Late Silurian bimodal volcanism of southwestern ... - GeoScienceWorld

Late Silurian bimodal volcanism of southwestern New Brunswick, Canada: Products of continental extension

Nancy A. Van Wagoner* Acadia University, Department of Geology, Wolfville, Nova Scotia B0P 1X0, Canada

Matthew I. Leybourne Department of Geosciences, University of Texas at Dallas, Box 830688, Richardson, Texas 75083-0688, USA

Kelsie A. Dadd Department of Earth and Planetary Sciences, Macquarie University, Sydney, New South Wales 2109, Australia

Diane K. Baldwin Wayne McNeil Acadia University, Department of Geology, Wolfville, Nova Scotia B0P 1X0, Canada

ABSTRACT Silurian–Devonian bimodal volcanism— a common feature of the northern Appalachians of the northeastern United States and eastern Canada—forms major belts of igneous rocks including the Coastal volcanic belt of New England and New Brunswick. The Coastal volcanic belt is one of the largest bimodal volcanic provinces in the world. Controversy remains regarding the tectonic setting of this volcanism; suggestions range from volcanic arc, to extensional continental rift, transpressional basin, backarc, or plume-modified subduction zone. The mapped Passamaquoddy Bay volcanic sequence of the Coastal volcanic belt of southwestern New Brunswick has a minimum thickness of 4 km and comprises 63 lithologic units that are divided into four distinct cycles of mafic-felsic volcanism. A preliminary U-Pb zircon date indicates an age of 423 Ma, similar to the age of the Cranberry Island bimodal volcanic series of the Coastal volcanic belt in Maine. Our analysis of 115 volcanic rocks for major and trace element concentrations indicates that the sequence is subalkalic and bimodal (basalt-rhyolite) with within-plate tectonic affinities. Most mafic rocks range in composition from basalt to basaltic andesite. The Mg# [defined as 100 3 MgO/(MgO 1 FeO), where FeO 5 FeOT 3 0.9] of the basalts ranges from 30 to 70, and there is a *E-mail: [email protected].

trend toward more primitive compositions upward in the section. The light rare earth element abundances of the basalts are enriched to levels of 30–130 times chondritic values. The flows are enriched in incompatible elements with respect to primitive mantle and show distinct Nb, Hf, and Sr anomalies. We interpret the basalts to be mantle melts modified by crustal contamination and mantle metasomatism from a previous subduction event. Most of the chemical variation can be explained by replenishment and fractionation of small magma chambers. The rhyolitic rocks are crustal melts, modified by crystal fractionation. Within each of the rhyolitic cycles there is an upward trend toward more evolved compositions as a result of an evolving magma chamber and magmachamber zonation. We suggest that the rocks were erupted in an extensional environment and that there was a significant zone of continental extension during the Late Silurian in the northern Appalachians. However, controversy remains regarding the precise nature of the tectonic setting and the duration and extent of this volcanism. Keywords: bimodal volcanism, Coastal volcanic belt, continental extension, northern Appalachian volcanism, petrogenesis. INTRODUCTION Silurian–Devonian bimodal volcanism is a common feature of the northern Appalachians

GSA Bulletin; April 2002; v. 114; no. 4; p. 400–418; 15 figures; 1 table.

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of the northeastern United States and eastern Canada (Bevier and Whalen, 1990; Boucot, 1968; Boucot et al., 1974; Dostal et al., 1989; Fyffe et al., 1999; Murphy et al., 1995; Seaman et al., 1999). These volcanic rocks form three major volcanic belts: the Coastal volcanic belt (Fig. 1), the Tobique volcanic belt in northern Maine, and the Piscataquis volcanic belt in northern New Brunswick and Quebec. The Coastal volcanic belt is one of the largest bimodal volcanic provinces in the world (Seaman et al., 1999). There is controversy regarding the tectonic setting of the Silurian–Devonian bimodal volcanism of the northern Appalachians. Several workers (Bird and Dewey, 1970; Bradley, 1983; Dewey and Kidd, 1974; Murphy et al., 1999; Wilson, 1966) suggested that the Coastal volcanic belt is an ancient volcanic arc related to the closing of the Iapetus Ocean. Gates and Moench (1981) suggested an extensional tectonic setting on the basis of their documentation of bimodal volcanism in the Machias-Eastport area in Maine. Since then, attempts to define the nature of the extensional environment have resulted in suggestions of a continental rift (Dostal et al., 1989), transpressional basins associated with transcurrent faulting (Murphy et al., 1995), backarc extension (Fyffe et al., 1999), and plume-modified subduction (Murphy et al., 1999). The complex tectonic history of the northern Appalachians can only be resolved though careful geological and geochemical studies. This study focuses on the geochemistry of the Passamaquoddy Bay volcanic sequence of

LATE SILURIAN BIMODAL VOLCANISM OF SOUTHWESTERN NEW BRUNSWICK

Figure 1. (A) Black box shows geographic setting of the study area. (B) Coastal, Piscataquis, and Tobique volcanic belts (after Dostal et al., 1989). Patterned areas show the location of the volcanic belt, and black areas show exposures of volcanic rocks. (C) Location map of the study area (modified after Fyffe and Fricker, 1987). Shown is the Passamaquoddy Bay volcanic sequence (PBVS), mapped in detail in Figure 2. The sequence is intruded in the north by the St. George batholith and overlain unconformably in the south by the Upper Devonian Perry Formation. southwestern New Brunswick, the thickest and most continuous part of the Coastal volcanic belt studied to date. This geochemical study follows detailed lithostratigraphic mapping (Van Wagoner et al., 1994) and is the first lithogeochemical study of the area. These data provide important constraints on tectonic models for the northern Appalachians and associated volcanism.

