Geochemistry of the metamorphosed Ordovician Taconian Magmatic Arc, Bronson Hill anticlinorium, western New England

Physics and Chemistry of the Earth 27 (2002) 5–45 www.elsevier.com/locate/pce

Geochemistry of the metamorphosed Ordovician Taconian Magmatic Arc, Bronson Hill anticlinorium, western New England Kurt Hollocher

a,*

, Jon Bull b, Peter Robinson

c,d

a

c

Geology Department, Union College, Schenectady, NY 12308, USA b 7021 Merrilee Lane, Dallas, TX 75214, USA Department of Geosciences, University of Massachusetts, Amherst, MA 01003, USA d Geological Survey of Norway, Trondheim N-7491, Norway

Abstract The Bronson Hill anticlinorium (BHA) in western New England is a north- to northeast-trending belt of gneiss domes containing a metamorphosed stratified sequence of Ordovician, Silurian, and Devonian sedimentary and volcanic rocks. These overlie mostly late Ordovician intrusive igneous rocks. The intrusive rocks, exposed in the dome centers, are dominated by felsic gneisses of tonalitic to granitic compositions. They are the exposed parts of a composite batholith that developed along the axis of the Taconian volcanic arc. The Fourmile Gneiss of the Pelham dome, and the Monson Gneiss of the Monson dome, covered in most detail here, are calcalkaline, dominantly tonalitic to granodioritic gneisses having 61–78% SiO2 . These are divided into two geochemical groups. The most abundant rocks have low Sr (
50–200 ppm), high Y (
10–50 ppm), low Al2 O3 , high FeOt , commonly negative Eu anomalies, and have flat to somewhat concave upward MREE and HREE patterns. These rocks were probably derived from melting of plagioclase–pyroxene–amphibole granulites in the deep crust. Rocks of the subordinate high-Sr group are similar in composition to adakites and have high Sr (
300–600 ppm), low Y (
1–13 ppm), higher Al2 O3 , lower FeOt , commonly no Eu anomalies, and are strongly depleted in HREE. These rocks were probably derived from a pyroxene–amphibole–garnet  plagioclase source at higher pressure. Amphibolites are low in abundance in the Fourmile and Monson Gneisses. Most are a low-Nb type (<6 ppm Nb), compositionally similar to typical calc-alkaline island arc basalts. These were probably derived by melting of spinel lherzolite in the mantle wedge under the arc. Gabbroic anorthosite and ultramafic rocks also occur in the Monson dome and appear to be cumulates from magmas similar to the low-Nb amphibolites. A high-Nb (6–12 ppm) amphibolite group was identified and is unique to the Monson Gneiss. These rocks are transitional between calc-alkaline and alkaline basalts, are strongly LREE-enriched (La is 46–440 times chondrite), are enriched in strongly incompatible elements, and two of seven samples have normative nepheline. A broader survey of available analytical data from BHA felsic gneisses shows that gneiss compositions vary from dominantly tonalitic and granodioritic in the southern BHA to dominantly granitic from the Croydon and Mascoma domes (west-central New Hampshire) northward. This variation suggests that the southern part of the arc exposed in the BHA was based on mafic crust, whereas the northern part was based on intermediate to felsic crust. Examination of radiometric dates of Taconian igneous rocks, and biostratigraphic constraints in the Taconic allochthons and in the Taconian foreland, are consistent with a model in which the BHA gneisses are the youngest igneous components of the Taconian arc proper. Collision of this composite arc with Laurentia reached a conclusion in the latest Caradoc or Ashgill of the late Ordovician or in the very earliest Silurian. Younger plutons continuing into the Silurian (e.g., Highlandcroft in part) may be related to magmas generated during post-collision delamination or detachment of the subducted Iapetus oceanic slab. Ó 2002 Elsevier Science Ltd. All rights reserved.

1. Introduction The Bronson Hill anticlinorium (BHA) is a north- to northeast-trending lithotectonic belt, extending from southern Connecticut into Maine, containing about 20 *

Corresponding author. E-mail address: [email protected] (K. Hollocher).

structural domes (Fig. 1) of Acadian (early Devonian), late Devonian to early Mississippian, or early Pennsylvanian age (Robinson et al., 1998). This belt contains a wide variety of metamorphosed plutonic, volcanic, and sedimentary rocks that range in age from late Precambrian to Devonian. Exposed in the cores of most domes are Ordovician massive to layered gneisses (referred to here as ‘‘dome gneisses’’), spanning a composition range

1474-7065/02/$ - see front matter Ó 2002 Elsevier Science Ltd. All rights reserved. PII: S 1 4 7 4 - 7 0 6 5 ( 0 1 ) 0 0 0 0 2 - X

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 1. Generalized geologic map of western and southern New England, emphasizing structures related to the Taconic Orogeny (adapted from Field, 1975; Lyons et al., 1986; Robinson, 1986; Aleinikoff and Green, 1987; Rogers et al., 1990; Sevigny and Hanson, 1993; Lyons et al., 1997). Numbers are analyzed samples in Table 3. SFB is the Shelburne Falls belt described in the text. WD is the Waterbury dome in Connecticut which contains in part gneisses of probable Ordovician age (Dietsch, 1989; Rodgers, 1985). Rectangle outline is the area of Fig. 2.

from ultramafic to granitic, that are the Oliverian Magma Series of Billings (1937). These rocks have historically been interpreted to be the metamorphosed plutonic roots of the Taconian volcanic arc that collided with Laurentia in middle to late Ordovician time (e.g., Robinson and Hall, 1980; Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985). The interpretation that the BHA represents an exposed portion of the Taconian arc is based on: (1) the elongate map pattern of the BHA; (2) the island arc

compositional affinities of the dome gneisses and related volcanics; (3) inferred or demonstrated middle to late Ordovician (Taconian) ages of the dome gneisses and related volcanics (e.g., Billings, 1937; Tucker and Robinson, 1990; Moench et al., 1995); and (4) the position of the BHA east of, and parallel to, the axes of Taconian foreland depositional basins and the belt of Taconian metamorphism and deformation in western New England, eastern New York, and southeastern Quebec. The dome gneisses, Ordovician volcanics, and related rocks

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

in the BHA are overlain by a relatively thin sequence of cover rocks of Silurian through Devonian ages. These metamorphosed Silurian and Devonian units and correlatives thicken rapidly east of the BHA in the Merrimack–central Maine synclinorium (e.g., Hatch et al., 1983; Thompson, 1985; Moench and Boudette, 1987). The location of this basin immediately east of the Taconian arc system, presence of volcanics having possible back-arc basin affinities (Hollocher, 1993), and recognition of another (though earlier) back-arc system in Newfoundland (O’Brian et al., 1997), suggests that this basin may have formed in a back-arc extensional environment. The basement for the Merrimack–central Maine synclinorium is unclear. Within the synclinorium in the granulite-facies region of central- and south-central Massachusetts and northern Connecticut there occur possible pre-Silurian gneisses exposed along Acadian thrust faults (Berry, 1992). These gneisses may represent basement beneath this portion of central New England. Farther to the east is the Massabesic Gneiss complex (Fig. 1; e.g., Aleinikoff and Green, 1987; Aleinikoff et al., 1995; Dorais and Wintsch, 1998), which has been suggested to be a part of the Avalon block margin. East of the Massabesic Gneiss complex is the Nashoba block (Fig. 1), a complex terrane that includes metamorphosed plutonic, volcanic, and sedimentary rocks, in part arcrelated. The stratified rocks extend from late Proterozoic to probably early Silurian, the plutons into the Mississippian (e.g., Acaster and Bickford, 1999). Both Massabesic and Nashoba rocks were assigned by Robinson et al. (1998) to a pre-Silurian ‘‘Medial New England’’ that also includes the Bronson Hill rocks discussed in this paper. This interpretation would have ‘‘Medial New England’’ rocks as the basement for the entire Merrimack–central Maine synclinorium. Dorais and Paige (2000) interpret the petrology, geochemistry, and isotope systematics of Devonian and younger plutons in central and northern New England as indicating only Grenvillian basement in the western part of the Merrimack–central Maine synclinorium, and Avalonian basement to the east. This interpretation omits consideration of late Precambrian (e.g., ‘‘Avalonian’’) rocks in the core of the Pelham dome (Tucker and Robinson, 1990). Because the Taconian volcanic arc was originally isolated from Avalon and Laurentia in the Ordovician, such basement must have been emplaced against Laurentia during the Ordovician and later collisions. The BHA dome gneisses have been studied in some detail, and numerous chemical and petrologic analyses are available (Billings and Wilson, 1965; Foland and Loiselle, 1981; Hodgkins, 1985; Leo, 1985, 1991; Leo et al., 1984; Pogorzelski, 1983; Webster and Wintsch, 1987; Bull, 1997; Fitz, 1996). This paper presents new data and interpretations for the origin of the Fourmile Gneiss in the Pelham dome and the Monson Gneiss in

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the Monson dome (Fig. 2), and summarizes the petrology, geochemistry, and age relationships of BHA dome gneisses in terms of magma origins and the role of the BHA magmatic arc in the Taconian orogeny.

2. Geologic setting The oldest unit exposed in the BHA is the Dry Hill Gneiss (613  3 Ma, Tucker and Robinson, 1990), which has been interpreted as a sequence of metamorphosed high-K rhyolites (Ashenden, 1973; Hodgkins, 1985; Robinson and Tucker, 1996). This and associated units are exposed only in the core of the Pelham dome (Fig. 2). Some Dry Hill Gneiss zircons have late Archean and early and middle Proterozoic inheritance (Tucker and Robinson, 1991; Robinson et al., 1992), indicating that production of the Dry Hill Gneiss magmas involved older continental material. The Monson and Fourmile gneisses range from ultramafic to granitic, but with a mafic–felsic bimodal distribution in which felsic rocks are overwhelmingly dominant. The felsic rocks are largely tonalitic and granodioritic gneisses, and range from moderately well foliated and layered to massive, and from metaluminous to slightly peraluminous rocks having up to
2% normative corundum. The Fourmile Gneiss overlies the older rocks in the Pelham dome (Dry Hill Gneiss) probably along a fault (Robinson and Tucker, 1996; Robinson et al., 1998), whereas the Monson Gneiss, to the east, occupies the entire exposed core of the Monson dome. The Monson and Fourmile gneisses are similar in appearance and structural position to granitoid dome gneisses that occupy the cores of many other domes in the BHA. These rocks are interpreted to have largely plutonic protoliths based on the massive texture of some varieties, typical meter scale homogeneity of massive and foliated varieties, angular inclusions resembling xenoliths, crosscutting relationships, and igneous chemical compositions. Layered gneisses are common in some areas and these have been interpreted by some as deformed layered volcanics (e.g., the Monson Gneiss, Leo et al., 1984). However, many of the layered rocks are in zones of high strain, and it seems likely that the majority formed from plutonic rocks during deformation (Robinson et al., 1989). The BHA dome gneisses therefore appear to be the exposed portions of a deformed composite batholith. Leo et al. (1984) report composite sample zircon ages of 456  10 Ma for the Glastonbury dome gneiss and 435  6 Ma for the Monson dome, though they cite possible problems with the accuracy of these values. The BHA dome gneisses were assigned a group age of 444  8 Ma by Zartman and Leo (1985); based on a composite of zircon data from gneisses in several domes. Tucker and Robinson (1990) reported higher-precision

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 2. Generalized geologic map of the principal study area. Sample numbers and sampling locations are indicated, and data are given in Tables 1–3 (adapted from Robinson et al., 1993).

zircon ages of several gneisses from the Pelham, Monson, Warwick, and Keene domes in the central BHA, including the Monson and Fourmile gneisses, that span an age range from 454þ3=2 to 442þ3=2 Ma. Moench et al. (1995) report ages of dome gneisses in several

domes in northern New Hampshire that span the range 456  3 to 441  5 Ma. Within the estimated uncertainties these ages are all in the Ordovician Caradoc and Ashgill series, 458–443 Ma, based on the time scale of Tucker and McKerrow (1995).

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Throughout most of the BHA the dome gneisses are overlain by the Ammonoosuc Volcanics (in the sense of Schumacher, 1988) and related units. The Ammonoosuc ranges in thickness up to 1200 m (Schumacher, 1988), and contains a laterally coherent internal stratigraphy that includes a lower member dominated by amphibolites, a middle garnet quartzite member, and an upper member dominated by felsic gneisses. Minor amounts of interbedded calc-silicate rock, marble, and other rocks of sedimentary and volcanoclastic origin also occur. The Ammonoosuc is commonly overlain by the Partridge Formation, which is dominated by rusty weathering mica schist but contains felsic and mafic volcanics near its base that are compositionally similar to the Ammonoosuc Volcanics (Hollocher, 1993). The volcanics in both of these units are bimodal, dominated by mafic rocks. The Ammonoosuc volcanics have yielded a concordant zircon age of 453  2 Ma from the upper member, and a felsic volcanic in the Partridge Formation has yielded an age of 449þ3=2 Ma (Tucker and Robinson, 1990). These rocks have long been interpreted as being volcanics of the Taconian volcanic arc, as supported by: (1) their stratigraphic position above the Ordovician dome gneisses; (2) interbedding with rocks having sedimentary protoliths; (3) inferred or determined ages that overlap ages of underlying dome gneisses; and (4) chemical compositions suggestive of island arc (e.g., Schumacher, 1988) or back-arc basin (Hollocher, 1993) origin. The contact between the Ammonoosuc Volcanics and the Partridge Formation with the underlying dome gneisses has had different interpretations, including an unconformity (Robinson, 1979; Schumacher, 1988), an intrusive contact (Leo et al., 1984; Leo, 1985, 1991), and a fault (Robinson and Tucker, 1996; Kohn and Spear, 1999). The along strike continuity of the internal stratigraphy of the Ammonoosuc Volcanics is strong evidence against the contact being intrusive (Schumacher, 1988). Zircon ages of the volcanics overlap those of the underlying gneisses (Tucker and Robinson, 1990), evidence against the contact being an unconformity. The Ammonoosuc Volcanics–Monson Gneiss contact is quite sharp (<1–2 m), but where well exposed it is a zone of anomalously high strain. It seems likely that the contact at the top of the dome gneisses is probably a low angle fault. It was suggested as a Silurian extensional detachment by Robinson and Tucker (1996) and as a Devonian thrust by Kohn and Spear (1999). The Ordovician and older rocks are overlain along an early Silurian unconformity by Silurian and Devonian rocks of mostly sedimentary origin. The Silurian units are relatively thin in the BHA, and include quartzite and calcareous rocks having shallow water sedimentary protoliths. The Devonian rocks are also relatively thin, though the top of the Devonian is not exposed in the

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BHA. These units thicken dramatically eastward into the Merrimack–central Maine synclinorium (Moench and Boudette, 1970, 1987; Hatch et al., 1983; Zen, 1983; Thompson, 1985). All of these rocks were deformed and metamorphosed during the Devonian Acadian Orogeny and later (Robinson et al., 1998), and no pre-Devonian metamorphism has been recognized in the BHA.

3. Field and analytical methods The gneisses exposed in the BHA are strongly deformed and metamorphosed, and in most cases original inhomogeneities such as xenoliths and dikes have been smeared out to form layers that extend across entire outcrops. Clear identification of the original structural relationships between different lithologies is rarely possible. Where the rocks are less strained, igneous features including dikes and angular xenoliths are visible. In addition, many of these rocks, particularly the felsic Monson Gneiss, were partially melted during Acadian metamorphism (Hollocher, 1988). The solidified melts occur as centimeter- to meter-scale dikes, sills, and irregular bodies that are generally coarser grained and more leucocratic than the host felsic gneisses, but have identical mineralogy. All these features make collection of representative homogeneous samples difficult. For geochemical and petrographic purposes the most important concern was to sample fresh, visibly homogeneous rock, and to avoid weathered rock, mixed rock, quartz veins, joints, and metamorphic partial melts. To this end, 2–10 kg samples of fresh gneiss, amphibolite, and ultramafic rocks were collected from relatively thick, homogeneous layers or masses, generally from clean outcrops where subtle differences in grain size and color could be distinguished. Sample preparation and analysis procedures are described by Hollocher (1993), and Bull (1997).

4. Petrology and field relations 4.1. Fourmile Gneiss The Fourmile Gneiss ranges from well foliated to massive tonalitic and granodioritic gneisses, with smaller quantities of granitic gneiss. The Fourmile Gneiss is cut by numerous pegmatite dikes (probably Acadian) in the southern part of the Pelham dome. Felsic rocks of the Fourmile Gneiss were divided by Bull (1997) into five units: muscovite-rich gneiss, microcline-rich gneiss, muscovite-poor gneiss, Tailrace gneiss (sampled and analyzed by Hodgkins, 1985), and northern gneiss. The common assemblage of these units is quartz–plagioclase–microcline–biotite–Fe–Ti oxides  muscovite  hornblende  titanite  garnet.

