant differential uplift across strike. For example, the peak pressure of sample TM825a from the Strafford dome (Fig. 4a) is 10 ± 1 kbar whereas the peak pressure of sample TM549 from the garnet isograd (Fig. 4a) is only 5 ± 1 kbar (Menard & Spear, 1994). The difference in peak pressures of 5 kbar implies some 18 km of structural relief between these samples. The present-day separation of these samples is of the order of only 7 km, requiring 10 km of differential uplift, presumably during emplacement of the domes.
The temperatures recorded at the time each rock experienced its peak pressure (525 ± 25°C and 480 ± 20°C, respectively) suggest a relatively steep geotherm, if it may be assumed that these P–T conditions occurred at the same time. In both rocks, the garnet core overgrows or is synchronous with the nappe-stage fabric, suggesting that the timing of the peak pressure conditions recorded by the garnet cores may have been similar. If so, the implied geothermal gradient was 2·5°C/km (45°C/18 km). Such a steep instantaneous geotherm is exactly what is predicted from one-dimensional crustal thickening models during the early stages of post-thickening relaxation (e.g. England & Thompson, 1984). Comparison of the P–T conditions recorded at the peak temperatures (7·5 ± 1·5 kbar, 580°C for sample TM825a and again 5 ± 1 kbar, 480°C for sample TM549) indicate a separation of only 10 km at this time and an instantaneous geothermal gradient of 10°C/km (100°C/10 km). The change from an early geothermal gradient of 2·5°C/km to 10°C/km is best explained as the relaxation of a geotherm by thermal conduction that had been perturbed by crustal thickening (e.g. England & Thompson, 1984).
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MERRIMACK AND BRONSON HILL BELTS P–T EVOLUTION |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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The Merrimack belt includes the Merrimack (Central Maine) synclinorium, a series of Silurian to Devonian metasedimentary rocks that were extensively intruded by early Acadian plutons (the Kinsman quartz monzonite and Bethlehem gneiss) as well as younger plutons. Deformation consisted of west-directed folding, thrust faulting and nappe formation (e.g. Thompson et al., 1968; Thompson, 1985; Chamberlain, 1986; Robinson et al., 1991). Between the Merrimack and eastern Vermont belts lies the Bronson Hill belt. The oldest rocks of the Bronson Hill belt are the metaigneous suites of the Ordovician Oliverian magma series, which crop out in the cores of a series of gneiss domes. Mantling the domes are metavolcanic and metasedimentary rocks with ages ranging from late Ordovician to Devonian. These rocks have also been deformed by isoclinal folding and a series of west-directed thrust faults that place higher-grade metamorphic rocks upon lower-grade rocks forming an inverted metamorphic sequence (Chapman, 1953; Thompson et al., 1968; Spear, 1992, 1993).
Metamorphic parageneses in the Bronson Hill belt of western New Hampshire are intimately associated with structural level (Chapman, 1953; Thompson et al., 1968; Spear et al., 1990, 1995; Spear, 1992, 1993; this study), which has led to a reinterpretation of the structural evolution of the belt. Large-scale, west-vergent recumbent folds have long been recognized in the area (e.g. Thompson et al., 1968; Thompson & Rosenfeld, 1979; Robinson et al., 1991), but recent mapping and the association of distinct metamorphic parageneses with structural level suggests that thrust faults are also common in the region. Figure 5 shows our structural interpretation in a part of southwestern New Hampshire. The Fall Mountain and Chesham Pond thrust faults are shown essentially as mapped by Thompson et al. (1968) and Chamberlain (1986), respectively. The Skitchewaug nappe outliers in the western map region of Fig. 5 are shown to be consistent, with minor modifications, with recent mapping by Armstrong et al. (1997). The root zone of the Skitchewaug thrust fault is shown to be consistent with stratigraphic mapping and the distribution of pseudomorphs (discussed below). The existence of a structural break in the position of the Skitchewaug thrust fault was first inferred by Spear (1992, 1993) and was subsequently verified by geological mapping (e.g. Armstrong et al., 1997). A thrust fault has also been inferred around the borders of the Alstead and Keene domes based on metamorphic disparity between the cover sequence and the underlying rocks (Kohn & Spear, 1999).
Finally, a major décollement has been inferred in the western part of the region just to the east of the Chicken Yard line–Monroe line, here called the Western New Hampshire Boundary Thrust (WNHBT). This thrust fault floors a sequence of garnet- and staurolite-grade rocks that stretches from Massachusetts to northern New Hampshire, which will here be referred to colloquially as the Big Staurolite nappe. The existence of the WNHBT is inferred from the sharp metamorphic discontinuity between rocks of the Big Staurolite nappe and the underlying garnet- and chlorite-zone rocks (see, e.g. Figs 15–17, below), which requires considerable post-metamorphic displacement. It is a testament to the difficulty of purely stratigraphic mapping in this terrane that this fault has not previously been identified, because the metamorphic parageneses absolutely require its existence. Reconnaissance field mapping by Spear and Cheney has identified a shear zone of several meters width with west-vergent kinematic indicators in the position of this fault, but further mapping is required to trace its extent. Additionally, mapping by Armstrong (1995; see also Armstrong et al., 1997) has identified a major shear zone near Bellows Falls, VT, in approximately the required metamorphic position (the Westminster West Shear Zone), and the WNHBT may also be related to this structure.