GEOLOGIC SETTING In southwestern New Brunswick, bimodal volcanic and sedimentary rocks in the Passamaquoddy Bay area (PBVS in Fig. 1) make up the eastern part of the Coastal volcanic belt as defined by Boucot (1968) and Boucot et al. (1974) and cover an area of ;185 km2. Dostal et al. (1989) included the Devonian

Geological Society of America Bulletin, April 2002

rocks in the Cobequid and Antigonish Highlands of Nova Scotia (Murphy, 1988) within the Coastal volcanic belt; however, these sequences are younger than those of the Passamaquoddy Bay volcanic sequence area and are therefore no longer considered part of the Coastal volcanic belt. The Passamaquoddy Bay volcanic sequence is unconformably overlain in the south by subaerial clastic rocks of the Upper Devonian Perry Formation and intruded to the north by the St. George batholith (Figs. 2 and 3). The area covered in this paper includes 63 lithologic units (Fig. 3). The sequence reveals four cycles of mafic-felsic volcanism, in which mafic volcanism accompanied and apparently preceded rhyolitic volcanism (Figs. 2 and 3). Felsic volcanic rocks predominate in all but the final cycle, which is dominated by sedimentary rocks and may represent the waning stages of volcanism in the zone. Although the lowermost mafic-rhyolitic cycle is the first cycle recognized in the map area, it does not represent the initiation of volcanism. A volcaniclastic unit at the base of the sequence comprises both felsic and mafic volcanic clasts derived from older and probably cogenetic volcanic deposits. The base of the volcanic sequence is not exposed owing to intrusion of the St. George batholith. The volcanic sequence has a minimum thickness of ;4 km and includes fine-grained, littoral-facies sedimentary rocks. The sequence lacks major unconformities, indicating that volcanism and sedimentation were continuous. The excellent preservation of glass shards in the volcaniclastic sedimentary rocks indicates that sedimentation rates were rapid. Reworked volcanic rocks are rare, and those derived from mafic sources predominate. There were a multitude of vent areas for the rhyolitic units, but the mafic flows probably originated from a single rift or vent area (Van Wagoner et al., 1994). Brief descriptions of the petrography and mineral assemblages are presented here to provide context for the geochemistry. Detailed descriptions are in Van Wagoner et al. (1994). Mafic Rocks Mafic volcanic rocks occur as flows and rare volcaniclastic units at four stratigraphic levels within the volcanic complex, at the base of each of the four volcanic cycles (Figs. 2 and 3). Only the flows are discussed in this paper. Mafic lava flows from the three lower cycles are similar in mineral assemblage, texture, and structure. The flows at the top of the sequence contrast with these lower flows and

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Figure 2. Map of the study area, showing groupings of lithologic units according to volcanic cycle; cycle 1 is the oldest, and cycle 4 is the youngest. reflect a slightly different style of volcanism. The units range in maximum thickness from ;12 to 460 m and comprise many individual flows. Flows from the lower three cycles thin to the south and change from being relatively massive in the north to comprising pahoehoe toes in the south, indicating a northern source area. The basalt flows typically occur as a series of stacked flows 1–12 m thick with interbedded peperitic breccia (Dadd and Van Wagoner, in press). In contrast, flows of cycle 4 occur as several thick flows, typically ;12 m thick, separated by sedimentary horizons. The flows in cycles 1–3 typically contain rare plagioclase phenocrysts in a pilotaxitic groundmass of plagioclase and clinopyroxene microlites and glass now altered to chlorite, epidote, actinolite, and rare quartz. Plagioclase is typically altered to saussurite. Flows in cycle 4 typically have a subophitic texture and comprise plagioclase, clinopyrox-

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ene, and opaque minerals, secondary chlorite, epidote, actinolite, and rare quartz and calcite. Several flows in cycle 4 (west) contain olivine phenocrysts (6%–12%) and rare clinopyroxene phenocrysts. The olivine phenocrysts are altered to serpentine minerals and iddingsite. Flows in cycle 4 (east) contain rare chlorite and iron oxide pseudomorphs after olivine. All flows are highly to sparsely amygdaloidal; the abundance of amygdales increases upward in most flows and represents up to 85% of the volume of unit D3mf1 (see Fig. 3 for unit explanations). The cycle 4 flows also have a high amygdale concentration at their base where they directly overlie sedimentary rocks. Samples chosen for geochemical analyses were taken from the massive, central part of flows where possible. Felsic Rocks The felsic volcanic rocks occur as flows and domes, as bedded and massive vitric, lithic-


vitric, crystal-vitric, and lapilli tuffs, and as densely welded tuffs. The flows and domes vary from flow-banded to massive and from sparsely porphyritic to aphyric. Peperitic breccia is associated with several of the domes, and its occurrence at the top and base of the domes indicates that they are intrusive (Dadd and Van Wagoner, in press). The felsic pyroclastic rocks can be divided into massive and bedded types. The vitric-rich massive tuffs were probably deposited by pyroclastic-flow processes, whereas the crystal- and lithic-rich bedded tuffs show bedding characteristics typical of both fall and surge deposits (Van Wagoner et al., 1994). The felsic flows and domes of cycles 1–3 are aphyric to sparsely porphyritic with plagioclase and, less commonly, sanidine phenocrysts that are rarely clustered in a devitrified microgranophyric to felty groundmass. Some samples contain ferromagnesian minerals altered to chlorite, opaque minerals, and

Figure 3. Composite section of the stratigraphy of the map area. The four cycles of mafic-felsic volcanism are indicated on the left. Arrows indicate the major component of flow direction in the plane of the section. Circled dots indicate flow direction perpendicular to the page.



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Age There is considerable uncertainty regarding the age of the Passamaquoddy Bay volcanic sequence. A recent U-Pb zircon age from a felsic tuff from the area indicates crystallization at 423 6 1 Ma (Van Wagoner et al., 2001). Several earlier workers (Cumming, 1967; Hay, 1967; Ruitenberg, 1968) correlated this sequence with volcanic formations in Maine ranging in age from Wenlockian to Gedinnian (Silurian to Early Devonian). Pickerill and Pajari (1976) confirmed the correlation to the Early Devonian Eastport Formation of Maine on the basis of similar lithology and faunal assemblages and proposed that the New Brunswick rocks also be termed ‘‘Eastport Formation.’’ However, two isotopic age determinations make the latter correlation suspect. The St. George batholith intrudes rocks of the Passamaquoddy Bay volcanic sequence in the north of the study area. The Utopia Granite, a part of the St. George batholith, yielded a UPb (zircon) age of 430 6 3 Ma (Bevier, 1989) and a 40Ar/39Ar total gas age of 418 6 5 Ma (McLeod, 1988). These isotopic dates, therefore, suggest an even older age for the Passamaquoddy Bay volcanic sequence (.430 Ma). The Cranberry Island series in Maine, a bimodal volcanic sequence that is also part of the Coastal volcanic belt, has a U-Pb age of 424 Ma (Seaman et al., 1995). ANALYTICAL METHODS Figure 4. Geochemical discrimination diagrams, showing the bimodal and subalkalic nature of the Passamaquoddy volcanic sequence. Diagram is from Winchester and Floyd (1977).