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Secondary chlorite, epidote, and sericite are common but together generally amount to <1%. Hornblende is found only in the northern gneiss, kyanite is found in one muscovite-rich sample (3), and cummingtonite is found in one muscovite-poor sample (24). Muscovite is absent from the Tailrace and hornblende gneisses. The muscovite-rich and microcline-rich gneisses are yellow-weathering, and are restricted to the lower part of the Fourmile Gneiss in the southern part of the Pelham dome. These rocks are interpreted by Bull (1997) to have undergone pre-metamorphic hydrothermal alteration that made them anomalously peraluminous by removal of Ca and Na. The mostly gray-weathering muscovite-poor gneiss is restricted to the upper part of the unit in the southern part of the Pelham dome. The contact is gradational between the lower muscovite-rich and microcline-rich gneisses and the upper muscovitepoor gneiss. The gray-weathering Tailrace and northern gneisses are restricted to the northern part of the Pelham dome. The muscovite-poor, Tailrace, and northern gneisses are interpreted by Bull (1997) to be have nominally original igneous compositions. Amphibolites are most abundant in the lower part of the Fourmile Gneiss, interlayered with the muscoviterich and microcline-rich gneisses. Garnet- or gedritebearing amphibolites have been observed in the field and are interpreted to be mafic igneous rocks that were hydrothermally altered before metamorphism, based on criteria of Hollocher (1985a,b) and Schumacher (1988). In the upper part of the Fourmile Gneiss amphibolites are rare and gedrite is absent. None of the sampled Fourmile amphibolites (Table 1) are interpreted as having undergone significant alteration. 4.2. Monson Gneiss The Monson dome is largely composed of grayweathering tonalitic to granodioritic gneisses. The rocks range from homogeneous on scales of tens of meters, to strongly layered on a centimeter scale. Most rocks are weakly foliated (Fig. 3). Some of the less deformed rocks have clear igneous structural features including xenoliths and dikes. In some outcrops relatively massive rocks having xenoliths, dikes, or boudins grade into highly deformed rock with strong compositional layering, demonstrating that, at least in some areas, layered gneisses are strained plutonic rocks rather than volcanics. In support of plutonic origin, no rocks having unambiguous sedimentary protoliths have been found in the Monson or Fourmile Gneisses. The only possible exceptions are rare layers or boudins of epidote-rich calc-silicate rock in the Monson Gneiss that may have been hydrothermal veins. One of us (Robinson) interprets layering in a calc-silicate on Parker Island in the Quabbin Reservoir (Fig. 2) to be possible sedimentary bedding.

Fig. 3. Field photographs showing typical Monson Gneiss. (a) Typical massive Monson Gneiss, showing a xenolith of light-colored biotite tonalite gneiss (sample 100) in the darker-colored hornblende–biotite tonalite gneiss (sample 95). (b) Fine-grained, thick amphibolite (sample 80) in layered felsic gneiss and amphibolite. The thick amphibolite crosscuts compositional layering and is therefore a dike.

The Monson tonalitic gneisses have the assemblage plagioclase–quartz–biotite–K-feldspar–Fe–Ti oxides  hornblende  garnet. The hornblende–biotite gneisses are generally the most mafic, biotite gneisses are intermediate, and garnet–biotite gneisses are the most felsic. Hornblende and garnet are mutually exclusive. Other felsic rock types are rare. One outcrop contained a white, 2 m thick body of alaskite gneiss having the assemblage microcline–quartz–plagioclase–biotite–magnetite–garnet, with a color index of only
2% (sample 118, not plotted or discussed further below). One body of gray microcline augen gneiss is exposed on the east side of Mt. Zion Island and the island between the baffle dams in the Quabbin Reservoir (sample 117). This body is at least 1 km long and P 10 m thick, and has the assemblage quartz–microcline–plagioclase–biotite– hornblende–Fe–Ti oxides, with secondary epidote. Most of the felsic rocks examined in this study are very fresh

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

and generally contain <1% secondary minerals including sericite, chlorite, and epidote. The mafic rocks are largely amphibolites with the assemblage plagioclase–hornblende–Fe–Ti oxides  augite  quartz  biotite  titanite, with minor secondary sericite, chlorite, and epidote. Garnet, cummingtonite, and orthoamphiboles are absent in the mafic Monson rocks, in contrast to their common occurrence in the overlying Ammonoosuc Volcanics and volcanics in the Partridge Formation (Schumacher, 1988; Hollocher, 1993). The Monson amphibolites usually occur as layers, boudins, and crosscutting dikes in the felsic rocks (Fig. 3(b)). These bodies are typically centimeters to a meter thick, although irregular bodies up to a few tens of meters across are found. Most of the amphibolites have grain sizes of
1 mm, and some have patches of pure polycrystalline plagioclase or single plagioclase crystals up to 1 cm long that are probably remnants of plagioclase phenocrysts. Other amphibolite bodies are coarse-grained with typical grain sizes of 3–5 mm. None of the fine-grained amphibolites contain augite, but most of the coarse amphibolites do. Coarse- and finegrained amphibolites are commonly in close proximity in the outcrop. Other mafic to ultramafic rocks that occur in smaller amounts include hornblende pyroxenite, gabbroic anorthosite, olivine hornblendite, and hornblendite. Hornblende pyroxenite occurs in bodies up to several meters across, commonly associated with amphibolite. Gabbroic anorthosite was found as a localized set of large boulders on the east side of the Quabbin Reservoir, and as layers up to
1 m thick associated with amphibolite and tonalitic gneiss. It is similar in appearance to some tonalite, but it contains augite and has a slightly bluish color that is particularly visible on cloudy days. In thin section the anorthosite commonly has visible patches of
1 lm thick plagioclase exsolution lamellae in plagioclase host crystals. Hornblendite occurs as bodies up to several meters across, but more commonly as smaller boudins and xenoliths, small patches in amphibolite, and as selvages surrounding hornblende pyroxenite and amphibolite where these are in contact with felsic gneiss. A large body of olivine– orthopyroxene–augite hornblendite is also known (sample C36 of Tracy et al., 1984). 4.3. Other dome gneisses Four other samples were collected and analyzed from elsewhere in the BHA: granitic gneiss (sample 119) from the Owl’s Head dome, a granitic gneiss (120) from the Warwick dome, and a coarse-grained amphibolite (sample 121) and a tonalitic gneiss (sample 122) from the Keene Dome. These samples are included as appropriate in the figures and text, but will not be discussed further.

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4.4. Summary of the BHA dome gneisses The gneisses in domes elsewhere in the BHA are described in detail by others (Billings and Wilson, 1965; Foland and Loiselle, 1981; Pogorzelski, 1983; Leo et al., 1984; Leo, 1985, 1991; Webster and Wintsch, 1987). These gneisses are overwhelmingly dominated by massive to weakly foliated or layered felsic rocks that range in composition from tonalite to granite. Mafic rocks are subordinate, and ultramafic rocks are rare. The whole BHA dome gneiss suite is calc-alkaline with most felsic gneisses being weakly peraluminous or metaluminous, with generally <5% normative diopside and <2% normative corundum. The dome gneiss suite is therefore plutonic, calc-alkaline, bimodal, dominated by felsic rocks, and inhomogeneous on a variety of scales.

5. Geochemistry 5.1. Overview Data for the 122 new analyzed samples are given in Tables 1–3. Fig. 4 is an AFM diagram showing the new analyses along with other BHA dome gneisses and Partridge Formation volcanics for comparison. Fourmile and Monson amphibolites straddle the boundary between tholeiitic and calc-alkaline fields, but predominate on the calc-alkaline side of the boundary in contrast to the tholeiitic Partridge amphibolites. Most of the ultramafic rocks, one hornblendite and four hornblende pyroxenites are Mg-rich and are probably cumulates. All plotted felsic rocks are calc-alkaline and extend from relatively mafic, alkali-poor compositions dominated by hornblende tonalitic gneisses with
61% SiO2 , to felsic compositions dominated by alkali-rich biotite granite gneisses with
78% SiO2 . A striking feature of this diagram is the composition gap between mafic and felsic rocks, demonstrating the bimodal distribution of compositions that is obvious in the field, and strong evidence that the mafic and felsic rocks are not cogenetic. Fig. 5 is a K2 O–SiO2 diagram showing all available data for BHA dome gneisses. The Fourmile and Monson rocks lie mostly in the tholeiitic and calc-alkaline series, as do almost all samples from the southern domes (south of the Mascoma and Lebanon domes; Fig. 1). Gneisses in the northern domes are mostly in the high-K calc-alkaline series and include examples of alkali granite and syenite gneisses. 5.2. Amphibolites The mafic rocks have experienced kyanite and sillimanite grade metamorphism, so pre- and synmetamorphic chemical alteration is a concern. Fig. 6 shows all

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 1 Chemical analyses of rocks in the Fourmile Gneiss Sample

Amphibolites 56

57

Hornblende gneisses 58

60

59

62

61

48

49

55

50.95 0.46 16.02 2.64 9.84 0.23 6.25 11.37 1.79 0.52 0.03

50.99 0.71 18.10 1.01 8.20 0.18 6.46 9.75 3.67 0.48 0.13

51.07 0.95 17.00 1.56 7.39 0.20 7.38 10.54 3.36 0.32 0.13

51.65 1.04 14.35 3.40 10.01 0.31 5.51 10.10 1.81 1.00 0.14

51.81 0.60 17.07 1.59 9.93 0.31 6.17 7.97 4.16 0.27 0.05

52.01 0.42 15.19 2.92 7.38 0.23 8.70 8.90 3.57 0.44 0.07

55.66 0.69 15.50 2.97 7.97 0.23 5.16 7.13 4.51 0.22 0.06

63.37 0.50 15.73 0.40 5.99 0.18 2.83 5.23 2.83 2.59 0.08

63.75 0.51 15.73 0.47 4.71 0.14 2.76 6.32 4.10 0.87 0.08

67.14 0.41 15.18 0.53 4.32 0.18 1.22 4.16 3.55 2.71 0.09

Total Mg=ðMg þ Fet Þ Feþ3 =Fet Ca/(Ca + Na)

100.11 0.48 0.19 0.78

99.67 0.56 0.10 0.60

99.91 0.60 0.16 0.63

99.31 0.43 0.23 0.76

99.93 0.49 0.13 0.51

99.83 0.61 0.26 0.58

100.12 0.46 0.25 0.47

99.73 0.44 0.06 0.50

99.45 0.49 0.08 0.46

99.48 0.31 0.10 0.39

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb

419 79 29 97 15 10.6 117 9.6 13 1.11

264 87 35 69 16 6.7 217 14.7 31 3.76

397 0 12 83 18 3.5 317 23.9 64 3.63

448 3 26 114 15 22.3 199 21.6 41 0.55

316 27 27 79 14 5.1 175 13.7 18 0.94

240 150 46 76 17 13.2 165 12.0 23 1.11

296 10 15 76 15 7.4 145 17.5 32 1.21

169 21 10 81 15 93.7 163 22.7 100 5.92

165 22 14 46

630

304

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5

Ba

58

19

53

222

36

59

85

22.2 247 25.5 110 5.60

La Ce Pr Nd Sm Eu

3.01 4.9 0.65 3.3 0.91 0.32

4.29 7.7 1.10 5.8 1.77 0.68

11.14 26.1 3.79 16.8 3.82 0.89

8.85 16.6 2.18 9.9 2.62 0.83

6.18 8.4 1.54 6.7 1.78 0.55

4.48 9.6 1.17 5.5 1.58 0.55

2.24 5.3 0.89 4.8 1.75 0.64

30.66 56.1 6.54 24.5 4.46 0.82

25.62 54.4 6.91 27.9 5.59 1.00

Gd Dy Er Yb Lu

1.18 1.45 0.93 0.97 0.12

2.14 2.67 1.65 1.55 0.23

4.20 4.70 2.89 2.87 0.46

3.05 3.83 2.42 2.33 0.35

2.16 2.52 1.45 1.39 0.23

1.75 2.13 1.34 1.37 0.21

2.32 3.10 2.05 2.03 0.34

4.22 3.92 2.41 2.36 0.38

5.18 4.79 2.91 2.79 0.43

Hf Pb Th U

0.39 7.4 0.49 0.50

0.99 19.6 0.78 0.85

2.17 12.6 2.44 0.74

1.42 5.6 1.60 0.74

0.56 11.2 0.94 1.03

0.78 5.1 1.79 0.50

1.16 3.9 1.06 0.39

3.38 15.3 9.52 2.18

3.60 17.8 9.20 3.12

15 1.3 1.1

13 2.4 1.2

9 1.8 1.1

13 1.6 1.3

8 0.7 0.9

7 6.1 1.4

10 5.0 1.5

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

12 1.3 1.0

14 1.8 1.0

86 7 14 94

144 126 7.51

24 55

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

13

Table 1 (Continued) Sample

Hornblende gneisses 53

51

Microcline-rich gneisses

54

52

50

18

21

19

20

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5

67.37 0.45 15.52 0.55 4.44 0.13 1.40 4.37 3.40 2.11 0.11

67.47 0.34 15.22 0.51 4.14 0.10 1.72 4.28 2.71 2.88 0.07

67.61 0.42 15.38 0.54 4.39 0.12 1.53 4.12 3.41 2.42 0.09

68.09 0.32 15.03 0.47 3.84 0.09 1.74 3.55 2.82 3.33 0.07

71.73 0.26 13.74 0.35 2.83 0.07 1.20 2.98 2.78 3.51 0.04

76.66 0.07 13.00 0.36 0.50 0.04 0.42 0.56 3.11 5.32 0.03

78.27 0.07 12.01 0.23 0.60 0.02 0.50 0.53 2.38 5.14 0.02

78.47 0.07 11.92 0.12 0.97 0.05 0.31 0.40 1.75 5.95 0.02

79.05 0.07 12.04 0.20 0.59 0.05 0.24 0.61 2.05 5.47 0.01

Total Mg=ðMg þ Fet Þ Feþ3 =Fet Ca/(Ca + Na)

99.85 0.34 0.10 0.41

99.45 0.40 0.10 0.47

100.03 0.36 0.10 0.40

99.34 0.42 0.10 0.41

99.49 0.40 0.10 0.37

100.06 0.47 0.39 0.09

99.77 0.53 0.26 0.11

100.02 0.34 0.10 0.11

100.39 0.36 0.23 0.14

86 12 4 68

97 15 7 40

86 6 4 64

88 13 7 36

61 12 6 38

195

209

181

245

130

136 7.25

105 5.44

115 7.15

2 0 3 26 10 147.4 52 16.1 74 5.86

2 0 6 26 9 142.6 65 19.5 100 3.65

2 0 8 22 9 134.7 87 10.7 85 4.47

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb

98 5.62

119 6.89

Ba La Ce Pr Nd Sm Eu

739 16 31

24 50

16 25

25 46

33 10 122.8 82 13.0 93 4.01 1001

829

1194

25.34 48.1 5.18 18.3 3.32 0.48

23.56 46.5 5.25 19.9 3.94 0.59

29.62 56.3 6.27 22.6 4.05 0.65

28.77 51.4 5.49 19.6 3.32 0.64

Gd Dy Er Yb Lu

3.09 2.97 2.15 2.55 0.41

3.36 2.57 1.42 1.25 0.17

3.69 3.78 2.67 2.88 0.48

2.71 2.16 1.52 2.00 0.33

Hf Pb Th U

2.42 45.3 8.48 2.52

2.52 51.9 8.32 2.02

3.12 32.3 8.35 1.68

2.72 39.6 8.37 2.39

3 4.9 1.0

6 9.6 2.2

3 5.1 1.0

8 6.7 1.1

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

26 67

1 0

14

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 1 (Continued) Sample

Muscovite-poor gneisses 63

38

22

36

23

26

32

33

39

24

40

74.61 0.16 14.10 0.67 1.50 0.08 0.61 2.59 3.93 2.15 0.04

75.74 0.16 13.37 0.54 1.48 0.05 0.75 2.23 4.28 1.32 0.04

76.13 0.10 13.73 0.15 1.13 0.08 0.42 2.27 4.42 1.59 0.02

76.31 0.19 13.67 0.16 1.74 0.10 0.83 1.72 4.87 0.83 0.04

76.58 0.13 12.97 0.38 1.33 0.06 0.41 1.87 3.77 2.58 0.04

76.89 0.12 13.05 0.71 0.99 0.06 0.48 1.87 3.84 2.46 0.03

77.12 0.10 13.05 0.21 1.10 0.06 0.44 1.91 3.62 2.55 0.02

77.15 0.10 12.97 0.31 0.79 0.04 0.55 1.71 4.49 1.18 0.01

77.26 0.12 12.87 0.63 1.01 0.05 0.41 1.92 3.85 2.34 0.02

77.33 0.17 12.53 0.83 0.70 0.03 0.87 1.80 4.90 0.28 0.02

77.44 0.09 12.52 0.13 1.19 0.05 0.49 1.07 4.44 2.23 0.01

99.96 0.41 0.25 0.22

100.04 0.37 0.11 0.22

100.46 0.44 0.08 0.16

100.12 0.31 0.21 0.22

100.49 0.34 0.39 0.21

100.18 0.38 0.15 0.23

99.31 0.48 0.26 0.17

100.50 0.32 0.36 0.22

99.46 0.52 0.52 0.17

99.67 0.40 0.09

0.26

100.42 0.34 0.29 0.27

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb

16 1 7 35 19 106.0 342 2.3 71 2.50

28 0 7 42 13 94.7 105 10.3 70 4.05

20 0 5 36 12 58.3 118 5.3 85 5.51

10 0 4 36 10 72.6 74 8.8 76 4.53

7 0 5 21 13 54.2 90 5.9 106 4.96

14 0 5 13 11 6.9 159 16.9 112 1.89

3 0 4 61 12 80.6 56 26.6 80 6.94

Ba

860

294

171

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5 Total Mg=ðMg þ Fet Þ Feþ3 =Fet Ca/(Ca + Na)