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In southwestern New Hampshire in the vicinity of Fall Mountain (Figs 3 and 5), there is evidence for three tectonic foliations. Bedding (S0) and a bedding-parallel foliation (S1) are both overprinted and commonly eradicated by a dominant penetrative foliation or crenulation cleavage (S2). This foliation is interpreted to be associated with early west-vergent isoclinal folding (see Thompson et al., 1968; Robinson et al., 1991) and locally has been reactivated during later thrusting, although it is important to note that S2 may not be of the same age or origin in rocks of all structural levels. Towards the WNHBT, a third tectonic fabric (S3) is common. The S3 fabric ranges in intensity from kink bands to penetrative shear fabric that completely disrupts the S2 fabric. Greenschist-facies alteration is common near the WNHBT and the intensity of the alteration correlates with the intensity of the S3 deformation, suggesting that fluids responsible for this alteration gained access along shear zones. In contrast, only two tectonic fabrics have been observed in west–central and northwestern New Hampshire [e.g. in the vicinity of the Orfordville belt and Mascoma dome (Fig. 3) and in the Littleton area (Fig. 1)]. In addition to bedding (S0), a strong foliation (S1) is axial planar to isoclinal folds, and again is thought to be the result of west-directed fold and thrust nappes (Rumble, 1969; Spear & Rumble, 1986; Kohn et al., 1992). A cross-cutting cleavage (S2) is locally present and thought to be the result of post-nappe (D2) upright folding.
Each of the nappes has experienced a different P–T history. Most importantly, the part of the P–T paths along which the dominant fabric (S2 in southwestern New Hampshire and S1 in west–central and northern New Hampshire) is developed differs between nappes, and this relationship can be used to infer the depth and thermal conditions of the crust during deformation. The pertinent observations will be presented below, beginning with the lowest (westernmost) structural level and working structurally upwards (eastward).
Low-grade belt
A zone of chlorite–biotite-grade rocks containing the assemblage quartz + chlorite ± biotite + muscovite + albite + K-feldspar occurs at the boundary between the Vermont and New Hampshire sequences (the Chicken Yard line–Monroe line; Figs 2, 3 and 5). The Chicken Yard line has been variously described as an unconformity, a normal stratigraphic succession, and a fault (Trzcienski et al., 1992; Thompson et al., 1997). Mylonites up to several meters thick occur at the Chicken Yard line with strain features such as mica fish and pyrite cubes with asymmetric tails. Quartz deformation was ductile whereas feldspar deformation was brittle, suggesting that mylonitization occurred at temperatures between 300 and 400°C, consistent with the metamorphic grade. The presence of greenschist-facies recrystallization in the mylonitic rocks near the Chicken Yard line suggests that at least some of the shearing took place under these conditions.
At the latitude of Fall Mountain, isograds near the Chicken Yard line are closely spaced; locally the garnet zone occurs only a few tens of meters east of the Chicken Yard line (Figs 2 and 5). Garnet-bearing assemblages (localities BF-53 and 93-19) contain garnet + biotite + chlorite + muscovite + plagioclase + quartz ± graphite. Biotite is porphyroblastic and pre-dates much of the deformation inasmuch as it is deformed into fish-like structures. Garnet is small (0·5 mm diameter) and generally makes up only a few modal percent of the assemblage. The dominant foliation in rocks of the garnet zone (S2) is a penetrative crenulation cleavage. Figure 6a illustrates small pressure shadows of the S2 fabric developed around a low-grade garnet from sample BF-53, suggesting that garnet growth in this sample pre-dated S2. Figure 6 also shows X-ray maps of the chemical zoning in this garnet. Mg, Fe, Mn and Ca are very nearly unzoned [Xprp = 0·06; Xalm = 0·76; Xsps = 0·14; Xgrs = 0·04; Fe/(Fe + Mg) = 0·93] because of limited progress on the garnet-producing reaction chlorite + quartz = garnet + H2O. Plagioclase is zoned from approximately An31–33 in the core to An2–7 on the rim. The P–T conditions of garnet crystallization for two samples from the garnet zone (BF-53 and 93-19) are estimated to be 475–530°C, 3–4 kbar based on garnet–biotite geothermometry and garnet–plagioclase–muscovite–biotite–quartz geobarometry (Fig. 6f). The lack of significant reaction progress involving garnet in these samples precludes determination of the P–T path.
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The garnet isograd roughly parallels the Chicken Yard line to the north until it enters the Orfordville belt, where it displays a distinct northward bulge (Fig. 2). This is believed to be largely due to the distribution of bulk compositions suitable for the formation of garnet: in the Orfordville belt the isograd is drawn based on the occurrence of garnet in felsic metavolcanic rocks. Therefore, the garnet isograd in the southern part of the Orfordville belt should not be taken as indicative of constancy of peak metamorphic temperature. Along the western part of the Orfordville belt, the garnet and staurolite–kyanite zones are juxtaposed against chlorite–biotite-grade rocks along the Ammonoosuc fault. Interpreted originally as a west-side-up thrust fault (Billings, 1937; Hadley, 1942), it is now believed to be a normal fault (west-side-down) (e.g. Thompson et al., 1968).