calcite. The lowermost felsic unit (D1ff1) is a flow-banded rhyolitic dome with a trachytic texture. It contains up to 70% altered feldspar microlites in a recrystallized, flow-aligned, silicic groundmass. The felsic flows and domes of cycle 4 are similar to those of cycles 1–3, but also contain altered ferromagnesian minerals including biotite and pyroxene. All flows in cycles 1–4 contain rare zircon. Several domes in cycle 4 have contrasting mineral assemblages and textures. Unit D4ff3 is an intrusive dome containing ,5% plagioclase phenocrysts in a pilotaxitic groundmass of devitrified glass with clinopyroxene and plagioclase microlites, opaque minerals, and rare zircon. Secondary chlorite occurs in the groundmass. Unit D4td1 is a dacitic to rhyolitic dome. It is sparsely porphyritic with phenocrysts (2%–5%), commonly clustered, of plagioclase, altered ferromagnesian minerals, rare

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fresh clinopyroxene, and rare sanidine in a finegrained groundmass of feldspar, altered ferromagnesian minerals, minor quartz, and rare apatite and zircon. Unit D4td2, is nonporphyritic to highly porphyritic with phenocrysts of plagioclase, clinopyroxene, hornblende, and opaque minerals in a microcrystalline groundmass of feldspar, clinopyroxene, hornblende, and opaque minerals. Clinopyroxene phenocrysts have rare rims of biotite. The vitric tuffs, and crystal- and lithic-vitric tuffs (.5% crystals and lithic clasts, respectively) are composed of pumice, crystals, and lithic clasts in a matrix of granophyric, recrystallized, and devitrified glass. Crystals, including plagioclase, sanidine, and rare euhedral zircon, are commonly broken and enriched in the matrix relative to pumice. Minor accidental lithic clasts include vesicular basalt, pink rhyolite, and rare mudstone.


Rocks were selected to characterize unit geochemistry and geochemical stratigraphy. We analyzed major and trace element concentrations in 115 of the least altered volcanic rocks: 47 from mafic flow units, 43 from felsic flow units, 18 from felsic pyroclastic units, and 7 from intermediate domes. The felsic pyroclastic rocks chosen for analysis were primarily vitric tuffs composed of essential material, glass, pumice, and crystals and should, therefore, provide an estimate of magmatic composition. Rock chips were handpicked to remove lithic clasts and altered material before pulverizing in a tungsten carbide mill. Analyses were recalculated to 100% volatilefree before being plotted on diagrams. Major and trace element analyses using a Philips PW1400 sequential X-ray fluorescence spectrometer were performed at St. Mary’s University, Nova Scotia. The standards used for major element analysis were HFL-1 and GDV-1A and U.S. Geological Survey sample N-1. The internal standard used at St. Mary’s University for trace element determination was HFL-1. Loss on ignition values (LOI)


of the analyses have high loss on ignition (LOI) values, and those with LOI . 6 were not considered further. The regional metamorphic grade of the study area is lower greenschist to subgreenschist facies, as indicated by the presence of secondary quartz 1 epidote 1 chlorite 1 albite 6 actinolite 6 pumpellyite 6 prehnite in the volcanic and sedimentary rocks (Van Wagoner et al., 1994). Greenschist-facies metamorphic mineral assemblages occur in amygdules, veins, altered matrix material, plagioclase phenocrysts and vitric pyroclasts. A contact-metamorphic aureole occurs adjacent to the St. George batholith; the metamorphic grade reaches the hornblende hornfels facies. Because of the alteration and metamorphism that has affected these rocks, our interpretations of the petrogenesis and tectonic setting are based largely on the least mobile trace elements (Zr, Nb, Y, and Ti) and REEs. The systematic behavior of these elements in the Passamaquoddy Bay volcanic sequence suggests that they were immobile during alteration and metamorphism. GEOCHEMISTRY

Figure 5. Chemical variation diagrams for the basaltic rocks, showing major element oxides plotted against Mg#, defined as 100 3 MgO/(MgO 1 FeO), where FeO 5 FeOT 3 0.9.

were determined by heating the samples to 1050 8C. Rare earth element (REE) and additional trace element (Li, Be, Sc, Mo, Cs, Hf, Ta, Tl, Bi, and U) concentrations were determined in 45 samples by using standard HF 1 HNO3 digestion and analysis by inductively coupled plasma–mass spectrometry (ICP-MS). The analyses were performed at Memorial University of Newfoundland following techniques of Jenner et al. (1990). Of these 45 samples, 23 were from mafic flows, 4 from felsic pyroclastic rocks, 15 from felsic lava flows, and 3 from intermediate domes.

ALTERATION AND METAMORPHISM Rocks of the study area are both slightly altered and metamorphosed, typical of volcanic rocks of this age. Alteration is usually the result of changes due to the interaction of the igneous rocks with groundwater, seawater, or hydrothermal fluids. The alteration processes particularly affect concentrations of SiO2, Fe2O3, Na2O, and K2O. During early devitrification and hydration, glassy volcanic rocks may undergo chemical mobility. This can lead to loss of Na and K from the glassy groundmass of rhyolites (Cerling et al., 1985). Many


The volcanic rocks of the Passamaquoddy Bay area form a bimodal (mafic and felsic), subalkalic suite with a SiO2 gap between ;58 and 64 wt%. Mafic rocks range in composition from basalt to basaltic andesite. Several samples are andesitic. Felsic volcanic rocks range in composition from rhyodacite to rhyolite; some appear as pantelleritic as a result of elevated Zr/TiO2 ratios (Fig. 4). The Nb/Y ratio is considered a good indicator of chemical affinity (Pearce and Cann, 1973; Winchester and Floyd, 1977), being less affected by alteration and metamorphism than the alkalic elements on the basis of which volcanic rocks are typically classified. The Nb/Y ratio signifies that both the mafic and felsic rocks are subalkalic. The Nb/Y ratio, however, is sensitive to crustal contamination of mafic melts and subduction-related mantle metasomatism; contamination leads to lower Nb/Y ratios (e.g., Wilson, 1989). The subalkalic signature of the mafic volcanic rocks is also supported by the Zr/TiO2 ratios. On the basis of Zr/TiO2, most of the felsic rocks plot in the fields of rhyolite and rhyodacite-dacite. Some of the felsic rocks, however, are classified as comendite-pantellerite owing to their high Zr concentrations (Fig. 4).