73.73 0.15 15.23

27

1.17 0.03 0.40 2.74 4.35 2.35 0.05 100.20 0.38

27 1 36 11 52.5 79 16.6 72

373

22 0 4 19 12 48.5 115 19.8 77 3.31

12 0 23 12 86.0 110 9.7 90 3.41

13 1 4 21 11 78.5 149 19.6 88 3.77

171

450

498

669

96

14 0 4 20 79.0 169 11.1 83 3.05 532

69

373

La Ce Pr Nd Sm Eu

10.28 17.6 1.96 7.3 1.48 0.40

20.20 42.1 3.27 11.0 2.02 0.39

10.94 14.0 2.64 9.6 1.81 0.35

41.60 39.7 8.26 28.5 4.69 0.62

19.63 38.1 4.41 17.8 3.87 0.71

20.82 46.3 4.02 14.1 2.49 0.42

32.81 66.7 6.76 24.8 4.63 0.79

4.81 6.3 0.76 2.5 0.59 0.26

5.30 9.4 1.12 1.12 0.85 0.44

20.95 44.6 4.63 4.63 3.00 0.51

20.84 42.6 5.22 5.22 4.29 0.89

26.34 52.1 6.17 4.1 5.11 0.42

Gd Dy Er Yb Lu

1.30 0.88 0.34 0.30 0.05

2.02 1.78 1.20 1.45 0.25

1.40 1.12 0.62 0.65 0.11

4.21 3.09 1.96 2.75 0.49

4.51 6.08 4.55 5.25 0.82

2.27 1.94 1.61 2.47 0.46

4.61 4.11 2.05 1.73 0.28

0.50 0.94 1.11 1.70 0.32

0.91 1.13 0.85 0.99 0.17

2.41 2.04 1.48 2.01 0.37

4.20 3.80 2.39 2.51 0.43

6.08 8.79 6.32 6.38 1.04

Hf Pb Th U

2.45 3.7 3.71 0.94

2.18 21.8 10.26 2.07

3.02 17.0 6.22 1.09

2.49 23.0 8.69 0.94

3.96 16.6 6.32 3.59

4.00 15.3 8.17 1.36

3.23 20.8 8.56 1.32

3.47 30.0 10.24 1.25

4.53 19.6 1.34 1.84

3.82 24.0 6.72 1.18

3.68 3.9 6.78 1.26

3.59 20.5 8.69 2.64

10 7.5 1.1

22 5.6 1.7

5 3.7 1.2

6 1.9 0.7

11 4.9 0.7

8 10.0 2.1

8 1.0 0.2

15 2.5 0.7

15 5.7 1.0

9 4.4 1.4

2 2.1 0.8

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

149 15.1 3.5

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

15

Table 1 (Continued) Sample

Muscovite-poor gneisses 42

35

34

29

28

30

41

37

43

31

25

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5

77.58 0.06 12.73 0.38 0.68 0.06 0.17 1.05 4.24 2.97 0.01

77.81 0.14 12.43 0.84 1.19 0.11 0.39 2.15 3.85 1.16 0.02

77.83 0.10 12.80 0.33 0.70 0.03 0.55 1.72 4.63 1.09 0.01

77.89 0.07 12.81 0.19 1.00 0.06 0.26 1.90 4.11 1.82 0.03

77.90 0.08 12.64 0.13 0.90 0.04 0.56 1.75 4.58 1.06 0.01

78.02 0.08 13.20 0.06 0.68 0.02 0.37 1.82 4.83 1.16 0.01

78.09 0.06 12.51 0.52 0.55 0.09 0.23 1.08 4.44 2.11 0.01

78.10 0.04 12.75 0.33 0.89 0.06 0.62 2.81 2.47 2.29 0.01

78.16 0.12 12.19 0.08 1.00 0.03 0.94 0.54 5.92 0.74 0.02

78.33 0.17 12.23 0.68 1.30 0.04 0.67 1.79 4.49 0.70 0.02

78.45 0.20 11.66 1.00 1.38 0.04 0.61 2.36 3.59 0.57 0.02

Total Mg=ðMg þ Fet Þ Feþ3 =Fet Ca/(Ca + Na)

99.93 0.23 0.33 0.12

100.09 0.26 0.39 0.24

99.80 0.50 0.30 0.17

100.14 0.28 0.15 0.20

99.66 0.50 0.12 0.17

100.25 0.48 0.07 0.17

99.70 0.29 0.46 0.12

100.38 0.48 0.25 0.39

99.74 0.61 0.07 0.05

100.42 0.38 0.32 0.18

99.89 0.32 0.39 0.27

9 0

10 0 2 21 13 51.0 87 4.6 111 4.90

6 0 0 21 11 49.8 114 10.8 131 2.08

13 0

5 0

3 0 5 55 12 68.8 44 28.4 78 5.05

0 0 3 52 11 97.4 60 18.9 83 7.46

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb Ba

2 0 23 12 80.3 47 29.1 91 3.62 509

43 11 46.4 92 19.0 107 3.73 232

84

356

31 13 54.5 83 15.7 104 5.97 75

21 43.3 122 50.2 90 5.53 169

363

635

6 0 25 11 22.9 38 26.7 121 3.22 93

10 0

10 0

18 10 17.0 93 9.9 106 4.55

11 10 18.5 145 9.5 119 3.44

319

195

La Ce Pr Nd Sm Eu

14.19 32.4 3.19 11.7 2.69 0.53

16.63 35.7 3.89 14.6 2.98 0.73

5.79 6.0 0.70 2.5 0.54 0.40

7.27 11.6 1.43 5.4 0.99 0.40

3.64 4.4 0.50 2.0 0.64 0.41

14.47 33.1 3.67 15.0 3.46 0.58

24.19 48.7 6.06 23.3 4.72 0.44

17.56 25.5 3.87 14.6 3.01 0.37

24.01 45.6 5.70 23.2 5.04 0.72

12.38 24.8 2.58 9.2 1.70 0.42

6.18 11.6 1.17 4.7 0.83 0.47

Gd Dy Er Yb Lu

3.48 6.59 5.52 6.50 1.05

3.04 3.43 2.64 3.20 0.50

0.61 0.88 0.70 0.85 0.15

0.90 1.45 1.93 3.03 0.50

1.30 2.74 2.34 2.53 0.41

4.65 7.26 5.54 6.00 0.97

4.73 4.75 3.09 3.65 0.57

2.81 2.88 2.22 2.61 0.42

5.34 6.17 4.06 4.33 0.71

1.69 1.62 1.42 1.97 0.35

0.86 1.38 1.40 1.79 0.28

Hf Pb Th U

4.22 18.4 9.44 3.66

3.59 15.8 4.32 1.20

4.31 19.3 0.83 1.60

3.97 25.1 7.65 0.83

4.40 19.5 1.57 1.73

3.84 28.3 9.86 3.45

2.91 16.8 7.08 2.38

2.80 22.4 10.02 2.45

5.24 6.6 7.18 2.32

3.69 7.8 5.10 1.84

3.02 7.8 2.92 1.66

2 1.3 0.4

5 2.9 0.8

19 1.8 0.6

10 1.0 0.2

5 0.5 0.4

2 1.4 0.6

2 3.5 1.0

3 2.5 0.9

1 2.7 1.0

9 3.3 0.7

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

15 1.7 0.4

16

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 1 (Continued) Sample

Muscovite-rich gneisses 7

2

13

11

12

6

4

8

15

5

14

74.83 0.19 14.16 0.20 2.01 0.03 1.33 1.07 4.54 1.89 0.04

76.33 0.22 13.14 0.63 1.09 0.05 0.88 1.64 4.79 0.89 0.03

76.53 0.09 13.36 0.10 1.30 0.09 0.40 1.41 3.49 3.40 0.02

77.43 0.05 13.76 0.05 0.73 0.02 0.56 0.90 3.96 1.90 0.11

77.43 0.17 12.64 0.66 1.35 0.05 0.65 1.50 4.76 1.11 0.03

77.75 0.16 12.34 0.31 1.19 0.04 0.90 1.26 4.02 1.48 0.02

77.89 0.07 12.33 0.23 0.79 0.08 0.21 1.32 3.62 2.79 0.01

78.16 0.12 13.20 0.33 0.99 0.01 0.54 1.17 4.59 1.25 0.03

78.18 0.07 12.38 0.49 0.80 0.07 0.15 1.29 4.07 2.60 0.01

78.19 0.08 12.68 0.25 0.79 0.07 0.30 1.35 4.41 1.58 0.02

78.23 0.05 12.67 0.14 0.90 0.08 0.32 0.71 3.82 3.36 0.02

Total Mg=ðMg þ Fet Þ Feþ3 =Fet Ca/(Ca + Na)

100.29 0.52 0.08 0.12

99.69 0.49 0.34 0.16

100.17 0.34 0.06 0.18

99.45 0.56 0.06 0.11

100.36 0.37 0.31 0.15

99.46 0.52 0.19 0.15

99.34 0.27 0.21 0.17

100.40 0.43 0.23 0.12

100.12 0.18 0.36 0.15

99.72 0.34 0.22 0.14

100.29 0.35 0.12 0.09

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb

25 0 2 60 13 84.4 143 13.1 89 4.23

22 0 5 21 12 24.3 107 21.3 94 2.94

2 0 6 56 12 101.6 61 19.0 91 5.32

2 0 1 27 11 73.4 64 27.6 95 2.65

4 0 4 44 10 47.2 222 8.6 114 5.27

2 0 5 25 12 41.5 75 17.1 80 4.43

3 0 7 59 11 105.9 36 28.8 82 5.73

Ba

310

177

783

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5

6 0 20 12 56.4 97 10.1 58 5.10 328

11 0 6 24 12 62.1 70 20.6 104 4.68

9 0 24 11 51.4 66 21.1 109 4.04

178

162

644

352

3 0 6 30

65 94 3.33 603

268

565

La Ce Pr Nd Sm Eu

16.80 33.7 3.56 12.5 2.21 0.62

17.05 34.7 4.09 15.6 3.17 0.56

35.42 54.5 7.51 26.4 4.79 0.58

17.96 34.2 3.80 13.7 2.79 0.43

18.78 35.7 4.23 17.2 3.65 0.52

16.28 35.5 4.08 16.2 3.30 0.58

7.96 14.2 1.80 6.8 1.64 0.38

48.78 86.5 8.93 30.0 4.23 1.22

16.81 33.6 4.17 15.4 3.11 0.56

28.12 54.1 6.11 22.7 4.09 0.49

26.19 54.1 6.26 24.2 4.76 0.27

Gd Dy Er Yb Lu

2.26 2.14 1.47 1.57 0.25

3.23 3.41 2.34 2.70 0.43

3.99 3.99 2.58 2.95 0.48

2.45 2.17 1.21 1.20 0.20

3.86 3.94 2.49 2.80 0.50

3.58 3.58 2.28 2.58 0.40

2.15 4.05 3.81 4.79 0.83

3.86 2.05 0.99 0.95 0.13

2.80 3.41 2.86 3.62 0.61

3.81 3.36 2.69 3.72 0.68

5.10 5.98 4.40 4.77 0.80

Hf Pb Th U

2.49 55.1 6.02 1.85

3.39 13.2 5.82 2.55

3.22 29.5 12.62 3.46

2.49 52.6 5.98 1.48

3.63 10.1 6.85 2.31

2.87 18.7 5.78 2.29

3.82 23.7 6.23 1.51

3.19 44.8 11.56 1.80

3.32 6.48 1.00

2.91 39.4 8.93 1.65

3.47 28.4 10.81 2.22

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

11 5.6 1.2

5 3.3 1.0

3 4.8 1.1

10 7.4 1.6

3 3.3 1.1

3 3.6 1.1

2 0.8 0.4

26 23.7 3.3

2.4 0.6

4 3.8 0.8

1 2.9 0.9

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

17

Table 1 (Continued) Sample

Muscovite-rich gneisses 9

17

3

Tailrace gneisses 10

16

1

44

45

46

47

SiO2 TiO2 Al2 O3 Fe2 O3 FeO MnO MgO CaO Na2 O K2 O P2 O5

78.98 0.09 12.44 0.14 0.89 0.03 0.30 0.75 5.82 0.42 0.01

79.07 0.07 11.84 0.16 0.80 0.03 0.45 0.87 3.57 2.46 0.02

79.13 0.21 12.52 0.05 0.79 0.02 0.96 0.79 4.62 0.82 0.03

79.20 0.05 12.72 0.05 0.75 0.01 0.50 0.96 3.86 1.82 0.19

79.34 0.11 11.97 0.38 0.69 0.02 0.60 1.27 4.91 0.67 0.02

81.13 0.06 11.69 0.05 0.75 0.01 0.44 0.68 4.03 1.50 0.03

69.00 0.36 16.20 0.80 2.40 0.06 0.91 4.18 4.01 1.92 0.11

69.70 0.28 15.30 0.60 2.80 0.07 1.18 3.75 3.27 2.51 0.09

69.70 0.29 15.70 0.70 2.70 0.12 1.20 3.84 3.30 2.44 0.11

71.50 0.23 15.10 1.00 1.70 0.07 0.75 3.16 3.36 2.98 0.08

Total Mg=ðMg þ Fet ) Feþ3 =Fet Ca/(Ca + Na)

99.87 0.35 0.12 0.07

99.34 0.46 0.15 0.12

99.94 0.67 0.05 0.09

100.12 0.53 0.06 0.12

99.98 0.51 0.33 0.12

100.37 0.50 0.06 0.08

99.95 0.34 0.23 0.37

99.55 0.39 0.16 0.39

100.10 0.39 0.19 0.39

99.93 0.34 0.35 0.34

Trace elements (ppm) V Cr Ni Zn Ga Rb Sr Y Zr Nb

1 0 5 11 11 7.6 35 33.4 129 4.14

4 0 11 32 10 72.3 79 16.4 84 5.01

14 0 2 17 9 23.2 50 15.5 100 4.78

1 0

1 0

28 11 58.2 147 13.4 77 4.60

6 0 4 15 10 30.8 78 26.0 124 3.92

41 5 5 31 14 61.0 336 9.0 166 6.00

62 5 6 28 14 90.0 274 13.0 108 8.00

56 6 2 54 15 111.0 306 14.0 102 9.00

44 4 5 30 13 94.0 271 13.0 118 8.00

Ba

188

265

352

402

86

26 9 48.1 140 12.0 70 3.19 306

1168

1173

1146

1158

La Ce Pr Nd Sm Eu

21.86 47.3 5.57 21.4 4.28 0.37

23.71 47.6 5.07 17.5 3.11 0.44

16.50 34.4 3.89 14.8 2.77 0.44

28.71 54.8 6.05 22.4 4.22 0.77

18.08 35.9 4.41 17.6 3.92 0.64

27.60 51.3 5.49 19.5 3.17 0.43

22.07 36.8 3.97 14.5 2.37 1.13

36.04 52.6 5.54 19.2 3.24 0.69

33.81 57.5 6.01 21.1 3.49 0.75

29.06 47.0 4.85 16.3 2.71 0.62

Gd Dy Er Yb Lu

4.58 5.71 4.34 5.10 0.80

3.02 2.68 1.70 1.77 0.30

2.67 2.67 1.68 1.86 0.28

3.73 3.02 1.46 1.37 0.22

4.44 5.26 3.49 3.57 0.57

2.89 2.17 1.39 1.46 0.25

2.07 1.61 0.99 1.13 0.19

2.73 2.34 1.41 1.51 0.25

2.96 2.60 1.65 1.86 0.31

2.39 2.22 1.48 1.62 0.28

Hf Pb Th U

4.41 6.2 7.85 1.80

2.55 36.5 9.70 2.75

3.26 7.5 5.63 1.22

3.12 47.8 9.73 2.08

4.12 10.5 6.84 2.04

2.66 61.4 9.63 2.10

4.76 13.0 5.28 1.91

3.53 18.0 15.63 2.79

3.63 30.0 18.73 3.87

3.97 29.0 15.65 4.61

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

1 2.4 0.7

5 7.0 1.4

3 4.8 1.2

11 10.4 2.2

3 2.6 1.0

12 9.1 1.6

37 8.4 1.5

21 9.0 1.5

22 8.0 1.3

21 7.5 1.2

18

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 2 Chemical analyses of rocks in the monson gneiss Sample

Hornblendite

Hornblende pyroxenite

Anorthosite

Low Nb amphibolite

64

66

69

70

65

67

68

71

72

73

74

75

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

41.44 1.69 12.48 21.10 0.25 7.95 10.97 1.41 0.89 0.01

49.03 0.38 9.80 10.38 0.26 15.34 11.73 1.26 0.47 0.01

51.81 0.31 10.65 6.96 0.17 11.99 16.01 1.30 0.29 0.02

52.64 0.31 9.39 6.00 0.16 12.98 16.54 1.57 0.27 0.02

53.26 0.20 8.48 5.52 0.16 13.59 16.48 1.39 0.25 0.01

48.75 0.12 26.24 2.92 0.07 4.12 15.65 1.93 0.26 0.01

45.95 1.13 17.01 12.55 0.24 7.27 12.24 1.88 0.46 0.12

46.56 1.02 15.11 13.57 0.25 8.02 11.32 2.03 0.77 0.09

48.13 2.13 15.96 11.31 0.21 7.16 10.18 3.81 0.31 0.25

48.29 0.86 17.60 10.06 0.19 7.38 12.32 1.96 0.50 0.14

48.85 0.70 16.72 10.01 0.21 8.03 12.76 1.59 0.33 0.06

49.41 1.22 18.44 7.90 0.15 6.40 11.80 3.25 0.72 0.12

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

98.19 0.40 0.81

98.66 0.72 0.84

99.51 0.75 0.87

99.88 0.79 0.85

99.34 0.81 0.87

100.07 0.72 0.82

98.85 0.51 0.78

98.74 0.51 0.75

99.45 0.53 0.60

99.30 0.57 0.78

99.26 0.59 0.82

99.41 0.59

Trace elements (ppm) Be 0.7 1.5 1.2 2.1 Sc 26 80 68 62 V 1264 224 221 228 176 Cr 10 1922 1314 2562 2292 Ni 24 485 202 228 269 Zn 145 129 75 58 66 Ga 20 14 11 8 8 Rb 8.2 3.8 2.5 2.3 3.6 Sr 36 12 70 68 56 Y 7.5 14.0 13.0 8.6 10.0 Zr 12 31 17 22 26 Nb 1.10 1.30 0.80 1.00 0.90