Orfordville belt
In the northern half of the map area of Fig. 3, a sequence of rocks collectively referred to as the Orfordville belt crops out. Originally mapped as a separate formation believed to be the oldest rocks in New Hampshire (Hadley, 1942), they are now correlated with other Ordovician, Silurian and Devonian rocks of the Bronson Hill sequence (Thompson et al., 1968).
The P–T evolution of rocks within the Orfordville belt has been described by Spear & Rumble (1986), Kohn et al. (1992) and Florence et al. (1993). P–T paths are typically clockwise with maximum pressures of 6–7 kbar and maximum temperatures of 475–575°C, depending on the metamorphic grade (see Kohn et al., 1992, fig. 7; Florence et al., 1993, fig. 15). The metamorphic gradient within the belt is normal, with higher-grade rocks exposed in deeper structural levels. If the chlorite–biotite-grade rocks to the west of the Ammonoosuc fault are representative of high structural levels of the Orfordville belt rocks, then the range of metamorphic temperatures implied by staurolite–kyanite-grade rocks in the deep structural levels (575°C) and chlorite–biotite-grade rocks in the high structural levels (450°C) must have been of the order of 125°C. At 25°C/km this suggests 5 km of structural throw across the Ammonoosuc fault.
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Garnet growth in schists and felsic metavolcanics of the Orfordville belt was, at least in part, synchronous with development of the nappe fabric whereas staurolite and kyanite typically overgrow this fabric [samples 77-15A, 79-149D, 68-422V in fig. 4 of Spear & Rumble (1986)]. Chemical zoning profiles from the syntectonic parts of these garnets typically show decreasing Mn, antithetically increasing Fe, and little change in Ca or Fe/(Fe + Mg). Plagioclase inclusions typically become more albitic from core to rim. The decrease in XAn at nearly constant XGrs and Fe/(Fe + Mg) results in calculated P–T paths of nearly isothermal loading for the syntectonic portions of the garnets. These P–T paths suggest that metamorphic recrystallization occurred in response to loading from higher-level nappes, and the rotated garnets from which the P–T paths are derived suggest that there was reactivation of the dominant fabric in these rocks during this higher-level nappe emplacement.
Big Staurolite nappe
To the east of the garnet zone at the latitude of Fall Mountain, across the WNHBT, lies what is here called the Big Staurolite nappe. This nappe can be traced from central Massachusetts to northern New Hampshire and is informally named after the characteristic metamorphic paragenesis of late (post-D2), large staurolite porphyroblasts, which are common in many rocks of the nappe. Rocks of the Big Staurolite nappe are equivalent to rocks of the Hardscrabble, Garnet Hill, and Salmon Hole Brook synclines in the Orfordville and Littleton areas of New Hampshire, respectively.
Much of the Big Staurolite nappe is in the staurolite zone and is characterized by distinctively large staurolite crystals (1–5 cm) that typically overgrow the dominant (S2) foliation in the rocks (Fig. 7). In some samples, a well-developed crenulation cleavage is preserved within staurolite crystals (Fig. 7b) whereas in other samples, S2 is penetrative and staurolite overgrows a straight foliation (Fig. 7a). Fabrics within and surrounding garnet crystals suggest that garnet growth in different samples was pre-, syn-, or post-tectonic. For example, garnet crystals in both samples in Fig. 7 are pre-tectonic whereas a sample illustrated by Kohn et al. (1992, sample D84-1C, fig. 8) is syn-tectonic.
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Peak metamorphic conditions in the Big Staurolite nappe near Fall Mountain are in the staurolite zone. Along strike to the north of Fall Mountain (Fig. 3), the metamorphic grade decreases to garnet and, locally, biotite grade (locations MC-9 and D84-1; Fig. 3), and increases again to the staurolite zone (again with large staurolite crystals) near Littleton, New Hampshire (Fig. 1). From this along-strike variation, it is surmised that a near-vertical metamorphic gradient exists within this nappe. Furthermore, near Littleton, the Ammonoosuc fault has juxtaposed low-grade equivalents (chlorite–biotite zone) of the large staurolite rocks of the Salmon Hole Brook syncline against the Walker Mountain syncline (Billings, 1937, 1992; Moench, 1989, 1992). If these rocks are indeed equivalent, the range in temperatures within the Big Staurolite nappe in the Littleton region must have been greater than 100°C (i.e. 575°C at the base and <475°C at the top).