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ation is the preferred mechanism for Cr depletion as its abundance is not expected to vary under conditions of increasing partial melting (Pearce and Norry, 1979). FeO*, TiO2, and V increase with decreasing Mg# to Mg# ø 45, below which their abundances drop sharply, indicating the onset of iron-titanium oxide fractionation. High- and low-Ti groups can be distinguished, with a break at ;2% TiO2. Unlike many continental provinces where two Ti groups are distinguished (e.g., Fodor and Vetter, 1985; Fodor et al., 1990), this grouping does not appear to indicate two separate and contrasting melts, but more likely represents fractionation processes. The CaO/Al2O3 ratio is relatively constant, with decreasing Mg# to Mg# ø 40, below which the ratio decreases. The constant CaO/Al2O3 ratio probably indicates contemporaneous clinopyroxene, plagioclase, and olivine fractionation. Below Mg# ø 40, olivine is no longer a fractionating phase, consistent with the observed low Ni values, and the drop in the CaO/Al2O3 ratio may indicate a decrease in the ratio of clinopyroxene to plagioclase in the melt. Plagioclase crystallization is also indicated by the decrease in Sr with decreasing Mg#. Positive correlations of P2O5 and Y with Mg# suggest that apatite was not a fractionating phase. Cycle 4 (east) and cycle 4 (west) are best distinguished on the basis of P2O5, Y, and Zr variations with respect to Mg# (Figs. 5 and 6). Mafic rocks in cycle 4 (west) have higher P2O5, Y, and Zr concentrations at a given Mg# compared to cycle 4 (east). One flow from cycle 4 (west) for which Mg# ø 47 plots consistently with many analyses from cycle 4 (east). This sample is from a small outcrop and may be from an intrusion within the outcrop area of D4mf2. Figure 6. Chemical variation diagrams for the basaltic rocks, showing trace elements plotted against Mg#, defined as 100 3 MgO/(MgO 1 FeO), where FeO 5 FeOT 3 0.9.

Mafic Flows Major and Trace Elements The mafic rocks span a large range in Mg# [defined as 100 3 MgO/(MgO 1 FeO), where FeO 5 FeOT 3 0.9], from more primitive compositions with Mg# as large as 70 to evolved compositions with Mg# as small as 30. There is an overall trend to more primitive compositions upward in the sequence. The most primitive compositions occur in cycles 4 and 3, and the most evolved in cycles 1 and 3 (Figs. 5 and 6). Cycle 4 mafic rocks are divided into two groups, an eastern group and

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a western group, on the basis of outcrop distribution (Van Wagoner et al., 1994). The two groups have similar major element abundances but differ in their minor and trace element and REE abundances (Figs. 5 and 6). Major and trace element chemical variation in the mafic rocks can be largely explained by crystal fractionation involving olivine, clinopyroxene, iron-titanium oxides, and plagioclase. There is an increase in SiO2 and the incompatible trace elements (Nb, Y, and Zr) with decreasing Mg#. Ni and Cr decrease with decreasing Mg#, consistent with fractionation of olivine and clinopyroxene. Crystal fraction-


Rare Earth Elements All analyses are light rare earth element (LREE)–enriched with abundances from 30 to 130 times chondritic abundances (Fig. 7). Heavy REE (HREE) abundances are from 10 to 40 times chondritic abundances. Flows from cycles 1–3 have higher chondritenormalized REE abundances than cycle 4 flows (Figs. 7 and 8). Chondrite-normalized REE profiles are similar for all flows, except those from cycle 4 (west). Cycle 4 (west) flows have steep slopes with (Ce/Yb)n 5 6.0– 7.8, and cycles 1 through 4 (east) have less steep slopes with (Ce/Yb)n 5 2.0–4.0 (Fig. 8). Therefore the flows most enriched in REEs are found in cycle 1, the most evolved group, and those least enriched are found within cycle 4 (east). A flow from cycle 1 has a slight neg-


Figure 7. (A, C, and E) Chondrite-normalized REE abundances and (B, D, and F) primitive mantle–normalized trace element spider diagrams for the basaltic rocks of the Passamaquoddy Bay volcanic sequence; the Cranberry Island series basalts (shaded region in C) (Seaman et al., 1999) are shown for comparison. Chondrite-normalizing values are from Nakamura (1974), and primitive mantle– normalizing values are from Sun and McDonough (1989). ative Eu anomaly, suggesting that plagioclase was a fractionating phase, whereas the analysis from cycle 3 has a positive anomaly, indicating that some plagioclase accumulation has occurred. Compared to primitive mantle, mafic flows from cycles 1–3 are more enriched than those from cycle 4. Most flows display negative anomalies for Nb, Hf, and Sr (Fig. 7).

Felsic Volcanic Rocks In order to study the geochemical processes operating in a subvolcanic magma chamber, samples are best taken systematically from base to top through the volcanic sequence, from proximal to distal parts of ash-flow sheets (e.g., Hildreth, 1979; Smith, 1979), and from accompanying lava flows (e.g., Smith, 1979). In


young deposits, pumice separates from these samples provide the most reliable indicators of the true magma composition and avoid the problems of mechanical sorting of the pyroclastic deposit during magma fragmentation and transport (Wolff, 1985). In ancient deposits, such as those of the Passamaquoddy Bay area, however, the processes of lithification, alteration, and metamorphism make such sampling

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Figure 8. REE ratio diagrams for the mafic rocks, showing the variation in (A) LREE/ HREE and (B) LREE/LREE, plotted against (LREE/MREE) (MREE—middle REE). [La/ Ce]n is relatively uniform (note different scales of y-axes), but the basalts at the top of the section show significant variation in [Ce/Yb]n. (C and D) REE ratios are also plotted against total REE concentration.

more difficult. Pumice clasts are usually difficult to separate from a deposit, and whole-rock samples provide the best alternative. In our study, the results of thin-section examination of felsic pyroclastic rocks indicates that most were unsuitable for geochemical analysis because of crystal enrichment in some samples and the presence of small mafic, felsic, and sedimentary lithic clasts that could not be separated before pulverizing. Therefore, only vitric-rich, crystal-poor, and lithic-free samples were used in this study. As many of the samples from felsic pyroclastic rocks could not be used, our data set (Table 1) consists mostly of samples from felsic lava flows. These samples may not cover the range of compositions present in the magma chamber and may represent samples of the chamber taken at periods of disequilibrium between major ash-flow eruptions (Smith, 1979). Coherent patterns on both major (Fig. 9) and trace element diagrams (Fig. 10), however, suggest that the samples represent discrete stages in the evolution of the magma chamber,