1.7 22 71 93 40 42 16 6.4 183 4.4 12 0.70

461 81 44 130 19 15.0 117 26.0 33 1.90

2.9 35 300 93 57 101 20 1.5 281 34.0 175 5.40

314 180 65 96 19 3.7 140 24.0 40 1.90

322 238 75 96 17 3.2 129 23.0 24 1.20

202 281 64 67 16 4.4 254 22.0 97 2.50

Mo Sn Cs Ba

458 68 49 123 20 3.6 150 24.0 24 1.00

0.05 128

0.14 0.80 0.06 44

0.11 0.24 0.05 44

0.08 0.25 0.03 73

0.09 0.27 0.07 85

0.09 0.22 0.54 90

0.48

La Ce Pr Nd Sm Eu

1.79 3.0 0.48 2.7 0.83 0.34

4.34 14.1 1.58 7.2 1.79 0.74

2.69 6.9 1.14 5.8 1.79 0.63

2.55 5.8 0.78 3.6 0.96 0.36

2.67 5.7 0.90 4.2 1.16 0.42

2.58 4.8 0.60 2.4 0.61 0.28

3.46 10.6 1.76 9.7 2.89 0.93

4.21 11.9 1.98 10.7 3.13 0.81

10.12 26.0 3.90 19.5 5.56 1.97

5.86 16.6 2.45 12.2 3.07 0.98

3.12 9.7 1.59 8.7 2.72 0.88

5.44 14.3 2.23 10.9 2.95 1.10

Gd Dy Er Yb Lu

1.12 1.26 0.77 0.74 0.11

1.95 2.06 1.33 1.55 0.26

2.22 2.43 1.52 1.48 0.24

1.21 1.41 0.87 0.91 0.15

1.45 1.68 1.05 1.11 0.18

0.69 0.78 0.48 0.52 0.09

3.79 4.14 2.69 2.55 0.39

4.08 4.33 2.88 2.74 0.40

6.20 6.15 3.48 3.30 0.53

3.68 3.99 2.51 2.47 0.38

3.63 4.07 2.58 2.49 0.37

3.83 4.12 2.41 2.19 0.34

Hf Pb Th U

0.37 3.0 0.28 0.66

0.86 2.2 0.33 0.19

0.60 3.3 0.31 0.28

0.74 3.0 0.98 0.20

0.61 2.8 0.42 0.09

0.31 9.2 0.21 0.20

0.62 5.0 0.04 0.07

1.16 5.0 0.17 0.26

1.39 5.2 0.93 0.23

1.43 7.0 0.17 0.11

0.84 3.0 0.03 0.01

1.45 9.0 0.62 0.20

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

4.8 1.05 1.23

0.9 2.35 1.02

5.4 1.21 1.21

7.9 1.64 1.07

5.6 1.33 1.05

41.6 2.39 1.07

6.3 1.08 1.20

4.5 1.12 1.20

8.3 2.04 1.52

5.8 1.74 1.20

5.6 1.01 1.18

11.5 1.69 1.41

76

112

0.06 106

68

34

69

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

19

Table 2 (Continued) Sample

Low Nb amphibolite 76

77

78

79

80

81

82

83

84

85

86

87

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

49.49 1.38 16.52 9.00 0.19 7.83 11.22 3.26 0.56 0.15

49.72 0.50 15.51 9.82 0.19 9.10 12.40 1.44 0.41 0.04

49.82 1.37 16.65 9.07 0.19 7.83 10.57 3.11 0.74 0.14

49.84 1.95 15.38 10.91 0.20 6.43 10.54 3.18 0.36 0.21

49.91 1.56 16.21 9.47 0.18 7.05 10.59 3.58 0.56 0.18

49.97 1.86 15.59 10.51 0.20 5.98 10.66 3.30 0.73 0.19

49.98 0.62 17.13 8.92 0.17 8.02 12.35 1.88 0.49 0.09

50.03 1.60 14.88 12.01 0.28 6.72 8.19 3.99 0.24 0.15

50.75 0.47 15.42 9.39 0.18 8.96 12.16 1.47 0.40 0.04

53.60 0.31 16.65 8.30 0.17 6.53 11.07 2.18 0.76 0.02

55.04 0.47 17.16 8.92 0.15 4.16 9.64 2.94 0.78 0.03

55.28 0.65 15.93 10.73 0.19 4.14 8.76 2.93 0.48 0.05

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

99.60 0.61 0.66

99.13 0.62 0.83

99.49 0.61 0.65

99.00 0.51 0.65

99.29 0.57 0.62

98.99 0.50 0.64

99.65 0.62 0.78

98.09 0.50 0.53

99.24 0.63 0.82

99.59 0.58 0.74

99.29 0.45 0.64

99.14 0.41

266 191 76 77 16 4.2 129 19.0 24 2.10

3.4 46 383 90 34 186 19 2.2 111 33.0 68 2.20

266 439 94 78 15 3.4 88 10.2 19 1.00

227 117 32 74 13 3.2 113 8.9 26 1.90

263 10 14 63 16 6.3 124 11.0 34 1.10

337 1 7 85 16 3.7 124 14.0 39 1.50

101

145

106

Trace elements (ppm) Be Sc V 213 Cr 268 Ni 87 Zn 99 Ga 17 Rb 3.1 Sr 273 Y 24.0 Zr 106 Nb 3.20

280 463 101 79 13 3.0 89 10.3 17 1.40

Mo Sn Cs Ba

97

La Ce Pr Nd Sm Eu

6.67 15.9 2.49 12.8 3.48 1.36

2.02 5.1 0.69 3.5 0.94 0.35

Gd Dy Er Yb Lu

4.08 4.49 2.59 2.40 0.37

Hf Pb Th U Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

215 268 95 89 19 9.2 213 22.0 109 3.10

272 192 64 92 20 6.6 222 37.0 162 3.90

236 230 69 82 18 2.2 233 29.0 133 3.80

276 200 43 93 19 3.5 177 35.0 134 3.40

96

59

57

48

0.20 1.34 0.12 57

8.25 19.3 2.82 13.9 3.76 1.46

8.26 20.5 2.96 14.4 4.04 1.32

7.92 18.6 2.78 14.2 3.75 1.27

7.93 20.9 3.31 16.6 4.69 1.72

5.54 14.9 2.28 9.9 2.51 0.77

8.38 21.0 3.09 15.1 4.37 1.51

2.25 5.3 0.82 3.9 1.15 0.45

5.63 11.3 1.32 5.7 1.22 0.40

4.93 10.4 1.37 6.1 1.59 0.50

3.45 9.0 1.30 6.2 1.69 0.57

1.27 1.57 0.99 1.05 0.15

4.25 4.50 2.71 2.50 0.39

5.07 6.02 3.46 3.27 0.49

4.88 4.86 2.91 2.71 0.40

5.74 6.46 3.79 3.56 0.55

2.84 3.19 1.98 1.96 0.31

5.28 5.82 3.50 3.52 0.58

1.48 1.78 1.12 1.11 0.18

1.39 1.52 0.99 1.06 0.17

1.84 2.06 1.41 1.45 0.23

2.16 2.47 1.57 1.62 0.26

2.55 10.0 0.47 0.20

0.43 4.0 0.55 0.09

2.77 15.0 0.81 0.68

3.74 5.0 1.09 0.25

3.23 4.0 0.94 0.28

3.91 7.0 0.86 0.29

0.95 5.0 0.19 0.17

1.35 3.3 0.81 0.28

0.58 6.0 0.41 0.12

0.88 12.0 1.26 0.33

1.04 6.0 1.40 0.43

1.30 6.0 1.68 0.43

11.4 1.71 1.37

8.6 1.26 0.98

9.7 1.99 1.37

6.0 1.62 1.25

8.0 1.78 1.45

5.1 1.52 1.30

6.8 1.96 1.17

3.4 1.54 1.21

8.6 1.24 1.08

12.7 2.75 1.06

11.3 1.85 1.02

8.9 1.43 1.08

43

153

54

20

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 2 (Continued) Sample

Low Nb amphibolite 88

89

High-Sr tonalitic gneisses

90

91

92

93

94

95

96

97

98

99

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

47.46 0.82 19.54 10.05 0.17 5.55 10.12 3.04 1.87 0.44

47.67 1.21 18.05 10.28 0.20 6.69 10.96 3.01 1.05 0.16

51.08 0.44 17.16 7.16 0.18 7.59 12.21 2.82 0.74 0.08

53.84 0.41 11.04 9.10 0.25 12.28 9.42 1.56 1.21 0.14

54.90 0.64 16.18 8.15 0.17 6.10 8.93 2.67 1.52 0.26

55.38 0.54 15.48 7.24 0.20 7.76 8.45 3.08 1.43 0.11

57.04 0.35 13.61 8.18 0.25 7.25 8.95 2.68 0.87 0.05

61.11 0.64 17.26 5.71 0.14 2.94 6.81 3.51 1.21 0.20

63.30 0.46 18.44 3.54 0.07 2.15 5.79 4.19 1.56 0.16

66.34 0.36 17.26 3.52 0.06 1.66 5.73 3.67 1.12 0.11

66.59 0.37 17.42 3.12 0.07 1.53 5.56 4.36 0.83 0.08

67.16 0.33 17.64 2.90 0.06 1.50 5.07 4.30 1.12 0.10

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

99.06 0.50 0.65

99.28 0.54 0.67

99.46 0.65 0.71

99.25 0.71 0.77

99.52 0.57 0.65

99.67 0.66 0.60

99.23 0.61 0.65

99.53 0.48 0.52

99.66 0.52 0.43

99.83 0.46 0.46

99.93 0.47 0.41

100.18 0.48

Trace elements (ppm) Be Sc V 274 Cr 22 Ni 26 Zn 86 Ga 21 Rb 45.0 Sr 1258 Y 26.0 Zr 162 Nb 12.00

363 103 74 111 19 13.0 677 16.0 92 7.10

215 245 48 76 17 3.5 843 16.0 58 6.80

166 523 169 111 16 28.0 112 33.0 80 8.00

201 161 54 75 16 32.0 978 24.0 67 11.00

163 349 115 77 16 36.0 478 22.0 57 9.20

177 205 53 105 15 8.1 368 15.0 107 6.30

129 35 18 68 20 27.0 421 13.0 127 4.40

77 13 10 59 20 44.0 610 4.5 74 5.20

67 9 10 49 18 28.0 487 8.2 67 3.20

52 3 6 48 18 29.0 394 5.5 35 2.10

51 2 3 51 19 28.0 487 3.7 49 2.70

Mo Sn Cs Ba

297

258

476

545

177

438

719

666

539

568

1788

1107

La Ce Pr Nd Sm Eu

135.45 276.1 33.77 131.9 18.45 3.71

34.65 71.6 8.92 34.3 5.67 1.58

30.75 66.0 9.83 40.1 7.10 1.82

21.27 59.4 8.55 35.9 7.01 1.46

42.66 111.2 16.01 69.2 12.43 2.62

34.50 86.6 12.64 54.3 9.70 1.95

14.38 33.8 5.33 22.3 4.19 1.16

17.12 37.4 4.63 19.2 4.15 0.82

26.60 27.5 4.56 16.5 2.06 0.70

8.20 15.6 2.01 8.5 2.11 0.64

7.75 15.7 1.86 7.4 1.68 0.61

19.62 35.0 3.68 12.2 1.75 0.50

Gd Dy Er Yb Lu

11.71 6.16 2.56 2.00 0.30

4.57 3.37 1.57 1.39 0.22

5.58 3.51 1.66 1.67 0.27

6.18 5.47 3.42 3.89 0.62

8.91 5.12 2.27 1.93 0.30

6.93 4.58 2.20 1.97 0.30

3.38 2.88 1.64 1.71 0.29

3.51 2.62 1.33 1.16 0.15

1.48 0.95 0.45 0.48 0.08

1.86 1.64 0.90 0.86 0.13

1.56 1.35 0.60 0.52 0.08

1.23 0.78 0.35 0.29 0.04

Hf Pb Th U

3.50 18.0 15.35 2.40

2.09 13.0 4.52 1.68

1.86 12.0 3.68 1.96

4.17 6.0 2.23 0.98

3.06 14.0 5.41 1.84

2.23 12.0 3.39 1.13

2.48 12.0 1.47 0.82

2.98 15.0 5.51 0.95

2.09 15.0 2.85 0.42

1.11 13.0 1.51 0.43

0.92 7.0 1.69 0.36

0.93 11.0 6.00 0.22

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

48.4 35.71 4.72

42.3 13.32 2.65

52.7 10.23 2.70

3.4 3.95 1.28

40.8 14.90 3.73

21.7 11.38 2.84

24.5 5.11 1.60

32.4 8.35 2.44

135.6 14.80 2.49

59.4 4.69 1.75

71.6 7.78 2.42

131.6 31.22 3.42

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

21

Table 2 (Continued) Sample

High-Sr tonalitic gneisses 100

101

102

103

104

105

106

107

108

109

110

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

68.76 0.27 16.76 2.73 0.06 1.28 4.87 3.72 1.22 0.08

69.48 0.27 16.54 2.66 0.07 1.19 4.51 3.38 1.28 0.08

69.65 0.29 16.60 2.56 0.07 0.75 4.29 4.86 0.72 0.10

70.28 0.23 16.54 2.11 0.03 1.09 4.69 4.32 0.86 0.08

71.10 0.25 16.13 2.29 0.05 0.70 4.01 4.40 0.96 0.08

71.86 0.22 15.72 2.08 0.04 0.65 3.73 4.65 0.96 0.07

72.49 0.20 15.49 1.64 0.05 0.64 3.54 4.60 1.07 0.07

72.68 0.19 15.37 1.66 0.04 0.58 3.85 4.15 0.95 0.06

72.68 0.20 15.47 1.61 0.06 0.60 3.63 4.43 1.26 0.07

73.82 0.15 15.08 1.49 0.03 0.36 4.22 4.18 0.45 0.03

74.42 0.15 14.57 1.66 0.03 0.55 3.68 4.00 1.05 0.02

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

99.75 0.46 0.42

99.46 0.44 0.42

99.89 0.34 0.33

100.23 0.48 0.37

99.97 0.35 0.33

99.98 0.36 0.31

99.79 0.41 0.30

99.53 0.38 0.34

100.01 0.40 0.31

99.81 0.30 0.36

100.13 0.37 0.34

Trace elements (ppm) Be Sc V 41 Cr 5 Ni 5 Zn 42 Ga 17 Rb 32.0 Sr 489 Y 1.8 Zr 53 Nb 3.60

45 4 3 47 18 35.0 445 2.3 47 4.10

25 0 1 51 17 6.8 336 4.1 75 2.10

36 1 4 36 17 23.0 429 2.9 42 2.10

29 0 3 45 16 29.0 328 5.3 82 1.80

25 0 0 54 16 33.0 314 5.2 82 2.10

21 1 0 46 17 31.0 412 4.6 59 2.90

20 0 1 29 16 29.0 459 2.5 51 2.40

24 0 3 43 17 34.0 464 3.0 52 2.40

12 0 0 22 15 11.0 326 2.4 74 1.20

9 0 1 29 15 20.0 491 1.0 78 2.80

Mo Sn Cs Ba

681

504

474

448

519

612

659

242

718

623

1045

La Ce Pr Nd Sm Eu

12.78 22.8 2.29 7.9 1.10 0.37

14.88 24.7 2.68 8.7 1.33 0.44

4.52 10.2 1.20 5.5 1.29 0.39

4.39 11.6 0.97 3.5 0.73 0.37

6.97 14.6 1.77 6.9 1.46 0.53

10.67 21.5 2.47 9.3 1.76 0.51

12.06 21.7 2.27 8.1 1.24 0.38

9.39 18.4 1.71 6.3 0.84 0.16

11.33 20.8 2.31 8.3 1.47 0.50

3.03 5.9 0.56 2.4 0.41 0.26

23.40 48.8 4.01 12.2 1.56 0.57

Gd Dy Er Yb Lu

0.60 0.35 0.15 0.14 0.01

0.88 0.51 0.25 0.26 0.04

1.22 0.89 0.46 0.39 0.04

0.56 0.55 0.26 0.30 0.06

1.34 1.14 0.58 0.56 0.10

1.36 1.02 0.51 0.50 0.08

0.90 0.74 0.35 0.34 0.05

0.51 0.19 0.12 0.11

1.07 0.72 0.34 0.32 0.05

0.38 0.32 0.20 0.26

0.85 0.27 0.10 0.10 0.02

Hf Pb Th U

0.87 12.0 3.38 0.41

1.56 17.0 3.13 0.41

1.70 7.0 0.85 0.20

1.28 10.0 1.96 0.31

2.21 9.0 2.02 0.48

2.36 10.0 3.45 0.65

1.31 14.0 3.68 0.60

0.85 14.0 2.68 0.20

1.10 15.0 2.94 0.22

1.65 7.0 0.48 0.18

2.04 18.0 5.89 0.10

271.7 42.07 3.46

193.5 24.59 2.73

82.0 6.77 2.52

147.9 10.00 1.51

61.9 6.74 1.93

60.4 11.14 2.19

89.6 16.50 2.14

183.6 43.31 3.74

154.7 16.83 2.70

135.8 5.85 1.18

491.0 126.10 6.86

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

22

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 2 (Continued) Sample

Low-Sr tonalitic gneisses 111

112

113

114

115

116

Granodiorite

Alaskite

117

118

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

63.09 0.59 15.63 7.29 0.18 2.05 6.06 3.99 0.55 0.09

69.56 0.42 14.59 4.87 0.14 1.07 4.42 3.38 1.01 0.09

71.68 0.34 14.30 3.83 0.09 1.15 3.06 4.09 1.31 0.07

74.63 0.21 13.75 2.49 0.07 0.56 3.31 4.08 0.87 0.04

74.63 0.23 13.24 2.92 0.08 0.93 3.00 3.82 0.99 0.04

76.62 0.17 12.93 1.66 0.05 0.71 2.43 4.40 0.76 0.03

69.08 0.24 15.02 2.92 0.08 1.23 3.49 3.06 3.74 0.10

75.36 0.04 13.21 1.19 0.07 0.03 0.93 3.06 5.75 0.00

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

99.52 0.33 0.46

99.55 0.28 0.42

99.92 0.35 0.29

100.01 0.29 0.31

99.88 0.36 0.30

99.76 0.43 0.23

98.96 0.43 0.39

99.64 0.04 0.14

Trace elements (ppm) Be Sc V 79 Cr 0 Ni 5 Zn 86 Ga 17 Rb 9.9 Sr 144 Y 25.0 Zr 80 Nb 1.90