Metamorphic parageneses have been described for rocks of the Big Staurolite nappe by Spear et al. (1990; near Fall Mountain, NH: sample BF-18c), Florence et al. (1993; near Littleton, NH: samples 8835b, 8848, 9047c, 8874, 8809, K8826, LT2a), and Kohn et al. (1992; near Hanover, NH: samples MC-9b, D84-1c, K87-110A, K87-82G, and 77-12Y). Sample BF-52a is typical of these, and shows garnet growth zonation (Fig. 8) with high Mn cores and bell-shaped zoning profiles (XSps = 0·18–0·02). Ca decreases strongly in sample BF-52a (XGrs = 0·24–0·05) suggesting epidote was present in the assemblage during initial garnet growth, but garnets from other samples have lower Ca contents and are relatively unzoned in Ca [e.g. sample BF-18c; figs 11 and 12 of Spear et al. (1990)]. Fe/(Fe + Mg) decreases only slightly from core to rim (from 0·96 to 0·93). The small decrease in Fe/(Fe + Mg) in a garnet growing in the assemblage garnet + chlorite + biotite + muscovite + plagioclase + quartz indicates only a small increase in temperature during growth. Correlation of plagioclase zoning with changes in grossular content of garnet reveals increases in pressure during growth (Fig. 9). Significantly, along the entire strike of the Big Staurolite nappe from the Massachusetts border to Littleton, NH (over 150 km), P–T paths consistently show nearly isothermal loading of 1–4 kbar (3–15 km) during garnet growth (Fig. 9), which is interpreted to have been caused by the emplacement of higher-level nappes (the Skitchewaug, Fall Mountain, and Chesham Pond nappes). Furthermore, many garnets from the Big Staurolite nappe that grew during isothermal loading are syn-tectonic, suggesting local reactivation of foliation during nappe emplacement. Rocks in which garnet apparently postdates the dominant fabric are interpreted as having not experienced significant reactivation of foliation whereas rocks in which garnets apparently pre-date the dominant fabric must have experienced considerable flattening of the foliation following loading and garnet growth.
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Skitchewaug nappe
Rocks of the Skitchewaug nappe (Fig. 5) have a distinctive paragenesis. Common in the schists of this nappe are pseudomorphs comprising predominantly white micas that are after either andalusite or staurolite (Fig. 10; see also Spear et al., 1990, Fig. 9). In most samples, the pseudomorph reaction is complete but in others relict staurolite can still be found (Fig. 10a). In still other samples, a second generation of minerals has grown within the mica pseudomorph including the minerals staurolite, kyanite and fibrolitic sillimanite [e.g. fig. 9 of Spear et al. (1990)]. Another characteristic feature of many samples is the development of chlorite porphyroblasts in the matrix that cut across and include the S2 foliation.
Andalusite pseudomorphs (Fig. 10b) are restricted to the upper levels of the nappe whereas staurolite pseudomorphs (Fig. 10a) are found in the lower part. Near Fall Mountain (e.g. sample BF-12, Fig. 5), andalusite pseudomorphs occur within a few meters of the Bellows Falls pluton, which lies structurally above the nappe. The early andalusite and staurolite porphyroblasts that are unique to the Skitchewaug nappe are interpreted as contact metamorphic minerals and it is inferred that an inverted metamorphic and thermal gradient existed during this early contact metamorphic event.
The white mica that forms the pseudomorphs is not deformed, indicating that the pseudomorph reaction occurred following development of the dominant fabric, S2. Furthermore, S2 appears to be wrapped around the pseudomorphs (e.g. Fig. 10a), suggesting that the porphyroblasts were present at the time of S2 development. Therefore, the parageneses of these samples suggest the sequence (1) early porphyroblast formation, (2) development of S2 (the nappe stage fabric related to isoclinal folding), and (3) development of the pseudomorphs and matrix chlorite.
A P–T path for rocks from this structural level was calculated by Spear et al. (1990) from zoning in garnet and plagioclase from a sample containing pseudomorphs after andalusite [sample BF-86B; figs 9b, 14, 15 and 16 of Spear et al. (1990)—it should be noted that the sample location is misplaced in fig. 1 of that paper and should be closer to sample BF-89]. Garnet from this sample is zoned with slightly decreasing Mn and increasing Ca. Fe, Mg and Fe/(Fe + Mg) are relatively unzoned except at the rim where Fe/(Fe + Mg) increases from 0·85 to 0·89. Plagioclase zoning and inclusions within garnet indicate An32 in equilibrium with garnet core and An22 in equilibrium with garnet rim. The preferred model for the evolution of this sample is garnet core growth with the early assemblage garnet + biotite + andalusite + muscovite + plagioclase + quartz and later growth with the assemblage garnet + biotite + chlorite + muscovite + plagioclase + quartz. The P–T path [Fig. 11; see also fig. 16 of Spear et al. (1990)] shows a period of nearly isobaric heating through the andalusite field (contact metamorphism from the Bellows Falls pluton) followed by 2–3 kbar (7–10 km) of loading. The sharp change from isobaric heating to isothermal loading occurs at the assumed assemblage change. As with the lower-level nappes, the increase in pressure recorded in the P–T paths of the Skitchewaug nappe is interpreted to have occurred in response to the emplacement of higher-level nappes (the Fall Mountain and Chesham Pond nappes). The pseudomorphing reaction suggested by Spear (1992, 1993) is the retrograde progress of the typical prograde reactions garnet + chlorite + muscovite = staurolite + biotite + H2O or garnet + chlorite + muscovite = andalusite + biotite + H2O. Because the garnet from which the P–T path was calculated was produced by these reactions, the pseudomorph reactions must have also proceeded during loading. These reactions both have positive P–T slopes (Fig. 11b). Accordingly, an increase in pressure will stabilize the low-temperature assemblage, provided sufficient H2O is added to the rock. The source of the fluids necessary to drive these reactions to the left is not known, but the thrust faults associated with the loading are likely conduits for fluid migration. It should be noted also that the absence of deformation in the pseudomorphs themselves suggests little reactivation of S2 during loading.