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but probably do not span the range of compositions present prior to ash-flow events. Major and Trace Elements Variation diagrams use SiO2 as an indication of chemical evolution of the suite rather than Zr, a less mobile element, because of the role of crystal fractionation of zircon in these rocks. Major element versus SiO2 trends for each cycle overlap, suggesting that each cycle followed a similar chemical evolutionary path, although the cycle 4 dome shows the least overlap and trends toward the least evolved compositions of the felsic rocks. Overall, major elements TiO2, FeO*, CaO, Al2O3, MgO, and P2O5 decrease with increasing SiO2 (Fig. 9). However, within individual cycles, trends are variable (Fig. 9). The large ion lithophile elements (LILEs)—K2O, Rb, and Ba—show a scatter over the SiO2 range, reflecting the mobility of these elements. Owing to the alteration of all phenocryst phases except plagioclase, it is difficult to determine which fractionating mineral phases were important in controlling


the major element chemical variation. The decrease in TiO2 and FeO* suggests that irontitanium oxides were present, the decrease in CaO/Al2O3 ratio and Sr suggest plagioclase fractionation, and the decrease in P2O5 probably indicates apatite fractionation. Altered ferromagnesian minerals, including biotite, clinopyroxene, and rare amphibole, occur as rare phenocrysts, and these were probably also important in controlling chemical trends. Zr displays two trends with increasing SiO2 (Fig. 10). Cycle 3 shows increasing Zr with increasing SiO2, cycles 1 and 4 show decreasing Zr, and cycle 2 exhibits both behaviors. The decrease in Zr with increasing SiO2 likely indicates that zircon was a fractionating phase. In cycles 2 and 3, two groups can be recognized, a high-Zr group (Zr 5 420–1015 ppm) and a lower-Zr group (Zr 5 250–880 ppm). Zr abundances in the high-Zr group are extremely elevated for rocks that are classified as subalkalic. Several explanations are possible: (1) the high-Zr rhyolite may be a differentiate of the less evolved melts in which Zr behaved as an incompatible element; (2) there may have been two melts, reflecting source heterogeneity; (3) the high-Zr rhyolite may contain inherited zircon that was entrained from the source area; and (4) the variations may reflect zonation within the magma chamber. The behavior of Nb and Y in the rhyolitic rocks mimics that of Zr in most samples, suggesting that the high Zr is not controlled by inherited zircon. Similar abundances and trends on major element diagrams indicate that the high-Zr rocks are not more evolved than the low-Zr group. The differences, therefore, likely represent source heterogeneity. Several samples with high Zr/Y and Zr/Nb ratios may, however, contain anomalously high amounts of zircon. Rare Earth Elements All of the felsic rocks are LREE enriched; their LREE abundances range from 125 to 250 times chondritic abundances. The HREE abundances of the felsic rocks range from 15 to 50 times chondritic abundances. Rock units at the base of cycles 1, 2, and 4 have the flattest REE patterns and a small negative Eu anomaly. Upward within each cycle, the REE patterns become steeper, with both LREE and HREE enrichment, and the negative Eu anomaly increases in size (Figs. 11 and 12). These patterns are consistent with variation due to extraction from a zoned magma chamber, controlled by crystal fractionation. Two felsic flows (D1ff3 and D4ff4) have steeper MREE to HREE (MREE 5 middle REE) profiles than the other samples (Figs. 11


38.07 3.30 24.62 1.20 2.09 42.7 100.4 13.51 56.4 13.90 3.68 14.21 2.42 14.75 2.96 8.19 1.12 6.83 0.94 6.51 1.51 0.56 0.03 2.01

Li Be Sc Mo Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Tl Bi U

20.07 2.66 26.02 1.68 3.74 40.7 95.3 12.84 52.0 13.42 3.59 12.63 2.30 13.70 2.77 8.18 1.11 6.77 1.00 6.28 1.06 0.24 0.07 1.82

17 357 77 286 60 2 N.A. 27 136 11 0 N.A. 114 356

55.18 2.02 14.21 11.02 0.21 2.77 5.68 4.03 1.92 0.91 0.80 98.75

20.14 1.80 40.07 1.24 2.72 24.0 57.7 8.06 34.6 9.31 2.93 9.82 1.70 10.38 2.07 5.99 0.81 4.90 0.68 5.32 1.07 0.13 0.01 0.78

15 293 61 303 23 1 N.A. 21 127 18 12 9 404 249

49.61 3.62 13.38 13.44 0.21 4.51 8.28 3.34 0.92 0.57 1.30 99.18

19.55 1.62 40.98 0.76 3.55 24.2 59.2 8.41 36.9 9.71 3.50 9.63 1.69 10.25 2.00 5.77 0.74 4.62 0.68 4.22 2.04 0.22 0.05 0.62

14 261 54 312 32 3 N.A. 23 112 53 13 14 369 183

49.05 3.35 13.58 13.51 0.27 4.63 8.88 3.41 0.91 0.66 0.50 98.75

24.99 1.65 36.53 0.74 1.86 22.8 55.9 8.06 35.2 9.37 3.30 9.59 1.62 10.04 1.99 5.53 0.74 4.65 0.66 4.43 1.06 0.64 0.01 0.61

10 252 60 295 89 N.A. N.A. 21 144 19 20 16 323 324

48.55 3.27 13.62 12.91 0.27 4.70 8.37 2.51 2.56 0.72 1.20 98.68

15.32 2.08 49.83 1.60 7.34 32.3 75.0 10.41 44.1 11.38 5.20 11.12 1.94 11.95 2.49 6.97 0.99 6.41 1.00 9.61 0.84 0.53 0.07 2.03

16 469 68 266 75 6 N.A. 22 116 27 10 11 31 871

53.72 2.20 12.73 15.13 0.35 2.67 5.32 3.24 1.73 0.97 1.40 99.46

49.68 0.89 33.50 0.41 2.79 12.1 29.2 4.17 17.4 4.80 1.71 4.60 0.82 4.95 1.01 2.87 0.39 2.44 0.38 3.08 521.00 0.04 0.01 0.27