31 0 1 59 14 35.0 134 20.0 37 2.40

20 0 2 42 12 46.0 151 16.0 81 2.60

20 0 0 31 12 27.0 106 14.0 76 2.70

11 12 0 0 0 3 33 17 12 10 20.0 20.0 200 155 20.0 20.0 52 101 2.20 3.30

51 8 8 23 14 102.0 578 10.0 129 10.50

4 0 2 20 17 87.0 147 33.0 134 1.40

Mo Sn Cs Ba

89

390

413

336

120

La Ce Pr Nd Sm Eu

5.87 13.5 1.67 8.3 2.20 0.54

9.07 17.9 2.17 9.5 2.33 0.58

9.00 18.0 1.92 8.0 1.61 0.34

6.61 13.7 1.63 6.8 1.63 0.57

11.10 17.4 2.51 10.5 2.32 0.54

11.30 22.1 2.39 9.9 2.02 0.26

63.08 94.7 9.49 30.0 4.14 0.99

4.33 13.1 1.13 4.9 1.67 0.73

Gd Dy Er Yb Lu

3.15 3.69 2.49 2.57 0.37

2.79 3.60 2.38 2.50 0.39

1.76 2.00 1.55 1.91 0.30

1.77 2.38 1.66 1.86 0.31

2.57 3.15 2.36 2.75 0.44

2.44 2.86 2.21 2.64 0.40

3.18 2.00 1.18 1.26 0.22

2.27 4.90 4.55 5.93 0.93

Hf Pb Th U

1.84 5.4 1.58 0.39

1.27 8.2 2.25 0.33

2.20 7.0 3.01 0.30

3.54 9.1 3.62 0.83

2.02 9.0 2.34 0.42

2.69 5.0 4.09 1.36

4.50 35.0 29.28 3.44

4.63 31.0 6.16 0.65

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

5.8 1.35 0.99

6.7 1.85 0.90

9.4 2.44 0.74

7.6 1.90 0.77

10.0 1.63 0.75

7.8 2.16 0.75

57.8 19.43 2.04

4.5 0.57 0.31

140

1850

1910

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

23

Table 3 Chemical analyses of felsic gneisses in domes other than the Monson and Pelham domes Sample

119

120

121

122

SiO2 TiO2 Al2 O3 FeOt MnO MgO CaO Na2 O K2 O P2 O5

72.37 0.24 14.24 1.81 0.06 0.58 1.88 3.64 4.50 0.07

74.93 0.09 13.56 1.22 0.03 0.29 1.61 3.13 4.72 0.03

49.59 0.61 15.87 14.51 0.19 5.38 10.38 1.45 0.53 0.03

72.60 0.13 14.17 2.86 0.06 1.25 6.83 1.76 0.26 0.07

Total Mg=ðMg þ Fet Þ Ca/(Ca + Na)

99.39 0.36 0.22

99.61 0.30 0.22

98.54 0.40 0.80

99.99 0.44 0.68

Trace elements (ppm) V Cr Ni Zn Ga

28 0 2.7 19 13

11 0 7 10 12

679 0.7 8 95 18

53 0.4 2 23 11

134 254 13.0 148 5.6

158 267 12.3 74 10.2

4.9 97 7.5 24 1.6

8 136 6.3 104 0.6

Rb Sr Y Zr Nb Cs Ba

1901

7.5 1285

71

71

La Ce Pr Nd Sm Eu

54.8 91.7 9.15 29.2 4.35 0.93

28.4 44.7 4.35 14.2 2.19 0.42

7.1 14.9 1.71 7.0 1.32 0.45

10.5 19.6 2.16 8.0 1.36 0.35

Gd Dy Er Yb Lu

3.07 2.59 1.62 1.82 0.32

1.85 1.88 1.32 1.64 0.30

1.32 1.34 0.83 0.92 0.14

1.26 1.23 0.81 0.93 0.15

Hf Pb Th U

4.32 17 23.7 4.4

2.96 69 31.2 5.6

0.92

3.60

1.28 0.26

1.64 0.41

Sr/Y CeðnÞ =YbðnÞ GdðnÞ =YbðnÞ

20 13.0 1.4

22 7.0 0.9

Monson and Fourmile mafic rocks plotted with respect to two variables that are sensitive to different chemical alteration processes. For comparison, fresh rocks from younger island arcs and back-arc basins are also plotted. All Monson and Fourmile mafic rocks lie well within the field defined by plagioclase, augite, and OPX þ olivine, and all but one lie within the field of fresh rocks. This one amphibolite (sample 91), the hornblendites, and the hornblende pyroxenites lie on

13 4.2 1.2

22 5.4 1.1

the same general trend as the fresh rocks toward higher mafic mineral content, or toward lower mafic mineral content (only the gabbroic anorthosite). The Fourmile and Monson mafic rocks have apparently not undergone substantial pre- or synmetamorphic alteration as defined by Fig. 6. This is similar to the conclusion for many amphibolites in the Ammonoosuc Volcanics (Schumacher, 1988) and in the Partridge Formation (Hollocher, 1993).

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 4. AFM diagram showing compositions of dome gneisses in the Bronson Hill anticlinorium and volcanics in the Partridge Formation. Data are from Tables 1–3, and from Billings and Wilson (1965), Foland and Loiselle (1981), Hodgkins (1985), Hollocher (1993), Leo (1985, 1991), Leo et al. (1984), Pogorzelski (1983), and Webster and Wintsch (1987). The tholeiitic–calc-alkaline discriminant line is from Irvine and Baragar (1971).

Fig. 5. K2 O–SiO2 chemical variation diagram showing data from the Monson and Fourmile gneisses (Tables 1 and 2). The dome gneisses from elsewhere in the BHA are from Table 3 and from Billings and Wilson (1965); Foland and Loiselle (1981), Hodgkins (1985), Leo (1985, 1991), Leo et al. (1984), Pogorzelski (1983), and Webster and Wintsch (1987). Igneous series discriminant lines are adapted from the Basaltic Volcanism Study Project (1981, Fig. 1.2.7.1).

Based on the dominant occurrence of Monson and Fourmile amphibolites in calc-alkaline fields, and on the lack of a clear iron-enrichment trend (Figs. 4 and 5), we conclude that the Monson and Fourmile amphibolites have calc-alkaline affinities,. The chemical variation diagrams in Figs. 7–10 show, however, that

Fig. 6. This diagram distinguishes fresh mafic igneous rocks from those that have undergone strong chemical alteration by calcite precipitation, low-temperature hydrothermal alteration by sea water or subaerial weathering, and high temperature hydrothermal alteration by sea water (after Schumacher, 1988). Other arc basalts are 361 Tertiary, Quaternary, and Holocene island arc and back-arc basin basalts from Arculus (1976), Kay (1977), Tarney et al. (1977), Basaltic Volcanism Study Project (1981), Dixon and Stern (1983), McKee et al. (1983), Conrad and Kay (1984), Hawkins and Melchior (1985), Reagan and Meijer (1984), Hickey et al. (1986), Nye and Reid (1986), Reagan and Gill (1989), Bacon (1990), Hughes (1990), Leeman et al. (1990), Baker et al. (1991), and Clynne (1993).

the Monson and Fourmile amphibolites do not generally have clear compositional trends indicative, for example, of crystal fractionation. This probably indicates that these rocks did not originate from a homogeneous source and probably had somewhat differing igneous histories. The large amount of compositional scatter of Monson and Fourmile amphibolites contrasts with ‘‘unaltered’’ amphibolites in the overlying Ammonoosuc Volcanics and Partridge Formation, which occur almost entirely in the tholeiitic field on the AFM diagram (Fig. 4) and have moderately well-defined iron-enrichment and differentiation trends (Schumacher, 1988; Hollocher, 1993). Amphibolites occur in two varieties that are distinguished on the basis of Nb and LREE concentrations, and REE patterns (Figs. 11 and 12; summarized in Table 4). The low-Nb amphibolites (<6 ppm Nb) are the most common in the Monson Gneiss, and include all Fourmile amphibolites. The low-Nb amphibolites (Fig. 12(b)) have straight, slightly LREE-enriched REE patterns, small or no Eu anomalies, La concentrations of 7–40 times chondrite, and Lu concentrations of 4–20 times chondrite. Amphibolites from the Ammonoosuc Volcanics and Partridge Formation are similar to the low-Nb amphibolites (Fig. 12(c)). The high-Nb amphibolites (6–12 ppm Nb) have curved, strongly LREEenriched REE patterns (Fig. 12(a)), have small negative to no Eu anomalies, and have REE concentrations of 46–437 times chondrite for La and 7–20 times chon-

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 7. Silica variation diagrams for selected major elements in the Fourmile Gneiss. Data are from Table 1.

drite for Lu. REE patterns of the high-Nb amphibolites are unusual, being steep for the middle REE and flatter toward both La and Lu. Both the low- and highNb amphibolite varieties include quartz- and olivinenormative samples. Unique are high-Nb samples 88 and 89 which have up to 4.6% normative nepheline. Combined with their LREE-enrichment, this suggests that the high-Nb amphibolites may be transitional between calk-alkaline and alkaline basalts (Fig. 5). This contrasts with the conclusion for amphibolite sample 12A from the Partridge Formation (Hollocher, 1993) which has high K2 O and is nepheline normative, but is otherwise identical to the other tholeiitic Partridge amphibolites. It was interpreted as having undergone pre-metamorphic low-temperature alteration in sea water, and was not considered to have had an alkalic basalt protolith. Fig. 13 shows spider diagrams for the Monson mafic rocks. Fig. 13(a) shows that the high-Nb amphibolites are enriched, relative to the low-Nb amphibolites, in all elements more incompatible than Er. The high-Nb amphibolites have large negative Nb anomalies, characteristic of arc basalts, as have the low-Nb amphibolites. Fig. 13(c) shows that the low-Nb amphibolites are, for

25

Fig. 8. Silica variation diagrams for selected major elements in the Monson Gneiss. Data are from Table 2.

the elements plotted, similar to Partridge Formation amphibolites. To summarize, the low-Nb Monson and Fourmile amphibolites have major and trace element characteristics that are typical of calc-alkaline arc basalts (comparisons with arc basalts of Marcelot et al., 1983; Baker, 1984; Briqueu et al., 1984; Thompson et al., 1984). The high-Nb amphibolites appear to be calcalkaline to weakly alkaline, and to have high concentrations of strongly incompatible elements compared to the low-Nb amphibolites. All Nb-rich amphibolites were collected from a 200 m stretch of shoreline on the east side of the Quabbin Reservoir (Fig. 2), suggesting that they probably have a limited aerial distribution and may be from a single intrusion that was deformed and dismembered during metamorphism. It is interesting that all low-Nb Monson amphibolites are finegrained (
1 mm), whereas the high-Nb amphibolites are coarse-grained (3–5 mm). The perfect partitioning of low- and high-Nb amphibolites into fine- and coarse-grained rocks, respectively, suggests that the present grain size was inherited. If so, the low-Nb amphibolites probably represent relatively thin dikes and other small, quickly cooled bodies, and the high-

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 9. Silica variation diagrams for selected trace elements in the Fourmile Gneiss. Data are from Table 1.

Nb amphibolites represent a more slowly cooled larger intrusive body or bodies. Compared to the amphibolites, the Monson gabbroic anorthosite is notable for being rich in CaO and Al2 O3 (plagioclase components) and low in most incompatible elements including TiO2 , P2 O5 , K2 O, and REE. The gabbroic anorthosite REE and spider patterns (Figs. 12(a) and 13(b)) resemble those of the low-Nb amphibolites, not those of the high-Nb amphibolites or the felsic gneisses. The five ultramafic rocks include two hornblendites (samples 64 and 65) and three hornblende pyroxenites (samples 66–68). All are low in Nb and have straight, almost flat REE patterns and no or very small positive Eu anomalies. The high-Mg hornblendite (sample 65) is similar to the hornblende pyroxenites in major element composition, and on AFM, REE, and spider diagrams (Figs. 4, 12(a), and 13(b)). In contrast, amphibolite sample 64 is much more Fe-rich and Mgpoor, has lower SiO2 , REE, Cr, and Ni, but higher V, Rb, Sr, and other strongly incompatible elements. This sample clearly has an origin different from the other ultramafic rocks, and will be discussed in more detail below.

Fig. 10. Silica variation diagrams for selected trace elements in the Monson Gneiss. Data are from Table 2.

5.3. Felsic Gneisses The felsic gneisses span a composition range of
61% to
78% SiO2 , and are generally low in FeOt , TiO2 , MgO, and CaO compared to the amphibolites, and relatively rich in Na2 O and K2 O. Figs. 7–10 show that the composition gap between amphibolites and felsic gneisses, seen in Fig. 4, also exists on silica variation diagrams in the range 57–61% SiO2 . With increasing SiO2 in the gneisses, the concentrations of elements that are compatible in felsic systems all decrease markedly (TiO2 , Al2 O3 , FeOt , MgO, CaO, V, Cr, Ni). However, the concentrations of likely incompatible elements (K2 O, Zr, La, Th) show no consistent pattern with increasing silica, and, particularly for the tonalitic rocks, are enriched only by small factors with respect to the low-Nb amphibolites. The muscovite-rich and microcline-rich Fourmile gneisses have relatively high-normative corundum, mostly 1–2% whereas all other felsic gneisses typically have 0–1%. As stated above, this is interpreted to be caused by loss of Ca and Na during pre-metamorphic hydrothermal alteration (Bull, 1997). The extent of chemical alteration can be evaluated using Fig. 14.

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

27

Fig. 11. La vs. Nb diagram showing how the high-Nb Monson amphibolites are distinguished from all other analyzed mafic rocks from the BHA, which are lower in Nb and LREE. Data are from Tables 1–3 and from Hollocher (1993).

Typical weathering and hydrothermal alteration processes tend to remove Ca and Na relative to Al, moving fresh felsic rock compositions toward the origin (see Schumacher, 1988, for a discussion). Although all Monson and Fourmile felsic gneisses in Fig. 14 lie within the field defined by fresh felsic island arc igneous rocks, the fields of Fourmile muscovite-rich and microclinerich gneisses extend more toward corundum-normative compositions than all other gneisses. This suggests that these samples were indeed altered, though not by enough to cause them to leave the nominal igneous composition field. Fig. 15 shows that the Monson Gneisses and other rocks in the BHA can be divided into two groups on the basis of Sr and Y concentrations (see also Table 4). The high-Sr gneisses have Sr concentrations generally 300– 600 ppm and Y concentrations generally 1–13 ppm, whereas the low-Sr gneisses generally have 50–200 ppm Sr and 10–50 ppm Y. These two groups are clearly differentiated in the Monson Gneiss and in gneisses from the Killingworth Dome (Webster and Wintsch, 1987). The vast majority of dome gneisses outside of the Monson and Killingworth domes, and all analyzed felsic gneisses in the Ammonoosuc Volcanics and Partridge Formation, are of the low-Sr variety. Fig. 8 shows that low-Sr Monson Gneisses also tend to have lower Al2 O3 and higher FeOt concentrations than high-Sr Monson Gneisses of similar silica content. Figs. 16(a), (b) and 17(a), (b) show REE and spider diagrams for the Monson and Fourmile gneisses. Both the low-Sr and high-Sr gneisses are LREE-enriched,

Fig. 12. REE diagrams for mafic rocks in the Monson and Fourmile gneisses (normalizing factors from Boynton, 1984). (a) REE patterns for the Monson high-Nb amphibolites, gabbroic anorthosite, hornblendite, and hornblende pyroxenite. The field of low-Nb Monson and Fourmile amphibolites is shown for comparison. (b) REE patterns of representative low-Nb Monson and Fourmile amphibolites, with the field of high-Nb amphibolites shown for reference. (c) REE patterns of representative amphibolites from the Ammonoosuc Volcanics (Leo, 1985, 1991) and Partridge Formation (Hollocher, 1993). These are similar in trace element composition to the low-Nb Monson and Fourmile amphibolites. The field of Monson and Fourmile amphibolites is shown for comparison.

spanning a La concentration in range of 10–200 times chondrite. Both also have prominent negative Nb anomalies characteristic of magmas in arc environments, and low-Ti concentrations characteristic of felsic igneous rocks in general. The low-Sr gneisses have flat to bowl-shaped MREE to HREE patterns, whereas in contrast the high-Sr gneisses are depleted in HREE compared to MREE. The low-Sr gneisses also have negative Eu and Sr anomalies, whereas the high-Sr gneisses have no significant Eu anomalies and have positive Sr anomalies. Figs. 16(c) and 17(c) show that the felsic gneisses in the Partridge Formation (and also in the Ammonoosuc Volcanics, not shown) have patterns similar to the low-Sr Monson and Fourmile gneisses, although with somewhat higher concentrations of the less incompatible elements.