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Fall Mountain nappe
Rocks from the Fall Mountain klippe have been the subject of intensive study (Spear et al., 1990; Spear & Kohn, 1996; Kohn et al., 1997) and rocks from the root zone have been considered by Spear et al. (1995) (Fig. 12). Rocks from both places have P–T paths that involve an episode of nearly isobaric heating in the andalusite field followed by loading of 2 kbar. Rocks from the Fall Mountain klippe underwent dehydration melting (Spear & Kohn, 1996; Kohn et al., 1997) as did some rocks from the root zone.
The early, low-pressure metamorphism experienced by rocks of the Fall Mountain nappe is evidenced by early andalusite porphyroblasts that have been pseudomorphed by sillimanite [e.g. fig. 2a of Spear et al. (1990)]. Andalusite pseudomorphs are most abundant in the vicinity of the Bethlehem gneiss (Bellows Falls pluton), and are interpreted as a contact metamorphic assemblage. The andalusite pseudomorphs are generally oriented in random planar arrays but sometimes form a mineral lineation and are locally folded by F2 (nappe stage) folds.
Five generations of garnet growth have been documented in rocks of the Fall Mountain klippe (Kohn et al., 1997) and two to three generations in rocks of the root zone (Spear et al., 1995). The early generation of garnet (Grt1) pre-dates the S2 fabric whereas the latest generations (Grt4 and Grt5) clearly postdate it [e.g. fig. 2c of Spear et al. (1990) and fig. 2 of Spear et al. (1995)]. The absence of obvious inclusion relationships makes it difficult to ascertain the relationship between fabric development and intermediate-generation garnets (Grt2 and Grt3). Some of the leucosomes from the migmatites are mildly deformed, displaying sigmoid shapes, but most are undeformed, indicating that development of S2 was over before the peak temperature was achieved. Late muscovites produced during crystallization of leucosomes crosscut the S2 fabric and are completely undeformed.
The loading experienced by rocks of the Fall Mountain nappe is interpreted to have occurred in response to emplacement of the higher-level Chesham Pond nappe, and the change in pressure recorded by rocks of the Fall Mountain nappe (
P = 2·5 kbar) suggests that the Chesham Pond nappe was 8 km thick. Additionally, inasmuch as the loading was followed by isobaric heating by as much as 100°C, it is inferred that the Chesham Pond nappe was hot (locally over 750°C) when it was emplaced.
Chesham Pond nappe
The highest structural level in central New England is the Chesham Pond nappe. Parageneses of rocks from this structural level have been described by Chamberlain (1986), Spear (1992) and Spear et al. (1999). Metamorphic grade reaches the cordierite + garnet zone and migmatites are typical. K-feldspar is common in these rocks, suggesting that partial melting occurred following muscovite breakdown to Al2SiO5 + K-feldspar, which requires a prograde P–T path below 4 kbar (Spear et al., 1999). Peak P–T conditions are 725°C, 3–4 kbar and the path is slightly counter-clockwise (Fig. 13). Significantly, the entire P–T evolution occurred at low pressure, and there is no evidence of a loading event as is seen in the Fall Mountain nappe P–T paths.
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Leucosomes in migmatites from the Chesham Pond nappe are locally concordant with the S2 fabric, and some leucosomes are sigmoidal. However, the bulk of the leucosomes are undeformed, and micas produced on crystallization of the leucosomes are completely undeformed. These observations suggest that deformation was over by the peak of metamorphism.
Peak temperatures experienced by rocks of the Chesham Pond and Fall Mountain nappes are similar, although the pressure at the peak temperature was lower in the structurally higher Chesham Pond nappe. Attainment of a peak temperature of 720–750°C at a pressure of 3 or 5 kbar requires input of heat, and the likely candidates are the Bethlehem gneiss and Kinsman quartz monzonite. A single pluton at a temperature of 900°C intruding a country rock at 500°C can achieve a maximum temperature in the contact aureole of 700°C, but the temperature of the contact aureole drops off to 600°C in a few hundred meters from the contact. Although this type of aureole is consistent with what is observed in the Skitchewaug nappe beneath the Bellows Falls pluton, it is too steep a gradient to be consistent with the regional low-pressure, high-temperature metamorphism seen in the Fall Mountain and Chesham Pond nappes. A possible explanation is that the high-grade metamorphism in these nappes was produced by overlapping thermal aureoles from both the Bethlehem gneiss and Kinsman quartz monzonite. Overlapping thermal aureoles require that both plutons intrude within a few tens of thousands of years of each other, because the time constant for thermal decay of these kilometer-thick, sheet-like plutons is only of the order of 30 kyr. More-or-less simultaneous intrusion is within the precision of the ages of crystallization of these plutons [i.e. the Bethlehem gneiss has been dated at 395–410 Ma (Aleinikoff & Moench, 1987; Moench, 1989; Moench & Aleinikoff, 1991; Kohn et al., 1992) and the Kinsman quartz monzonite has been dated at 402–413 Ma (Lyons & Livingston, 1977; Barreiro & Aleinikoff, 1985)]. A scenario that fits the available petrological data calls for the Bethlehem gneiss to intrude first, producing contact aureoles above and below. The Chesham Pond nappe is then emplaced with simultaneous intrusion of the Kinsman quartz monzonite, causing the increase in pressure seen in the Fall Mountain nappe followed by raised temperature that produced anatexis. Although field maps are insufficiently precise to permit verification because of poor outcrop exposure, it is likely that the Kinsman quartz monzonite intruded along the Chesham Pond thrust, and presumably helped lubricate the fault surface. The Fall Mountain thrust is coincident with the Bethlehem gneiss in places and it is similarly likely that thrust movement was aided by a partially molten intrusion. However, at least some of the movement on the Fall Mountain thrust must have occurred following cooling of the Fall Mountain nappe and Bethlehem gneiss to 550°C because rocks of the underlying Skitchewaug nappe reveal no temperature rise following loading by the Fall Mountain nappe, which would have occurred had the nappe been near its thermal peak of 725°C.