,5 133 22 282 5 N.A. N.A. 18 94 52 81 120 241 86

46.93 1.81 15.99 11.11 0.17 7.49 9.92 2.92 0.19 0.31 2.90 99.74

40.81 0.90 34.61 0.51 4.15 10.7 25.8 3.65 15.8 4.45 1.61 4.64 0.86 5.30 1.08 2.93 0.42 2.57 0.38 2.83 0.51 0.16 0.01 0.28

,5 136 23 400 24 N.A. N.A. 16 89 53 81 170 225 166

48.13 1.85 15.81 10.78 0.17 7.24 10.19 2.62 0.61 0.27 2.30 99.97

38.97 0.82 33.34 0.53 1.48 12.4 29.6 4.18 17.7 4.77 1.58 4.64 0.81 4.97 1.00 2.78 0.37 2.24 0.33 2.59 1.29 0.12 0.01 0.21

7 143 24 490 18 N.A. N.A. 15 107 42 83 120 242 176

47.52 1.82 15.72 11.00 0.18 7.33 9.32 3.29 0.48 0.30 2.50 99.46

25.27 1.53 47.97 0.57 1.01 19.1 45.2 6.42 27.8 7.64 2.44 7.76 1.41 8.74 1.78 4.93 0.67 3.98 0.56 4.05 1.02 0.12 0.03 0.63

10 247 44 297 11 N.A. N.A. 21 172 44 31 78 404 201

49.28 3.04 13.33 13.33 0.23 4.90 8.11 3.87 0.55 0.45 1.40 98.49

Note: Compound analyses are in wt%. Elemental analyses are in ppm. LOI—loss on ignition; N.A.—not available.

18 400 84 278 63 19 95 25 214 5 4 N.A. 76 405

56.47 1.81 14.36 10.74 0.21 2.43 4.94 3.78 2.08 0.83 1.20 98.85

Nb Zr Y Sr Rb Th Pb Ga Zn Cu Ni Cr V Ba

SiO2 TiO2 Al2O3 Fe2O3* MnO MgO CaO Na2O K2O P2O5 LOI Total

147.27 2.32 33.73 0.54 16.93 14.8 34.7 4.64 19.7 5.25 1.86 5.08 0.90 5.43 1.09 2.85 0.39 2.35 0.35 1.90 0.43 0.10 0.03 0.23

5 150 25 261 18 N.A. N.A. 18 97 54 79 141 315 64

45.04 2.07 17.12 11.24 0.23 5.12 9.09 1.53 0.44 0.34 8.90 101.12

61.18 0.74 33.21 0.70 3.95 10.4 25.3 3.68 15.7 4.56 1.57 4.62 0.84 5.07 1.01 2.85 0.40 2.38 0.34 2.12 0.41 0.02 0.01 0.19

9 135 24 265 ,5 N.A. N.A. 18 87 57 79 187 249 88

46.63 1.89 16.23 10.86 0.21 7.85 9.45 2.60 0.18 0.28 3.60 99.78

78.69 0.86 32.41 0.58 3.25 10.2 24.3 3.46 15.1 4.08 1.45 4.12 0.76 4.75 0.92 2.54 0.35 2.08 0.30 2.35 0.35 0.03 0.03 0.26

19 130 23 340 ,5 N.A. N.A. 16 84 43 82 189 235 46

47.87 1.78 16.29 10.19 0.18 7.50 8.79 3.01 0.18 0.28 4.70 100.77

104.47 1.35 34.83 0.46 3.46 10.2 24.9 3.63 15.6 4.47 1.57 4.72 0.86 5.40 1.60 2.97 0.41 2.61 0.37 3.35 0.62 0.02 0.06 0.37

5 144 24 374 ,5 N.A. N.A. 18 82 44 71 203 266 141

47.58 1.93 16.18 10.21 0.17 7.13 9.07 3.12 0.20 0.30 3.40 99.29

31.75 0.83 33.57 0.24 3.57 9.5 22.9 3.29 14.6 4.18 1.48 4.38 0.79 4.81 0.95 2.67 0.36 2.13 0.30 1.25 0.59 0.25 0.06 0.16

6 125 23 295 30 N.A. N.A. 16 115 25 94 146 259 203

46.98 1.68 16.74 11.18 0.19 8.43 9.28 2.75 0.67 0.27 1.80 99.97

50.10 0.93 34.86 0.47 4.26 11.9 29.0 4.21 18.2 4.79 1.57 4.88 0.85 5.20 1.03 2.88 0.38 2.43 0.35 2.86 0.55 0.07 0.01 0.27

,5 145 25 386 8 N.A. N.A. 15 91 42 84 138 246 188

47.02 1.92 16.46 11.04 0.18 7.28 8.66 3.30 0.33 0.34 3.30 99.83

21.76 1.02 48.34 0.44 1.58 20.0 47.2 6.54 27.1 7.03 2.24 6.72 1.21 7.65 1.52 4.26 0.60 3.49 0.51 2.95 0.89 0.11 0.03 0.54

11 221 43 375 21 ,10 ,10 21 115 48 41 122 364 140

49.10 2.54 14.08 13.20 0.21 5.75 9.65 3.77 0.53 0.41 0.80 100.04

21.70 1.23 46.47 0.44 1.56 19.8 47.4 6.62 27.1 6.93 2.34 7.21 1.28 7.93 1.59 4.27 0.56 3.42 0.51 3.01 0.79 0.12 0.04 0.55

8 217 41 377 21 ,10 ,10 20 116 44 39 113 373 148

49.06 2.53 14.14 12.56 0.21 5.78 9.66 3.62 0.55 0.42 0.60 99.13

30.41 0.52 33.74 0.67 6.94 11.8 28.4 4.08 17.1 4.68 1.65 4.52 0.81 5.06 1.00 2.79 0.38 2.32 0.33 2.30 0.90 0.11 0.01 0.16

5 140 25 344 17 ,10 ,10 19 94 25 89 128 250 87

47.05 1.81 15.97 11.24 0.17 7.63 10.59 2.54 0.38 0.31 1.90 99.59

9.89 1.58 44.59 1.15 8.84 22.7 52.4 7.14 29.7 7.97 2.45 7.86 1.42 8.81 1.81 5.28 0.73 4.74 0.75 5.74 1.12 0.17 0.05 0.88

10 267 49 269 36 ,10 ,10 19 127 39 25 42 426 225

51.64 2.82 13.55 13.36 0.21 4.77 8.72 3.08 1.03 0.44 0.01 99.63

53.93 0.96 35.61 0.47 2.78 39.0 93.1 12.62 50.2 11.07 3.05 8.84 1.30 7.08 1.35 3.69 0.49 3.02 0.45 4.16 0.67 0.13 0.13 1.45