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Table 4 Summary of amphibolite and gneiss chemical groups Low Nb amphibolites

High Nb amphibolites

Mafic rock groups, illustrated in Figs. 11, 12 and 13 Nb concentration < 6 ppm LREE pattern Flat to weakly enriched MREE–HREE pattern Flat Eu anomalies None to small positive or negative Ti anomalies None Hf and Zr anomalies Small negative Ba anomalies None Nb anomalies Negative Low-Sr felsic gneisses

High-Sr felsic gneisses

Felsic rock groups, illustrated in Figs. 15, 16 and 17 Sr concentration < 200 ppm typical Y concentration 5–60 ppm typical LREE pattern Enriched MREE–HREE pattern Flat to bowl-shaped Eu anomalies Strong negative to small positive Ti anomalies Low Ti, negative anomaly Sr anomalies Ba anomalies Nb anomalies

Negative None Negative

It was mentioned above that the Monson gneisses are mostly hornblende–biotite, biotite, and biotite– garnet tonalites. Both the high-Sr and low-Sr gneisses cover this petrologic range, and both also span similar ranges of SiO2 and of most other major elements. Some effort was spent trying to differentiate between the highSr and low-Sr gneisses in the field and in thin section, but no clear differences were seen. The five Fourmile Gneiss mineralogical and location groupings are caused by small differences in major element composition that, in part, resulted from pre-metamorphic hydrothermal alteration for the muscovite- and microcline-rich gneisses. These five groups of Fourmile Gneiss are not reliably distinguished based on trace element concentrations.

6. Origin of the monson and fourmile rocks 6.1. Low-Nb amphibolites The low-Nb amphibolites are basaltic and have relatively wide ranges of major and trace element concentrations. All, however, have negative Nb anomalies and many have negative Hf and Zr anomalies. These are characteristic of arc magmas and probably originate from a mantle source that is continuously depleted of incompatible elements by melt removal. Replenishment of this source with Nb, Hf, and Zr from the underlying, dehydrating, subducting slab is inefficient compared to REE and other incompatible elements (e.g., Pearce and Peate, 1995).

> 6 ppm Strongly enriched Negative slope None to negative Negative Large negative Small positive Negative

> 300 ppm typical 1–15 ppm typical Enriched Depleted None Low Ti, negative anomaly masked by low HREE concentrations Positive Positive Negative

Fig. 18(a) shows that Ni and Cr covary in the amphibolites, strongly indicating that these compatible elements are controlled by crystal fractionation. Similar relationships are seen among these and other compatible elements, including MgO, FeOt , and V. If all element concentrations are controlled by crystal fractionation, then incompatible element concentrations should increase smoothly as the compatible elements decrease. Fig. 18(b) shows that this idea is too simple. Although rocks having high Zr tend to have low Ni, the trend is not well defined and some amphibolites with low Ni also have low Zr. This pattern is also seen for other incompatible elements. This is strong evidence that the low-Nb magmas initially had a wide range of incompatible element concentrations, suggesting that these magmas were derived either by different degrees of partial melting, or from a mantle source that was inhomogeneous, or both. If the incompatible element variations were caused by differences in the degree of melting in the source region, then ratios of highly incompatible elements should be nearly constant. However, in these rocks most incompatible element ratios vary by factors of 3–6, indicating that the compositional variability was inherited from a source region that was inhomogeneous in time and/or space. It is unlikely that crustal contamination was responsible for the wide chemical variation in these rocks. The only known host rocks for the amphibolites are the calc-alkaline gneisses and, as can be inferred from Figs. 7–10, large changes in incompatible element concentrations would require large changes in silica content as well. The Monson and Fourmile low-Nb amphibolites and amphibolites of the Partridge Volcanics have similar

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 13. Spider diagrams for mafic rocks in the Monson and Fourmile gneisses. Elements are arranged from left to right in order of increasing compatibility in basalt systems (Hoffman, 1988; normalizing factors are those used by Hollocher, 1993). The mobile elements Rb, Ba, K, and Sr are shown in light type. (a) Comparison of the fields of Monson high-Nb amphibolites and Monson and Fourmile low-Nb amphibolites. (b) Monson gabbroic anorthosite and ultramafic rocks, with the field of low-Nb amphibolites shown for reference. (c) Comparison of the field of low-Nb Monson and Fourmile amphibolites with the field for ‘unaltered’ amphibolites from the Partridge Formation (Hollocher, 1993), which are similar in trace element composition to the low-Nb Monson and Fourmile amphibolites.

major and trace element compositions (Figs. 4, 11, 12(c), 13(c)). The Partridge amphibolites were interpreted as having been derived by partial melting of a spinel lherzolite source rock (Hollocher, 1993). Based on similarities of chemical composition, it is likely that the Monson and Fourmile low-Nb amphibolites were also derived from a similar, but more strongly inhomogeneous, spinel lherzolite source. 6.2. High-Nb amphibolites The high-Nb amphibolites have compositions similar to the low-Nb amphibolites in several respects: both seem to be largely calc-alkaline (Figs. 4 and 5), both have similar ranges of compatible element concentrations (Figs. 7–10, and 18), and both have negative Nb,

29

Fig. 14. This diagram distinguishes fresh felsic igneous rocks from those having undergone hydrothermal alteration or weathering (after Schumacher, 1988). Most alteration processes tend to remove Ca and Na from felsic rocks, moving rock compositions toward the origin. Severely altered rocks should lie outside the envelope of fresh igneous rocks. The field of fresh igneous rocks is defined by a set of 115 mostly calc-alkaline arc and back-arc basin felsic volcanic and plutonic rocks having SiO2 contents of 61–78%, from Bacon (1990), Arculus (1976), Bice (1985), Box and Patton (1989), Bryan et al. (1972), Clynne (1993), Dixon and Stern (1983), Hildreth (1981), Kay (1977), Leeman et al. (1990), Lonsdale and Hawkins (1985), Lowder and Carmichael (1970), Mahood (1981), Merzbacher and Eggler (1984), Noble et al. (1984), Putman and Alfors (1969), Reagan and Meijer (1984), Robinson et al. (1984), Sajona et al. (1993), Sims et al. (1985), Smith and Johnson (1981), Swanson and McDowell (1985), and Vennum and Rowley (1986).

Hf, and Zr anomalies Fig. 13). The high-Nb amphibolites, however, have substantially higher concentrations of Nb, Zr, Hf, LREE, and other elements more incompatible than Er. The negative Nb, Zr, and Hf anomalies are therefore relative to greatly enriched LREE. The overall similarity between the high- and low-Nb amphibolites in major element composition (Fig. 8) and overall shape (not slope or absolute value) of the spider diagrams (Fig. 13) suggests that the highNb amphibolites were derived from melting of a spinel lherzolite mantle source rock similar to that inferred for the low-Nb amphibolites. The high-Nb amphibolite source must, however, have been more enriched in strongly incompatible elements. Enrichment in Na2 O and K2 O were sufficient to make high-Nb samples 88 and 89 nepheline normative. Though we refer to these rocks as high-Nb amphibolites, they are not like the high-Nb basalts reported from modern subduction environments in association with adakites, that are thought to have been derived

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 15. Sr vs. Y diagram showing that the dome gneisses occur in two dominant groups: a high-Sr, low-Y group and a larger low-Sr, generally high-Y group. Data for the Fourmile and Monson Gneisses are from Tables 1 and 2. Data for the volcanics and other BHA dome gneisses are from Table 3 and Hollocher (1993), Leo (1985, 1991), Schumacher (1988), Webster and Wintsch (1987).

from direct melting of the subducting slab (e.g., Defant et al., 1992). Such basalts generally have >15 ppm Nb, >1.5% TiO2 , and >0.5% P2 O5 and lack negative Nb anomalies, unlike any of the high-Nb amphibolites we report. We interpret the high-Nb magmas as having been derived from a mantle wedge source that was more enriched in highly incompatible components from the subducting slab than was the source for the low-Nb magmas. Enrichment of the mantle wedge is generally believed to be caused by transfer of material from the subducted slab and oceanic sediments into the melting region of the overlying mantle wedge. Transfer of the chemical components upward is generally thought to be mediated by migrating fluids produced during dehydration of amphibole and other hydrous minerals in the subducting slab (Othman et al., 1989; Plank and Langmuir, 1993; Pearce and Parkinson, 1993; Pearce et al., 1995; Ernst, 1999). High-Nb amphibolite/low-Nb amphibolite enrichment factors are 5–8 for highly incompatible elements including LREE,
3.5 for Nb,
1.5 for Zr and Hf, and
1 for the HREE. The greater enrichment of Nb, Zr, and Hf compared to the HREE suggests that the high-Nb mantle source had been enriched both by H2 Orich fluids, in which Nb, Zr, and Hf probably have low solubility, and by silicate liquids, which are more effective at transporting Nb, Zr, and Hf. The lack of enrichment in HREE may imply garnet in the slab that supplied the enriching fluids and melts.

Fig. 16. REE diagrams for the felsic Monson and Fourmile gneisses (normalizing factors from Boynton, 1984). (a) Comparison of the fields of low-Sr and high-Sr groups of felsic gneisses. (b) REE patterns for representative low-Sr and high-Sr Monson and Fourmile gneisses. (c) Comparison of the field of low-Sr Monson and Fourmile gneisses with that of felsic volcanics in the Partridge Formation (Hollocher, 1993).

6.3. Ultramafic rocks and gabbroic anorthosite The three Monson hornblende pyroxenites and hornblendite sample 65 are chemically similar to one another and probably have similar origins. These rocks have low TiO2 , Al2 O3 , and incompatible elements, high Ni, Cr, and MgO, and relatively high SiO2 ,
52–53%, indicative of pyroxene-rich cumulates. These samples are interpreted to be cumulates from magmas similar in composition to the low-Nb amphibolites. This interpretation is supported by similarity in ultramafic rock mineralogy to that of the amphibolites, similarity in shape of REE and spider diagram patterns, and physical association of the ultramafic rocks with low-Nb amphibolites in the outcrop. Hornblendite sample 64 came from a large boudin in deformed tonalitic gneiss. It differs from the other four Monson ultramafic rocks in having much lower SiO2 , MgO, Cr, and Ni, and much higher TiO2 , FeOt , and V. Note that, in Fig. 4, a line passing through the trend of felsic rock data would pass approximately through this

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

31

rich plagioclase (labradorite), absent in the felsic rocks. 6.4. Low-Sr felsic gneisses

Fig. 17. Spider diagrams for the felsic Monson and Fourmile gneisses (for plotting details see Fig. 13). (a) Comparison of the fields for low-Sr and high-Sr groups of felsic gneisses. (b) Representative low-Sr and high-Sr Monson and Fourmile gneisses, the same samples as shown in Fig. 16(b). (c) Comparison of the field of low-Sr Monson and Fourmile gneisses with that of felsic volcanics in the Partridge Formation (Hollocher, 1993).

amphibolite sample. This suggests that sample 64 is a cumulate from the felsic magmas, and that crystal fractionation of hornblende controlled part of their compositional range. The Monson gabbroic anorthosite is interpreted to be a plagioclase-rich cumulate from magmas similar to the low-Nb amphibolites. This interpretation is supported by low-Nb concentration (0.7 ppm) and negative Nb anomaly, similarity of overall REE concentrations (Fig. 12(a)), small LREE enrichment compared to MREE and HREE, positive Eu and Sr anomalies, generally low concentrations of elements incompatible in plagioclase (e.g., TiO2 , K2 O, P2 O5 , Zr), and similarity of spider diagram patterns (Fig. 13(b)). These chemical characteristics can be easily modeled by simple plagioclase-rich crystal accumulation from a crystallizing low-Nb amphibolite magma. Derivation from the felsic rocks is unlikely because: (1) simple crystal fractionation models yield poor matches for the gabbroic anorthosite; and (2) the gabbroic anorthosites have abundant augite and Ca-

The composition gap and overwhelming volumes of felsic rock in the Monson and Fourmile gneisses clearly indicate that the mafic and felsic rocks are not cogenetic. Indeed, it has long been understood that most felsic magmas are derived from partial melting of deep crustal rocks (e.g., Roberts and Clemens, 1993), rather than by any direct process from ultramafic mantle rock or from basaltic magma. When the Monson and Fourmile felsic magmas were emplaced, the deep BHA crust was probably dominated by rocks of broadly basaltic composition: gabbros and related cumulates, amphibolites, and pyroxene granulites. These hypothetical source rocks were probably formed by the crystallization of mantle-derived basaltic magmas in the deep crust, and by depression of the lower crust as the volcanic pile on the surface accumulated. These processes permit a deep arc crust compositionally similar to mafic rocks that were emplaced at shallower levels (the amphibolites sampled in this study). With SiO2 increasing in the gneisses from
61% to
78% the content of compatible elements (FeOt , CaO, MgO, TiO2 , Cr, V, Ni) decreases substantially (Figs. 7– 10), indicating some control by crystal fractionation. However, over this same SiO2 range the incompatible elements (e.g., La, Zr, Th, K2 O) do not increase regularly. This suggests that source composition was the dominant control on magma composition, rather than crystal fractionation. Experiments have shown that melting basaltic rocks, leaving a pyroxene- or amphibole-rich restite, can generate metaluminous to weakly peraluminous felsic magmas much like the felsic Monson and Fourmile gneisses (see discussion and references in Gromet and Silver, 1987; Schumacher, 1988; Hollocher, 1993). Hollocher (1993) modeled felsic volcanics in the Partridge Formation (and by analogy the felsic Ammonoosuc Volcanics) as having been derived from melting of deep crustal basaltic rocks with compositions similar to the mafic Partridge volcanics. The felsic volcanics were modeled successfully by assuming batch melting of a pyroxene granulite source leaving an augite–plagioclase–olivine–orthopyroxene–ilmenite–magnetite–apatite restite assemblage. These models indicated that little or no garnet or amphibole were present in the restite. A similar approach was taken for modeling the origin of the Monson and Fourmile gneisses. The uncertainties of source composition and mineralogy, partition coefficients, and which, if any, felsic magma compositions represent original liquids, are

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Fig. 18. (a) Ni vs. Cr diagram for mafic rocks. The strong positive correlation between these two compatible elements in rocks with <1000 ppm Cr indicates control by crystal fractionation. (b) Ni vs. Zr diagram for mafic rocks. The large amount of scatter of Zr concentrations at various Ni concentration indicates inhomogeneous magma source rocks. Data for Monson and Fourmile rocks are from Tables 1 and 2; data for the Ammonoosuc and Partridge amphibolites are from Schumacher (1988) and Hollocher (1993).

large. Therefore, a simple-minded batch melting model was used initially and modified as necessary. The base model was constructed as described by Hollocher (1993) for the Partridge felsic volcanics, except that the source rock composition was taken to be the average of all lowNb Monson amphibolites, and the model target composition was the average low-Sr Monson Gneiss. The source rock restite mineralogy was first taken to be similar to the norm of the average low-Nb Monson amphibolite, differing in that normative orthoclase (an artifact of the norm calculation, not likely to have been present in the source as a mineral phase) was replaced by additional normative CPX. The model source restite therefore had the mineralogy: plagioclase 55.0%, CPX 20.9%, OPX 16.1%, Olivine 3.5%, ilmenite 1.9%, magnetite 2.3%, apatite 0.3%. Calculated liquids at 5%, 20%, and 40% partial melting are shown in Figs. 19(a) and 20(a). The most important points to note are that the model MREE and HREE are too high compared to the average low-Sr felsic gneiss, and the model REE patterns do not have the concave-up ‘‘saddle’’ shape common in the low-Sr dome gneisses. A second model was therefore constructed in which amphibole was made from normative olivine, plagioclase, and CPX in the proportions 35:25:40. In this model, olivine is exhausted first at
10% model amphibole. The resulting mineralogy for the second model restite was: plagioclase 52.5%, CPX 16.9%, OPX 16.1%, amphibole 10.2%, ilmenite 1.9%, magnetite 2.3%, apatite 0.3%. The results are shown in Figs.

19(b) and 20(b). The addition of amphibole results in model liquids reproducing the characteristic saddle shape of the low-Sr felsic gneiss REE patterns, and results in a good fit for most elements including the prominent negative Nb, Sr, Eu, and Ti anomalies. The Nb anomaly is inherited from the model source rock, the Sr and Eu anomalies are the result of plagioclase in the source restite, and the Ti anomaly is caused by TiO2 retained by mafic minerals in the source. The poor model fits for Th and U are a concern, but uncertainties in the partition coefficient values for these elements suggest that this is not presently a major problem. In any case, the model fits are not improved by inclusion of garnet, titanite, or biotite, or by including restite entrainment in the magmas, assimilation of a shale (schist) component, or by using different melt extraction models (e.g., fractional melting, dynamic melting). By adjusting the model restite mineralogy, particularly the amount of amphibole, plagioclase, and augite, mimicking source major element inhomogeneity, it was possible to closely reproduce the range of low-Sr Monson and Fourmile felsic gneiss compositions. These models do not support large amounts (>20%) of crystal fractionation. 6.5. High-Sr felsic gneisses A third model was constructed to evaluate the origin of the high-Sr Monson gneisses, using the same methods and source rock composition as described

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

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Fig. 19. REE diagrams showing the results of melting models for the felsic gneisses. In each case the lower crustal source rock had the composition of the average Monson and Fourmile low-Nb amphibolite, and involved different percentages of batch melting. Shaded regions are the fields of the low-Sr gneisses (a and b) and high-Sr gneisses (c). (a) Melting model for the low-Sr gneisses involving a plagioclasepyroxene restite. This was successful for modeling the felsic Partridge formation volcanics (Hollocher, 1993). (b) Same as (a), but involving a plagioclase-pyroxene amphibolite restite assemblage. (c) Melting model for the high-Sr gneisses, in which melting leaves a pyroxene– amphibole–garnet–plagioclase restite assemblage.