Alstead and Keene domes
The Bronson Hill anticlinorium contains a series of domes that are cored by gneisses of the Ordovician Oliverian magma series (the Alstead, Keene and Mascoma domes of Figs 3 and 5). The cover sequence to the domes (i.e. the Big Staurolite nappe; Fig. 5) has generally been interpreted as being para-autochthonous to the dome rocks. However, Kohn & Spear (1999) presented evidence that the Oliverian magmas intruded at pressures of 8–10 kbar and that the metamorphic grade (greenschist facies) of the dome rocks themselves never reached the same grade (amphibolite facies) as the cover sequence during Acadian metamorphism (Fig. 14). These observations require both a structural break between the dome rocks and the overlying metavolcanic and metasedimentary cover sequence and that the two rock suites were not juxtaposed until the cover rocks had cooled somewhat from their peak temperature.
Retrograde metamorphism
High-temperature ‘retrograde’ metamorphism is developed in some lithologies, and is especially prevalent in the Chesham Pond and Fall Mountain nappes. Rocks of both nappes locally experienced partial melting and the water released from the crystallization of leucosomes promoted hydration reactions during cooling. For example, in both the Fall Mountain klippe and the Chesham Pond nappe, retrograde progress of the reaction sillimanite + biotite = garnet + K-feldspar + melt is pervasive and in the Fall Mountain klippe, retrograde progress of muscovite + quartz = sillimanite + K-feldspar + melt has produced abundant late muscovite. In rocks that contain the subsolidus assemblage sillimanite + biotite + garnet + muscovite + quartz, cooling has produced late garnet and muscovite by the water-absent reaction sillimanite + biotite = garnet + muscovite (i.e. garnet generation G4 of Fig. 12; Spear et al., 1990, 1995, 1999; Kohn et al., 1997).
Greenschist-facies retrograde metamorphism is developed locally throughout western New Hampshire. In some localities it is pervasive (i.e. most rocks in some localities contain late chlorite alteration). For example, Kohn et al. (1997) described from the Fall Mountain klippe a fifth generation of garnet growth in rocks, generally in association with chlorite by the reaction sillimanite + biotite + H2O = garnet + chlorite + muscovite (i.e. Grt5, Fig. 12). Widespread chloritization is also evident in rocks at the base of the Big Staurolite nappe. The intensity of chloritization increases as the Western New Hampshire Boundary Thrust is approached to the extent that rocks at the base of the unit are nearly completely converted to greenschist-facies assemblages (although pseudomorphs after staurolite remain). Greenschist-facies assemblages (chlorite + albite + epidote after plagioclase + hornblende) are also developed in shear zones at the margin of the Alstead dome (Kohn & Spear, 1999).
In summary, high-grade retrogression was produced pervasively from melt crystallization and then locally where the appropriate water-absent assemblage could proceed to react. From this observation it does not appear that retrograde fluids permeated the nappe sequence during initial cooling [see also Kohn et al. (1997) for isotopic evidence of closed-system behavior at high temperature]. Greenschist-facies retrogression, however, is widespread, has affected all nappes, and is most intense in and near faults and shear zones. Furthermore, Kohn et al. (1997) presented stable isotope evidence that this late retrograde hydration occurred at 
475°C. From these observations, it is concluded that the final assembly of the nappe pile probably occurred at the time of greenschist-facies alteration and that fluid causing the retrogression probably gained access along the faults and shear zones.
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SUMMARY AND INTERPRETATION OF P–T EVOLUTION ACROSS CENTRAL NEW ENGLAND |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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Figures 15–17 summarize the maximum P–T conditions and P–T paths experienced by rocks across strike in central New England. Maximum pressures range from nearly 10 kbar in the eastern Vermont domes to 6 kbar in central New Hampshire (Figs 15 and 17). Maximum temperatures in eastern Vermont are only of the order of 600°C whereas maximum temperatures in central New Hampshire approach 750°C. The pattern of isograds shows strong correlation with structural position (Fig. 15d). Isograds in eastern Vermont are symmetrical about the domes, with grade increasing with structural depth. In western New Hampshire, isograds depicting early metamorphic history are coincident with inferred thrust faults. Only the sillimanite isograd crosses structural boundaries (see Figs 2 and 5).
P–T paths also show marked differences across strike (Fig. 16). Vermont P–T paths are clockwise whereas those in western New Hampshire are dominantly counter-clockwise. It is particularly significant that the lower nappes in the New Hampshire sequences all involve an episode of increasing pressure in their P–T paths, whereas the Chesham Pond nappe does not. The pressure increases are interpreted to have occurred during the emplacement of the overlying nappes. The absence of such a pressure increase in the Chesham Pond nappe suggests that it is the highest structural level.