8 220 35 614 23 ,10 ,10 19 131 33 122 266 250 339

45.83 1.91 16.29 10.37 0.24 8.66 8.77 2.85 0.90 0.50 2.90 99.22

48.87 1.04 33.55 0.57 3.15 29.6 69.1 9.26 37.8 8.62 2.50 7.22 1.08 6.35 1.22 3.29 0.45 2.95 0.43 3.63 0.68 0.31 0.01 0.91

6 172 32 519 44 ,10 ,10 19 99 28 106 220 244 322

47.38 1.64 15.66 10.57 0.22 7.54 9.56 2.95 1.36 0.39 2.50 99.77

27.38 0.99 32.79 0.65 1.74 34.1 82.2 11.23 47.2 11.01 3.16 8.68 1.28 7.09 1.34 3.66 0.48 2.94 0.45 3.28 0.63 0.10 0.09 1.10

,5 175 35 660 19 ,10 ,10 16 111 39 61 31 250 218

48.83 1.39 16.41 10.22 0.18 6.34 10.57 2.94 0.63 0.45 1.70 99.66

Sample no. 7M-202A 7T-235A 7T-176A 7T-220A 7M-110A 6N-127 8K-122B 8K-153A 8K-226A 8K-300A 8K-323A 8K-360A 8K-458A 8K-613A 8K-646A 8K-648A 8B-238A 8B-50A 8B-157A 8B-168A 8B-94C 8B-61A 8B-102A Unit D1mf1 D1mf1 D2mf1 D2mf1 D2mf1 D3mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf1 D4mf2 D4mf2 D4mf2

TABLE 1. REPRESENTATIVE ANALYSES OF MAFIC AND FELSIC VOLCANIC ROCKS FROM PASSAMAQUODDY BAY


409

410 TABLE 1. (Continued.)


10.86 4.30 18.34 0.99 0.87 50.8 116.0 14.58 54.9 13.57 2.74 12.73 2.30 13.84 2.76 8.27 1.19 7.97 1.19 10.57 0.09 0.49 0.04 2.55

Li Be Sc Mo Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Tl Bi U

11.07 2.96 18.35 1.73 1.15 60.8 133.3 16.25 59.3 14.14 1.24 13.99 2.53 15.58 3.28 9.59 1.38 8.63 1.30 7.79 2.04 0.71 0.04 4.13

24 296 96 90 100 40 6 20 90 N.A. 8 31 2 622

72.74 0.22 13.71 1.98 0.09 0.92 0.74 4.26 3.98 0.04 0.40 99.08

23.04 4.24 16.53 1.07 1.96 59.1 121.7 15.80 58.2 13.50 1.22 12.51 2.22 14.32 2.91 8.64 1.28 8.28 1.26 9.44 0.60 0.62 0.05 4.26

23 382 95 47 163 22 16 20 60 0 5 2 N.A. 568

72.58 0.26 13.76 2.46 0.06 0.80 0.49 4.43 4.44 0.05 0.90 100.23 20 362 81 145 29 39 9 21 36 2 5 19 4 215 10.58 4.78 5.49 1.17 2.21 45.3 95.8 11.60 39.0 8.97 0.58 8.56 1.79 12.43 2.80 9.09 1.42 9.52 1.46 13.08 1.26 0.84 0.09 6.94

22 498 94 50 184 30 11 19 48 10 8 14 5 246 6.34 4.31 3.23 1.75 1.67 43.7 89.1 10.96 37.2 8.50 0.63 7.22 1.36 8.42 1.85 6.11 0.98 6.86 1.02 10.89 0.66 0.68 0.03 5.31

20 414 68 55 124 29 9 15 30 6 5 24 8 244

72.62 69.38 73.16 0.26 0.33 0.28 13.59 15.89 13.70 2.94 3.07 2.79 0.08 0.05 0.02 0.85 0.89 0.68 2.04 0.15 0.35 5.57 4.44 4.67 0.89 5.35 4.15 0.05 0.05 0.04 0.40 0.60 0.50 99.29 100.20 100.34

14.03 31.27 4.36 3.49 10.54 15.53 2.18 1.68 1.17 3.97 56.6 47.0 126.1 104.8 15.76 13.09 58.6 49.5 14.19 12.01 1.95 1.89 12.58 11.40 2.40 1.87 14.90 10.21 3.03 2.00 9.49 5.68 1.42 0.75 9.14 4.77 1.35 0.63 11.87 2.93 1.50 1.44 0.74 0.78 0.02 0.01 3.53 3.78

24 611 103 77 152 35 23 22 249 7 8 25 10 628

71.12 0.35 13.58 3.83 0.09 1.04 0.50 3.61 4.90 0.05 0.30 99.37

9.54 4.80 6.14 3.82 0.65 55.3 124.9 16.02 59.4 14.21 2.30 13.46 2.45 14.79 2.97 8.44 1.15 7.55 1.08 9.72 1.02 0.80 0.21 2.58

28 746 103 61 99 19 N.A. 19 66 5 10 31 12 564

68.54 0.46 12.46 5.63 0.08 0.97 0.96 4.23 3.25 0.07 2.50 99.15 22 433 76 28 180 27 15 19 41 11 2 30 2 235

72.02 0.31 14.44 2.29 0.04 0.86 0.24 4.64 4.81 0.05 0.10 99.80 34 833 119 13 72 24 31 27 67 5 5 33 N.A. 82

74.15 0.25 12.76 3.88 0.03 0.67 0.08 5.42 2.42 0.03 0.30 99.99 39 904 124 51 207 N.A. N.A. 26 113 12 6 11 ,5 233

75.36 0.21 11.43 3.25 0.04 0.69 0.25 2.73 4.96 0.06 0.90 99.88

11.94 8.67 5.50 14.25 5.12 3.89 2.47 4.85 3.29 2.98 1.84 1.32 3.34 2.32 1.11 1.10 1.49 1.41 1.16 6.04 57.2 78.9 78.8 65.4 107.3 120.4 148.8 148.0 13.77 17.87 22.56 17.85 47.4 59.4 83.4 65.7 10.39 12.25 19.91 15.88 0.69 0.81 1.69 1.38 8.80 10.34 16.30 14.44 1.67 1.84 2.65 2.73 10.42 11.26 13.87 17.47 2.19 2.32 2.68 3.73 6.70 6.96 7.86 11.70 1.06 1.10 1.21 1.75 7.37 7.58 8.25 11.54 1.14 1.11 1.20 1.71 10.53 11.42 12.39 16.57 0.68 0.69 0.16 1.06 0.65 0.54 0.32 0.98 0.05 0.20 0.04 0.02 6.26 6.68 1.38 5.33

21 430 78 28 202 24 4 17 43 6 7 32 4 249

72.07 0.30 13.96 2.36 0.03 0.93 0.28 4.03 5.09 0.05 0.30 99.40

Note: Compound analyses are in wt%. Elemental analyses are in ppm. LOI—loss on ignition; N.A.—not available.