Fig. 20. Spider diagrams showing results of melting models for the felsic gneisses. In each case the lower crustal source rock had the composition of the average Monson and Fourmile low-Nb amphibolite, and involved different percentages of batch melting. Shaded regions are the fields of the low-Sr gneisses (a and b) and high-Sr gneisses (c). (a) Melting model for the low-Sr gneisses involving a plagioclasepyroxene restite. This was successful for modeling the felsic Partridge formation volcanics (Hollocher, 1993). (b) Same as (a), but involving a plagioclase-pyroxene amphibolite restite assemblage. (c) Melting model for the high-Sr gneisses, in which the restite has a pyroxene– amphibole–garnet–plagioclase assemblage.

above, but with the target composition being the average of all high-Sr Monson and Fourmile gneisses. A model using a HREE-depleted basaltic source rock was considered, but no such rocks have been found in Ordovician amphibolites from the BHA, and such rocks are rare in subduction zone environments (as summarized by Gromet and Silver, 1987). It was obvious from the low-Sr gneiss models that garnet was required in the source restite to accurately model the HREE-depletion of the high-Sr gneisses. The higher pressure implied by the presence of garnet also implies partial or complete loss of plagioclase from the model assemblage, and also implies the possible occurrence of sodic pyroxene or amphibole. Because of the larger number of degrees of freedom caused by these changes, we ran a large number of models that included a variety of plausible source restite modes (plagioclase 0–

55%, CPX 10–60%, OPX 0–16%, Olivine 0–10%, amphibole 0–50%, garnet 0–50%, ilmenite 0–2%, magnetite 0–2%, apatite 0–0.3%). For each starting mode, mineral proportions were adjusted to find a local leastsquares best fit to the average high-Sr gneiss target composition, based on the 20 elements in the spider diagrams (e.g., Fig. 20). Twenty percent batch melting of the source rock was used for all of these models. Although a rather wide range of initial source rock mineralogy yielded reasonably successful high-Sr model liquids, garnet was always 9.5–11.5% of the model restite modes. This result is caused in part by the fact that, of the 20 modeled elements, five are MREE and HREE which garnet strongly controls. The results for the most successful model are shown in Figs. 19(c) and 20(c) (source rock restite: plagioclase 3.0%, CPX 50%, OPX 15.8%, amphibole 20%, garnet 9.5%, ilmenite

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1.4%, apatite 0.3%). The REE are modeled quite well except for Eu, for which the model liquids have small positive anomalies rather than no anomaly. Sr, Nb, and Ti are modeled well. As with the low-Sr gneiss models, the fit for the most incompatible elements, particularly U and Th, was rather poor. Additional models were also tested to derive the compositions of several individual high-Sr gneiss samples. It was possible to successfully model different samples with restite mineralogy ranging from a plagioclase–pyroxene-5% garnet amphibolite, to a pyroxene50% garnet eclogite. However, some gneiss samples (e.g., samples 102, 103, 109) having particularly lowLREE concentrations require a LREE-depleted model source rock. Additionally, LREE-rich samples (e.g., 96, 99, 110) require either a somewhat LREE-enriched source or substantial crystal fractionation. Incompatible element concentrations are poorly correlated with compatible element concentrations in the high-Sr gneisses. This argues against compositional control dominated by crystal fractionation, and argues for source region inhomogeneity with respect to mineralogy and trace element composition. The low-Sr and high-Sr Monson gneisses appear to lie on two distinct trends on silica variation diagrams: the low-Sr gneisses have lower Al2 O3 and higher FeOt contents than the high-Sr gneisses at given SiO2 concentrations (Fig. 8). This difference is consistent with the source rock mineralogy predicted by trace element modeling. The low-Sr gneisses are derived from a plagioclase–pyroxene–amphibole source, so the resulting liquids should have low Al2 O3 which is retained by restite plagioclase, and high FeOt which is not effectively retained by Mg-rich pyroxenes and amphiboles compared to garnet. In contrast, the high-Sr gneisses are derived from a pyroxene–amphibole–garnet source, so these liquids should have higher Al2 O3 , which is not as effectively retained by garnet and other mafic minerals compared to plagioclase, and lower FeOt which is effectively retained by garnet. The Monson and Fourmile felsic gneiss data lie along relatively narrow, straight lines (Figs. 7–10); two lines (Fig. 8) in the case of Al2 O3 and FeOt . The straight, narrow trends suggest mechanical mixing lines between two end members: two magma types, magma and restite, or two magma source compositions. This idea is easily tested since all elements, including major, trace, compatible, and incompatible elements, should vary linearly between the two hypothetical end members. Examination of Tables 1 and 2, and mixing calculations (not shown), demonstrate that many trace and some major elements do not vary linearly between the hypothetical end members with increasing SiO2 , Al2 O3 , FeOt , MgO, TiO2 , or other measures. This supports the idea that the source rock for these magmas was inhomogeneous. The narrow, linear trends for

most major elements are presumably controlled by melting phase relations in the source, and probably to a lesser extent by polyminerallic fractional crystallization. Both the low-Sr and high-Sr felsic gneisses are therefore likely to have been partial melts of compositionally and mineralogically somewhat inhomogeneous mafic rocks. The low-Sr gneisses were derived by melting of a plagioclase–pyroxene–amphibole, garnet-free source rock, presumably at relatively low pressures (<10 kbar, after Wyllie and Wolf, 1993), as has also been proposed for compositionally similar rocks in subduction environments elsewhere (e.g., Drummond and Defant, 1990; Tepper et al., 1993). In contrast, the high-Sr gneisses were probably derived by melting of pyroxene–garnet  amphibole  plagioclase source rock, presumably at higher pressures (
10–22 kbar, Wyllie and Wolf, 1993). The proposed typical source mineralogy, having
10% garnet, is transitional between pyroxene granulites and amphibolites at lower pressure and eclogite at higher pressure. This suggests that the high-Sr source rocks could have been at intermediate pressures of
10 kb. Rocks compositionally similar to the high-Sr gneisses are uncommon in modern subduction environments, but they have been described from several locations (Kay, 1978; Gromet and Silver, 1987; Drummond and Defant, 1990; Defant et al., 1992; Sajona et al., 1993; Defant and Drummond, 1993; also see Defant and Kepezhinskas, 2001). Volcanics of this type have been named adakites (Defant et al., 1992; Sajona et al., 1993). In modern environments these rocks are generally similar to normal calc-alkaline andesites and dacites (or plutonic equivalents), but are higher in Sr, lower in Y, are strongly depleted in HREE relative to LREE, and have small or no Eu anomalies. These rocks are generally interpreted as having been derived by partial melting of garnetbearing subducted hot, young oceanic crust (garnet amphibolite or eclogite). Our modeling suggests that the high-Sr gneisses are derived from rocks having typically
10% garnet by weight (range 5–50%). This suggests that the source region for the high-Sr gneisses was largely transitional between amphibolite and typical eclogite. In most environments where both HREE-depleted magmas and the more normal magmas with flat MREE and HREE patterns occur, the HREE-depleted magmas are emplaced closer to the trench. In the example from the Peninsular Ranges batholith, however, Gromet and Silver (1987) concluded that the HREE-depleted rocks were emplaced farther from the trench. The abundance of high-Sr, HREE-depleted gneisses on the east side of the BHA in the Monson dome cannot, therefore, be used alone to distinguish between east- or west-directed subduction beneath the Taconian arc.

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7. Timing of arc collision Work by Tucker and McKerrow (1995) has improved radiometric time constraints on the Ordovician biostratigraphic time scales in both North America and Europe. For the purpose of this paper, their most important conclusions are: Arenig Series base is
485 Ma, Llanvirn Series base is
470 Ma, Llandeilo Series base is
464 Ma, Caradoc Series base is
458 Ma, Ashgill Series base is
449 Ma, and the Ashgill Series top, equivalent to the base of the Silurian, is
443 Ma. Although Tucker and McKerrow (1995) do not place uncertainties on these ages, we estimate that most are probably accurate to 2 or 3 Ma based on an examination of their radiometric constraints. We will use this radiometric time scale for the following discussion. Fig. 21 shows the timing of metamorphism, structural development, sedimentation, and igneous activity related to the Taconic Orogeny in the New England Appalachians. Taconian orogenesis was clearly a protracted episode lasting tens of millions of years. For much of this time the Taconian arc was distant from Laurentia, but during collision it involved extensive metamorphism and contraction of Laurentian basement and sedimentary cover rocks (e.g., Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985; Sutter et al., 1985; Drake et al., 1989; Tremblay, 1992; Cawood et al., 1995).

35

The ages of Taconian volcanics and dome gneisses in the Shelburne Falls belt, west of the Connecticut River (Figs. 1 and 22), are
479–452 Ma (Arenig to middle Caradoc; Karabinos and Tucker, 1992; Karabinos et al., 1996, 1998). These are mostly older than the 454–442 Ma ages of dome gneisses and volcanics in the central BHA (middle Caradoc to lowest Silurian; Tucker and Robinson, 1990) and dome gneisses in the northern BHA which span
456–441 Ma (lower Caradoc to lowest Silurian; Moench et al., 1995). In the northern BHA in New Hampshire and Maine, Moench (1993), Moench et al. (1995) and Drake et al. (1989) argue that the Ammonoosuc Volcanics and Partridge Formation are middle Ordovician. For the Ammonoosuc Volcanics this age is supported by a radiometric age of 467  3 Ma for one of the Chickwolnepy intrusions that cuts sheeted dikes that are interpreted to have fed the Ammonoosuc Volcanics (Fig. 21; Moench, 1993; Moench et al., 1995; Fitz, 1996), an age of 469  1:3 Ma for the Joslin Turn pluton that cuts Ammonoosuc Volcanics west of the Jefferson dome, an age of 465  6 Ma for a felsic volcanic correlated with the Ammonoosuc Volcanics (Moench and Aleinikoff, 2001), and an age of 461  8 Ma from a felsic layer in the Ammonoosuc Volcanics (Moench, 1993; Moench et al., 1995). These constraints yield a likely age of
469 Ma (Llanvirn) for the Ammonoosuc Volcanics in the northern BHA.

Fig. 21. Summary diagram for timing of the Taconic Orogeny in western New England with respect to metamorphism, constraints on thrust slice emplacement, and igneous rock crystallization ages. Age of geologic period and stage boundaries are those of Tucker and McKerrow (1995) and Tucker et al. (1998).  458 Ma Pb/Pb evaporation age now superceded by a 471 Ma U/Pb age on the same sample; oldest reliable age in Shelburne Falls dome is 479 Ma (oral communication from Paul Karabinos, 2001).  Excludes East Inlet pluton (430 Ma), thought by Robert Moench (oral communication, 2001) to be unrelated to the other Highlandcroft plutons.

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Fig. 22. Schematic tectonic and timing model for the BHA arc igneous activity with respect to collision. The accretionary wedge complex is shown to be emergent to limit sediment transfer from the arc into the trench region, as indicated by Nd and Pb isotopic evidence for a largely Laurentian source for forearc sediments (Andersen and Samson, 1995; Bock et al., 1998). This geometry still permits erosion of ultramafic rocks in the accretionary wedge to supply Cr- and Ni-rich sediment to the trench region (Garver et al., 1996). See text and the figure for references to age information.

The Partridge Formation in the northern BHA, overlying to interfingering with the Ammonoosuc Volcanics, is constrained by a 444  4 Ma date determined on a felsic tuff in the lower Quimby Formation (unit Oqv of Moench, 1996, his Fig. 2 and Stop 1). However, note that Oqv of Moench (1996) is identified as unit Oaux, bimodal volcanics in the Ammonoosuc Volcanics overlying a layer of Partridge Formation, by Lyons et al., 1997). If the Lyons et al. (1997) interpretation of the dated unit is correct, then the Ammonoosuc Volcanics in the northern BHA spans an astounding
25 Ma (
469–444  4 Ma). In the northern BHA the middle part of the Partridge Formation is constrained by graptolites assigned to the North American Climacograptus bicornis graptolite zone (Harwood and Berry, 1967; unit called the Kamankeag Formation by Osberg et al., 1985; but correlated with Partridge Formation by Moench and Boudette, 1987; and Moench, 1993, 1996). Tucker and McKerrow (1995) conclude that the best radiometric ages for the Millbrig and Deicke K-bentonites in Virginia are 453:1  1:3 and 454:5  0:5 Ma, respectively. These bentonites occur in the upper part of the European Diplograptus multidens graptolite zone, constraining the upper part of this zone to middle Caradoc at
454 Ma. The upper D. multidens zone correlates with the uppermost part of the North American C. bicornis graptolite zone (Finney, 1986), placing the Partridge Formation graptolites in the middle Caradoc (upper Ordovician as defined by Palmer, 1983; Haq and van Eysinga, 1998) at
454 Ma. The age of the lower Partridge Formation in the northern BHA is therefore between the
469 Ma age of the Ammonoosuc Volcanics and the
454 Ma graptolite age of the middle Partridge, thus Llanvirn to middle Caradoc.

The Llanvirn to middle Caradoc age for the Ammonoosuc Volcanics and Partridge Formation in the northern BHA is mostly older than the dates for these units in the central BHA in Massachusetts: 453  2 Ma (middle Caradoc) for the upper member of the Ammonoosuc Volcanics and 449þ3=2 Ma (latest Caradoc or earliest Ashgill) for a felsic volcanic near the base of the Partridge Formation (Tucker and Robinson, 1990). A simple explanation is that the northern BHA Ammonoosuc Volcanics and Partridge Formation are not correlative with younger units of the same name in the central BHA. Because the lower mafic member of the Ammonoosuc Volcanics in the central BHA (Schumacher, 1988) has never been precisely dated, it is possible that it alone is older and correlative with the northern BHA Ammonoosuc Volcanics. The relatively young ages of BHA dome gneisses and central BHA Ammonoosuc Volcanics and Partridge Formation, contrasting with older ages of the Chickwolnepy intrusions, Ammonoosuc Volcanics and Partridge Formation in the northern BHA, and older ages of dome gneisses and volcanics in the Shelburne Falls belt west of the BHA (Fig. 1) have led others to propose that the Taconian arc is a composite terrane with the BHA belt representing one of its youngest components (e.g., Karabinos and Tucker, 1992; Kim and Jacobi, 1996; Kusky et al., 1997; Karabinos et al., 1996, 1998). Kusky et al. (1997), for example, present a model for the Taconian orogeny that involves a complex amalgamation of arcs between Laurentia and Avalon. These are proposed to have formed and accreted to Laurentia or Avalon at times spanning the entire Ordovician. Though this particular model is principally based on the geology of southeastern Quebec, northwestern New Brunswick, and western Maine, their model includes collision of a Bronson Hill–Notre

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Dame arc with Laurentia in the middle Ordovician (Llanvirn;
470–464 Ma). Taconian metamorphic ages have been used to place constraints on Taconian orogenesis. These metamorphic ages range from Lower Ordovician through Silurian (Fig. 21; Laird et al., 1984; Drake et al., 1989; Hames et al., 1991; Sevigny and Hanson, 1995; Whitehead et al., 1996). Karabinos et al. (1998, 1999) use some metamorphic ages in western New England, and the older radiometric ages of Shelburne Falls belt volcanics and dome gneisses, to support a model in which the Shelburne Falls belt was a volcanic arc separate from and earlier than the BHA. They proposed that the Shelburne Falls belt arc collided with Laurentia in the middle Ordovician, causing much of the Taconian deformation and metamorphism in western New England, eastern New York State, and southeastern Quebec, including emplacement of the Taconic allochthons. They suggest that the BHA developed later above a west-dipping subduction zone east of the accreted Shelburne Falls belt, though we know of no firm structural or stratigraphic evidence for a west-dipping Ordovician subduction zone beneath the BHA. The hypothetical middle Ordovician collision of the Shelburne Falls arc is roughly contemporaneous with emplacement of ophiolites and the Ascot Complex onto Laurentian crust in southeastern Quebec (Tremblay, 1992; Pinet and Tremblay, 1995), and the time of emplacement of Taconic allochthons in Newfoundland (e.g., Cawood et al., 1995). The problem of the timing of collision of the Taconian arc with Laurentia therefore involves disparate evidence involving wide ranges of radiometric ages of Taconian magmas, ages of metamorphism, and biostratigraphic constraints. Here we concentrate on evidence that we believe gives the best constraints on the timing of arc collision with Laurentia in the New England region. As discussed by Ratcliffe et al. (1999a,b) and Robinson et al. (1998), some of the older Taconian metamorphic dates in New England contradict stratigraphic and fossil constraints. The Utica Shale of eastern New York overlies platform carbonate rocks and represents deposition on subsiding Laurentian crust during advance of the Taconian arc (Fig. 22). The Utica Shale, in the central and eastern Mohawk Valley, is constrained by fossils to the Corynoides americanus through the Climacograptus spiniferus graptolite zones (Fisher, 1977; Goldman, 1995). The uppermost European D. multidens zone is correlated with the lowermost North American C. americanus graptolite zone (Finney, 1986), placing the C. americanus zone close to
454 Ma in the middle Caradoc (as discussed above regarding the Partridge Formation). The C. americanus zone is overlain by the Orthograptus ruedemanni and C. spiniferus zones (Fisher, 1977; Finney, 1986), the latter being latest