Figure 17 summarizes the peak P–T conditions in a schematic diagram in which the different structural levels are stacked vertically. It should be noted that this representation is not meant to imply that the New Hampshire sequence rocks were ever as far west as the Chester dome in Vermont. Clearly apparent is the fall in temperature and pressure going upward from the Vermont domes towards the Chicken Yard line (a normal metamorphic sequence), and the discontinuities in pressure and temperature across thrust surfaces through the inverted metamorphic sequence in the New Hampshire rocks. It is important to emphasize that Fig. 17 does not represent a crustal thermal structure at any point in time, but is, rather, a representation of the maximum P and T conditions achieved in each structural level. Indeed, the marked P–T discontinuity at the Chicken Yard line virtually requires that the New Hampshire nappes were cool before being emplaced into their present position.
Tectonic summary
The combined petrological and textural observations presented above provide constraints on the tectonic assembly of the central New England metamorphic belt during the Acadian orogeny. The earliest metamorphism observed is the contact metamorphic aureoles associated with the Bethlehem gneiss (Bellows Falls pluton), which produced andalusite-bearing assemblages both above and below the pluton in the Fall Mountain and Skitchewaug nappes, respectively (Fig. 18a). Intrusion of the Kinsman quartz monzonite at a structurally higher level must have occurred only shortly thereafter, along with thrusting of the Chesham Pond nappe over the Fall Mountain nappe (Fig. 18b). Continued westward thrusting emplaced the Fall Mountain and Chesham Pond nappes on the Skitchewaug nappe (Fig. 18c).
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It is significant that the early contact metamorphic assemblage of the Fall Mountain nappe is similar to that of the underlying Skitchewaug nappe (andalusite + garnet + biotite), but then the Fall Mountain nappe experienced a considerable rise in temperature following loading whereas the Skitchewaug nappe experienced only minor heating or isobaric cooling. The simplest explanation is that the nappes were not vertically juxtaposed at the time of the peak of metamorphism in the Fall Mountain nappe. For this reason, the emplacement of the Fall Mountain nappe on the Skitchewaug nappe is inferred to have occurred following cooling of the high-grade assemblages in the upper nappes. The amount of time required for the cooling need not be great, inasmuch as the plutons are not thick and crystallization probably occurred within a few hundred thousand years of the time of intrusion.
Ages of the Bethlehem gneiss and Kinsman quartz monzonite fall in the range 393–413 Ma (Lyons & Livingston, 1977; Barreiro & Aleinikoff, 1985; Aleinikoff & Moench, 1987; Kohn et al., 1992) with considerable overlap between published ages. These plutons intrude the Littleton Formation (e.g. Hadley, 1942; Rumble, 1969), which contains Lower Emsian fossils (Boucot & Arndt, 1960; Boucot & Rumble, 1980). The Lower Emsian has recently been dated at 408 ± 2 Ma (Tucker et al., 1998), providing a tight upper age limit for plutonism. Metamorphic monazite and zircon ages of 408–392 Ma have been reported from central New Hampshire and Maine (Barreiro et al., 1988; Eusden & Barreiro, 1988; Smith & Barreiro, 1990; Zeitler et al., 1990). These ages are from rocks in structural levels above the Fall Mountain nappe (e.g. the Chesham Pond nappe) and are generally consistent with the model presented here that the metamorphic peak in these high-level nappes was caused by heat from the syn-tectonic plutons.
Emplacement of the Skitchewaug and higher nappes on the Big Staurolite nappe is the next event recorded in the rocks of western New Hampshire (Fig. 18d). It is not clear whether this happened immediately following the emplacement of the Fall Mountain nappe, or after a hiatus in deformation. Before the isothermal loading experienced by rocks of the Big Staurolite nappe, the P–T conditions were in the lower to middle garnet zone (Fig. 9), and there is some evidence that the geothermal gradient may have been slightly elevated, at least locally, at this time. For example, tails of some P–T paths in Fig. 9 are in the andalusite field and Florence et al. (1993) have described early andalusite followed by garnet + staurolite + kyanite + biotite assemblages from the Salmon Hole Brook Syncline in the Littleton area.
The structural evolution shown in Fig. 18 is that of an in-sequence thrust stack. Total shortening is difficult to estimate, but the current exposure of the nappes can be used to provide a minimum estimate. The distance across strike from the Chicken Yard line to the root zone of the Fall Mountain nappe near Gilsum, NH, is 20 km (Fig. 5). Assuming that before erosion the nappes were stacked vertically in the vicinity of the Chicken Yard line, a minimum shortening of 75% is implied (l/l = 60 km/80 km).
Onset of metamorphism in the Silurian–Devonian rocks of eastern Vermont is not well constrained. An Sm–Nd age of 378 Ma on the rim of a garnet from the pre-Silurian Gassetts schist on the west flank of the Chester dome (Vance & Holland, 1993) suggests that the rocks in the core of the domes were undergoing prograde metamorphism at this time, and it is reasonable to assume that the cover rocks were as well. However, ages of monazite from Silurian–Devonian rocks of 354 Ma (Wing et al., 1999; Ferry, 2000) suggest that metamorphism in the cover sequence may have been much younger than was previously thought.