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Caradoc. The Utica Shale therefore extends from middle to latest Caradoc (
454–449 Ma). The Schenectady Formation in part overlies the Utica Shale and is composed of interbedded dark gray shales and greywacke, representing deposits more proximal to the advancing Taconic allochthons than the Utica Shale. The Schenectady Formation is constrained by fossils to the C. spiniferous graptolite zone (Fisher, 1977), latest Caradoc and as young as
449 Ma. The Taconic belt east of the flat-lying sedimentary rocks of the Mohawk River valley contains the classic Taconic allochthons. These overrode parautochthonous rocks that are locally exposed in windows eroded through the thrust sheets. The parautochthonous Ira and Walloomsac Formations are shales and slates originally deposited on the Laurentian continental margin (Stanley and Ratcliffe, 1985). These two units are tightly constrained by fossils to the upper part of the C. bicornis graptolite zone (middle Caradoc, Finney, 1986; Ratcliffe et al., 1999b), as are the youngest rocks in the Giddings Brook slice. The Giddings Brook slice also overrides the Utica Shale in the eastern Mohawk Valley (Potter, 1972; Finney, 1986). Ratcliffe et al. (1999b) report conodonts from the West Bridgewater Formation on the east side of the Green Mountain massif in Vermont. This unit is interpreted to be a member of the cover sequence deposited in deep water on the Laurentian continental margin. The Periodon aculeatus and Protopanderodus conodonts constrain deposition of this unit to between late Arenig and lower Caradoc,
475 to
454 Ma. As discussed by Ratcliffe et al. (1999a,b), biostratigraphic constraints demand that overriding of the Laurentian continental margin by the Taconian accretionary wedge cannot have been earlier than
475 to
454 Ma for the West Bridgewater Formation in deep water east of the present Green Mountain massif. Similarly, overriding of the Ira and Walloomsac Formations or the easternmost Utica Shale on the subsided Laurentian margin cannot have occurred earlier than
454 Ma in the middle Caradoc. The Schenectady Formation is cut by a Taconian thrust, exposed along the north shore of the Mohawk River west of Lock 7 (Fisher et al., 1971; Kidd et al., 1995, Stops 1 and 2). This thrust separates flat-lying Schenectady Formation flysch to the west from similar but deformed rocks to the east (Normanskill Formation in Fisher et al., 1971; reclassified to Austin Glen Formation suggested by Kidd et al., 1995). This thrust must post-date deposition of the latest Caradoc Schenectady Formation, so faulting was latest Caradoc or younger. Subduction and advancement of the Taconian arc onto Laurentia therefore continued into latest Caradoc and possibly beyond the Caradoc–Ashgill Stage boundary at
449 Ma. This interpretation is consistent with data and models of other studies, including: C. Spiniferous zone

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graptolites in the matrix of m
elange underlying the Taconic allochthons (see discussion, p. 677 of Bradley and Kusky, 1986); Taconian flexural faulting and trench sedimentation extending into the uppermost Caradoc and lower Ashgill (Bradley and Kusky, 1986); and the orogen-scale analysis of Taconian arc convergence of Bradley (1989). Drake et al. (1989) summarize the occurrence of middle to late Ashgillian (latest Ordovician) and Silurian molasse deposits in Maine, New York, Pennsylvania, and New Jersey (e.g., the Lower Silurian Shawangunk Formation) that were derived from the Taconian orogen. By earliest Silurian the collision had ceased and the Taconian marginal basin on Laurentian crust had largely filled, allowing deposition of terriginous sediments westward. The structural and biostratigraphic constraints therefore place the end of Taconian convergence between latest Caradoc at
449 Ma and earliest Silurian at
443 Ma, encompasing the Ashgill Stage. Within radiometric and biostratigraphic age uncertainties, the BHA dome gneiss magmas (456–441 Ma, Fig. 21) therefore can have been emplaced (though barely) during normal subduction in the upper Ordovician. The ages of some intrusives, including some of the Highlandcroft plutons in northwestern New Hampshire (452–430 Ma, Lyons et al., 1986; Aleinikoff and Moench, 1987; Moench et al., 1995; Figs. 1 and 21), the Mount Hermon Gabbro of northern Massachusetts (Fig. 2; Silurian, Elbert, 1984), some intrusives in southwestern Connecticut (453–428 Ma, Sevigny and Hanson, 1993, 1995), and the Cortlandt Complex (
430 Ma, Domenick and Basu, 1982;
440 Ma, Bender et al., 1984) extend into middle Silurian and so post-date the apparent end of Taconian convergence. We suggest that the younger magmatism was related to partial delamination or full detachment and sinking of the subducted Iapetus Ocean lithospheric slab (Fig. 22). Mantle and crustal magmas would have been produced during: (A) sinking of the oceanic crust through the zone of amphibole dehydration at
100 km depth (e.g., Ernst, 1999), and/or (B) rising of asthenospheric mantle to fill the vacated space, with concurrent decompression melting and heating of the lower crust (e.g., Carpathians, G^ırbacea and Frisch, 1998; see also Wortel and Spakman, 2000). Vertical sinking of a detached slab or slab rollback during delamination would tend to cause the magmatic axis to shift westward from the original BHA magmatic axis, possibly explaining the location of the Highlandcroft and other late plutons within and west of the BHA. Dvorkin et al. (1993) indicate that narrow subducted slabs, or by implication detached slabs, should sink essentially vertically. Numerical models of Schott and Schmeling (1998) are consistent with vertical sinking of a detached slab or lithosphereic block, and with the slab

rollback conceptual model of G^ırbacea and Frisch (1998). Further, G^ırbacea and Frisch (1998) indicate that the delamination and rollback process can permit igneous activity to continue for at least 5–15 Ma after the end of subduction convergence. If this duration of post-collision magmatism is appropriate for the Taconian case, and if we assume that the end of collision was latest Caradoc (
449 Ma), then magmatism may have continued to
434 Ma, well into the Silurian and at least partially encompassing the ages of the younger plutons (in Connecticut as young as 428  2 Ma, Pumpkin Ground pluton, Sevigny and Hanson, 1993; for the Highlandcroft plutons as young as 430  4 Ma, East Inlet pluton, Lyons et al., 1986). We note also that postcollision magmatism in the Carpathians involves alkalic basalts (G^ırbacea and Frisch, 1998), possibly analogous to the alkalic basalt parent magmas of the Cortlandt Complex.

8. Discussion and summary The Fourmile Gneiss consists of several mineralogically distinct rock types, dominated by tonalitic and granodioritic gneisses with minor amphibolites of basaltic composition. In the southern part of the Pelham dome, the lower member of the Fourmile Gneiss consists of rusty-weathering muscovite-rich and microclinerich gneisses that are thought have undergone premetamorphic hydrothermal alteration resulting in minor loss of Ca and Na. These rocks, however, are not so altered as to lie outside the field of fresh igneous arc rocks (Fig. 14). The lower member of the Fourmile Gneiss contains the majority of amphibolites in this unit. These amphibolites commonly have gedrite and abundant garnet, indicative of pre-metamorphic chemical alteration (as discussed for other amphibolites nearby by Schumacher, 1988, and Hollocher, 1993). The analyzed Fourmile amphibolites did not contain large quantities of these minerals and were apparently not substantially altered (Fig. 6). The upper muscovite-poor member of the Fourmile Gneiss in the southern Pelham dome, is not thought to have undergone significant chemical alteration. In the northern Pelham dome the Fourmile Gneiss is dominated by the biotite-bearing Tailrace gneiss and the hornblende–biotite northern gneiss, which are also thought to be effectively unaltered. All five parts of the Fourmile Gneiss are interpreted to be deformed plutonic rocks, based on preserved massive, homogeneous outcrops that remain despite strong deformation, and on their association with similar gneisses of inferred plutonic origin in other domes. The Monson Gneiss is dominated by gray-weathering tonalitic rocks with a significant amount of amphibolite, and some ultramafic, anorthositic, and granodioritic rocks. None of the rocks sampled appear to have been

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

chemically altered to any significant degree, although some have apparently undergone partial melting (up to perhaps 2% liquid) during Acadian or younger metamorphism. The partial melts were avoided during sampling. In addition, the partial melts are tonalitic and similar in composition to the host gneisses (Hollocher, 1988), so small amounts of melt removal or addition were unlikely to have altered rock compositions significantly. In its less deformed portions the Monson Gneiss clearly resembles a deformed plutonic complex. The felsic Fourmile and Monson Gneisses occur in two chemically distinct groups. Both groups have negative Nb anomalies and other characteristics of calc-alkaline island arc magmas. The low-Sr group has low Sr and Al2 O3 , high Y and FeOt , flat MREE through HREE patterns, and typically have negative Eu anomalies. Modeling suggests that these were derived by melting of lower crustal mafic rocks leaving a plagioclase–pyroxene–amphibole restite. Compared to the low-Sr group, the high-Sr rocks have high Sr and Al2 O3 , low Y and FeOt , are strongly depleted in HREE, and typically have no Eu anomalies. Modeling suggests that these were derived by melting of deeper mafic rocks, leaving a pyroxene–garnet  amphibole  plagioclase restite. The typical model garnet proportion in the restite is
10%. We suggest that the garnet-bearing source rocks may be an underthrust portion of older mafic crust in the Taconian arc complex, perhaps a remnant from an earlier episode of subduction (Fig. 22). The Fourmile and Monson amphibolites also occur in two groups. The low-Nb group is most abundant and is compositionally similar to typical island arc tholeiitic and calc-alkaline basalts, including negative Nb anomalies characteristic of magmas derived from the mantle wedge in subduction zone environments. The high-Nb group is similar in many respects but with higher Nb, strongly enriched LREE and other incompatible elements, unusual REE patterns that are steep for MREE and flatten toward LREE and HREE, and two samples that are nepheline normative. The high-Nb amphibolites may be transitional between calc-alkaline and alkaline arc basalts. We have not found in the literature rocks with REE patterns similar to those of the high-Nb amphibolites. The felsic BHA gneisses show regional compositional differences (Fig. 23). Dome gneisses within and north of the Mascoma and Lebanon domes in west-central New Hampshire are dominantly granitic and are all of the low-Sr type. The southern BHA dome gneisses are mostly of the low-Sr type, but in contrast are dominantly tonalitic to granodioritic. The southern BHA gneisses include significant quantities of the high-Sr type rocks in the Monson and Killingworth domes. Roberts and Clemens (1993) conclude that ‘‘I-type’’ calc-alkaline igneous rocks, in general, inherit their K2 O contents largely from their source rocks. The largely granitic BHA dome gneisses in the northern domes were there-

39

Fig. 23. Map summarizing the aerial distribution of felsic BHA dome gneiss rock compositions. Patterns in the BHA domes are as follows. Black: dominantly granitic rocks in northern domes, probably derived from melting of garnet-free intermediate or felsic deep crustal rock. Horizontal lines: dominantly tonalitic and granodioritic rocks in southern domes, probably derived from melting of garnetfree mafic deep crustal rock. Checkered pattern: same as the other southern domes, but includes substantial quantities of high-Sr, lowY, HREE-depleted felsic rocks that are nearly absent from the rest of the BHA. These rocks were probably derived from melting of garnet-bearing deep crustal mafic rock. The chemical and lithologic data on which this figure is based are from Tables 1–3, and from Billings and Wilson (1965); Foland and Loiselle (1981), Hodgkins (1985), Leo (1985, 1991), Leo et al. (1984), Pogorzelski (1983), and Webster and Wintsch (1987). Inset graphs of the same type have the same scales.

fore derived from K2 O-rich intermediate or felsic crustal source rock of igneous composition, probably the remains of older continental crust on which the arc was built. The tonalitic and granodioritic gneisses in the southern domes were probably derived from a more mafic, probably largely basaltic crustal source involving

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K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

a much smaller proportion of felsic continental material. Such along-arc differences in igneous rock compositions have been described elsewhere (e.g., Kay et al., 1990). Deep crustal sources for the low-Sr felsic magmas were clearly present across the width and length of the BHA volcanic arc, whereas the presumably deeper garnetbearing source for the high-Sr magmas was available only in the southeastern part of the arc (Fig. 23). Samson (1994) and Andersen and Samson (1995) reported eNd values of )5.8 to þ1.0 for several felsic dome gneisses from the central BHA, and interpreted these to indicate involvement of old continental crust in the production of dome gneiss magmas, in addition to a mantle component. Initial 87 Sr/86 Sr ratios are available from early attempts to date the BHA dome gneisses using the Rb–Sr system (Brookins and Hurley, 1965; Naylor, 1969; Wintsch and Grant, 1980; Foland and Loiselle, 1981; Leo et al., 1984). These yielded a set of 87 Sr/86 Sr initial ratios that range from 0.7045 to 0.711 in the Mascoma, Lebanon, Glastonbury, Monson, and Killingworth domes. Though these initial ratios are of mostly of rather low precision, they are elevated above mantle values and so support the eNd evidence for involvement of an older felsic crustal component in the generation of the BHA magmas. 87 Sr/86 Sr initial ratios for the Ammonoosuc Volcanics are 0.705–0.706, in the same range as the dome gneisses (Brookins, 1968). If old continental material was present in the BHA felsic magma source regions, it is odd that no zircon inheritance has been reported from BHA dome gneisses (to our knowledge, excluding the Dry Hill Gneiss and related rocks in the Pelham dome; Brookins and Hurley, 1965; Naylor, 1969; Zartman and Leo, 1985; Leo et al., 1984; Aleinikoff and Moench, 1987; Tucker and Robinson, 1990; Moench et al., 1995). Although some of these workers (e.g., Tucker and Robinson, 1990) carefully selected clean prismatic zircons and zircon tips that were less likely to contain inherited cores, others used a variety of zircon shape and size fractions. Highlandcroft plutons (Fig. 1), in contrast, overlap in age the BHA igneous rocks (Fig. 21) but have abundant inherited zircon components (Lyons et al., 1983; Lyons et al., 1986), as do dome gneisses in the Shelburne Falls belt (Karabinos and Tucker, 1992). It is possible that the old continental component indicated by eNd values and 87 Sr/86 Sr initial ratios was indirectly transferred from low eNd , high 87 Sr/86 Sr subducted sediment into the deep arc crustal source regions for the BHA felsic magmas (a possibility in modern arcs, e.g., Othman et al., 1989), rather than directly by melting of old felsic crustal rock. Andersen and Samson (1995) and Bock et al. (1998) concluded that Taconian forearc sediments have essentially undiluted Laurentian Nd and Pb isotopic compositions. If such sediments were an important part of the material sub-

ducted beneath the Taconian arc, they may have contributed their old continental crust signature to the BHA felsic magmas. Inherited zircons, in contrast to whole-rock isotope ratios, must have been derived directly from the felsic magma source regions. Almost all BHA dome gneisses reported in the literature and in this paper have low Zr contents (median is 103 ppm for all 168 reported Zr analyses). Using the zircon saturation equation of Watson and Harrison (1983), we find that almost all BHA felsic gneisses are likely to have been undersaturated with respect to zircon if magmas ever approached their liquidus temperatures. Such temperatures would range from 850 °C to 1150 °C depending on the sample and on H2 O content (Ghiorso and Sack, 1995). We suggest that low-Zr concentrations in the source region and in derived magmas, and undersaturation of the felsic magmas with respect for zircon, led to complete dissolution of relic zircon. We have found only two Zr analyses for Highlandcroft rocks in the literature (Leo, 1991), but note that one of the two has 1598 ppm Zr (Leo, 1991), several times the saturation concentration at any reasonable magmatic temperature. In this case, zircon inheritance could hardly have been avoided. The Highlandcroft plutons, and perhaps the Shelburne Falls belt dome gneisses, may have been derived from source regions that were relatively Zr-rich. These sources may have included relatively pristine continental material, less diluted by large volumes of mafic arc magmas compared to the BHA belt sources. Several recent models for the Taconic Orogeny involve the collision of two or more arc segments along an irregular Laurentian margin (e.g., Tremblay, 1992; Pinet and Tremblay, 1995; Van der Pluijm et al., 1995; Kim and Jacobi, 1996; Kusky et al., 1997; Karabinos et al., 1998). We interpret current evidence to require, in the New England region, only one arc collision with the Laurentian margin, terminating in the latest Caradoc to earliest Silurian,
449 to
443 Ma. The scale of the Taconian orogen from northern Newfoundland to southern New England (
1700 km) or further south to Taconian rocks and structures in the central Appalachians (
2400 km; Drake et al., 1989) is similar to that of modern island arcs and arc complexes such as the Java–Timor segment of the Sunda arc (analogy made by Tremblay, 1992), the New Britain arc plus the Solomon islands, the Ryukyu Islands of Japan, the Kuril Islands system of Russia and Japan, and the Philippines. These arc complexes are inhomogeneous with respect to basement rock, size and location of sedimentary basins, age and volume of igneous rocks, and accretion history. Though more work is needed to understand the interaction of the Taconian arc complex with the eastern margin of Laurentia, there is clearly room within the orogen for different segments to have experienced different geologic histories.

K. Hollocher et al. / Physics and Chemistry of the Earth 27 (2002) 5–45

Acknowledgements Paul Karabinos and Johnathan Kim reviewed this manuscript and made numerous useful comments for improvement. Charles Mitchell, at the State University of New York at Buffalo, helped us understand graptolite biostratigraphy. The analytical work of Hollocher and Bull and the field work of Bull were supported by NSF grant 9 105 438 (to Robinson). To each of these persons and institutions we express our grateful acknowledgement.

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