The 8–11 kbar peak pressures in the vicinity of the domes in eastern Vermont require significant crustal thickening, and the question naturally arises: ‘What was emplaced on eastern Vermont to cause the crustal thickening?’ One possibility is that the New Hampshire nappe sequence was thrust directly on eastern Vermont to cause the Barrovian metamorphism, a model favored by Spear (1992, 1993). However, this scenario is no longer favored for several reasons. First, the peak pressures in garnet-zone rocks now exposed along the Chicken Yard and Monroe lines are only 3–4 kbar whereas the peak pressures of rocks in the overlying Big Staurolite nappe are 5–6 kbar. If the garnet-zone metamorphism was caused by emplacement of the New Hampshire nappe sequence, this could be accommodated only if 3–10 km of erosion of the New Hampshire nappes had occurred before their emplacement onto the garnet-zone rocks. Second, the 28–35 km of burial implied by the 8–11 kbar pressures experienced by rocks of the Vermont domes is significantly greater than the thickness of the New Hampshire nappe sequence, especially if the thickness is limited to 10–15 km based on the 3–4 kbar pressures of the garnet-zone rocks in the low-grade belt. Third, the apparent absence of high-grade metamorphism in the Oliverian gneisses of the Alstead and Keene domes (i.e. Kohn & Spear, 1999) requires that the rocks of the Big Staurolite and higher nappes were placed on top of these dome rocks after they cooled below 500°C. Fourth, the pervasive greenschist-facies alteration at the base of the Big Staurolite nappe and elsewhere throughout the New Hampshire nappe sequence suggests significant fluid infiltration at temperatures of 400–450°C, presumably along structural breaks. The simplest explanation is that this infiltration occurred when these faults were active and the nappe sequence had cooled to greenschist-facies conditions.
Our revised scenario invokes burial of rocks of the Vermont sequence beneath rocks that now form the low-grade belt, producing a more-or-less normal metamorphic gradient (Fig. 19). It is proposed that this might have been the Bronson Hill arc, with the New Hampshire nappes eastward of the arc (Fig. 19). Major recumbent fold nappes are mapped in eastern Vermont, and presumably formed during this partial subduction, contributing to crustal thickening. Additional burial must have occurred along thrust faults or distributed shear zones within the eastern Vermont belt. Substantial reflectors are observed beneath western New Hampshire in the COCORP seismic line (Ando et al., 1984), and it is possible that one or more of these reflectors represents the proposed thrust faults, although direct evidence is lacking.
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Following burial to 30 km depth, eastern Vermont experienced 7 km of differential uplift of the dome rocks relative to the lower-pressure rocks toward the Chicken Yard–Monroe lines. Finally, the New Hampshire nappe sequence was thrust westward along the Western New Hampshire Boundary Thrust fault accompanied by greenschist-facies alteration. It is also proposed that ramp anticlines associated with related thrust faults produced the Alstead and Keene domes.
These hypotheses are testable with thermochronology. For example, the timing of this late shearing is not known, but the greenschist-facies metamorphism that is produced in the shear zones, at the base of the Big Staurolite nappe, and, to a lesser degree, throughout the New Hampshire nappe sequence, suggests a temperature of 400–450°C, that is, some time after hornblende 40Ar/39Ar closure and before muscovite and biotite closure. Argon thermochronology suggests that these temperatures were attained some time between 300 and 350 Ma (Laird & Albee, 1981; Harrison et al., 1989; Spear & Harrison, 1989), which is substantially later than the 378 Ma Sm–Nd age on garnet from Gassetts (Vance & Holland, 1993), and later than the 354 Ma age of monazites from Vermont (Wing et al., 1999; Ferry 2000). This evidence suggests that the assembly of the central New England metamorphic belt was a protracted process that extended through the Devonian into the Carboniferous. Additional ages on the timing of the metamorphism in Vermont, the different structural levels of New Hampshire, and the greenschist-facies alteration should help constrain further the timing of this assembly.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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A combination of P–T studies and fabric analysis provides a framework within which to interpret the sequence of tectonic assembly in central New England. The model presented here is broadly consistent with prior geological mapping based on stratigraphic correlations, although there are numerous examples where the thrusts proposed here (e.g. Fig. 5) require reinterpretation of earlier mapping (e.g. Thompson et al., 1968; Thompson & Rosenfeld, 1979; Armstrong et al., 1997). Indeed, some of the results of this study are not at all obvious based solely on stratigraphic mapping. Clearly, considerable additional work needs to be done, especially in constraining the timing of different events. Testing of this model can be best achieved from detailed analysis of the timing of peak metamorphism across strike from petrologically well-characterized structural levels. Ultimately, any satisfactory model will have to reconcile all of these data.
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ACKNOWLEDGEMENTS |
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The results and interpretations presented in this paper have benefited from many years of discussions with numerous individuals. The authors are particularly indebted to the insights of J. B. Thompson, Jr, D. Rumble, III, and P. Robinson, whose work has influenced the thinking of all of the authors. Insightful and informative reviews were provided by T. Dempster and S. Harley. This work was supported by NSF grants EAR-8916417, EAR-9220094, and EAR-0073747 (to F.S.S.).
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FOOTNOTES |
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*Corresponding author. E-mail: spearf{at}rpi.edu
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REFERENCES |
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TOPABSTRACTINTRODUCTIONEASTERN VERMONT P-T EVOLUTIONMERRIMACK AND BRONSON HILL...SUMMARY AND INTERPRETATION OF...CONCLUSIONSREFERENCES |
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