24 678 91 127 127 4 N.A. 22 59 4 5 3 N.A. 496

66.38 0.48 13.67 6.72 0.15 0.91 1.52 5.10 3.95 0.09 0.90 99.87

Nb Zr Y Sr Rb Th Pb Ga Zn Cu Ni Cr V Ba

SiO2 TiO2 Al2O3 Fe2O3* MnO MgO CaO Na2O K2O P2O5 LOI Total

26.27 4.22 8.50 2.34 2.07 71.6 137.6 18.34 66.3 15.51 1.61 14.31 2.65 16.03 3.28 9.96 1.44 9.09 1.33 11.05 0.95 0.81 0.03 5.25

25 464 92 205 94 N.A. N.A. 19 47 17 ,5 10 6 521

71.84 0.27 13.77 2.64 0.05 0.86 1.00 4.61 3.47 0.10 0.80 99.41

11.20 3.42 13.47 2.30 2.09 47.2 102.7 13.21 50.0 12.66 2.60 12.16 2.36 15.21 3.25 9.63 1.40 9.10 1.41 10.49 0.32 0.50 0.09 3.43

18 539 93 148 103 31 ,10 20 116 ,5 6 5 ,5 469

67.58 0.49 14.72 5.32 0.13 0.96 2.41 4.29 3.16 0.17 0.60 99.83

12.72 3.93 4.70 2.14 2.72 44.8 99.2 12.41 46.6 11.93 1.42 11.52 2.29 14.79 3.12 9.44 1.39 8.86 1.34 5.94 0.41 0.66 0.13 4.37

20 285 95 62 147 37 ,10 19 83 ,5 5 ,5 ,5 560

72.70 0.20 13.80 2.25 0.06 0.62 0.95 4.42 4.20 0.10 0.30 99.60

20.72 3.56 3.26 1.17 2.95 51.7 107.4 14.26 53.4 13.47 1.32 12.94 2.52 15.87 3.38 10.27 1.56 9.79 1.50 7.85 0.71 0.50 0.05 4.51

,5 256 93 36 155 N.A. N.A. 19 78 11 ,5 12 ,5 546

73.41 0.13 13.34 1.68 0.04 0.66 0.50 4.93 4.11 0.09 0.30 99.19

31.61 3.27 16.14 13.21 4.04 49.6 109.7 14.08 52.3 12.65 1.69 10.40 1.79 10.58 2.01 5.47 0.72 4.33 0.61 2.93 0.66 0.84 0.02 3.46

18 363 73 106 150 N.A. N.A. 19 97 14 ,5 12 ,5 630

71.44 0.34 13.77 3.93 0.14 0.86 0.77 3.84 4.00 0.35 0.60 100.04

9.49 3.43 9.56 1.33 3.53 49.7 106.8 13.72 50.7 11.96 1.93 10.99 1.94 11.78 2.43 7.56 1.21 8.07 1.24 9.66 0.34 0.51 0.10 4.90

16 566 81 96 132 N.A. N.A. 19 76 10 ,5 14 ,5 621

70.05 0.29 14.05 2.55 0.07 0.70 1.34 4.88 3.70 0.09 1.70 99.42

23.86 3.66 14.35 2.45 1.71 41.0 91.4 11.68 45.1 11.68 2.49 11.56 2.14 13.91 2.87 8.79 1.30 8.39 1.31 9.77 1.10 0.51 0.11 3.17

15 547 81 147 110 N.A. N.A. 21 112 16 ,5 13 10 520

65.58 0.66 14.56 5.63 0.16 1.20 2.33 5.02 3.20 0.27 0.80 99.41

24.47 2.97 19.96 2.04 2.32 37.4 84.6 11.20 44.3 11.40 2.95 11.78 2.17 13.83 2.92 8.50 1.21 7.90 1.13 8.94 0.78 0.41 0.03 2.76

16 420 83 200 91 10 ,10 21 119 ,5 8 ,5 9 434

63.75 0.85 14.75 7.50 0.17 1.48 3.26 4.59 2.72 0.36 0.60 100.03

Sample no. 7T-210A 7M-20D 7T-245A 7M-19H 7M-131 6N-68 H-30B 7M-162B 6N-133 6N-80B HW-12 7T-10A 8K-181A 8B-96A 8B-156A 8K-594A 8K-131B 8K-567A 8K-574A 8B-35A Unit D1ff1 D1ff2 D1ff3 D1ff3 D1ff3 D2xlt1 D2tb1 D2ff1 D2ff4 D2ff4 D3vt1 D3ff1 D4vt1 D4ff2 D4ff3 D4ff3 D4ff4 D4td1/A D4td1/B D4td2

45.66 2.96 20.24 2.01 4.86 33.5 75.9 10.59 43.4 11.30 3.49 11.10 1.99 12.08 2.44 7.01 0.99 6.34 0.94 1.30 0.56 0.51 0.06 1.26

20 1364 76 262 110 14 17 21 102 10 6 18 N.A. 568

61.70 1.00 16.00 8.28 0.17 1.71 2.64 4.52 2.48 0.22 0.80 99.52

6N-61 Dsg

3.31 3.26 0.90 1.18 1.45 60.7 130.0 17.28 64.8 14.59 2.20 13.56 2.42 14.72 2.96 8.53 1.21 7.83 1.13 8.72 0.82 0.47 0.06 2.89

25 491 85 74 97 41 28 17 33 N.A. 4 28 1 632

76.02 0.16 12.22 1.26 0.06 0.79 0.75 3.99 3.73 0.02 0.30 99.30

45.82 3.59 5.15 3.94 11.84 43.3 74.7 12.09 44.6 10.32 1.12 8.83 1.34 7.14 1.29 3.52 0.51 3.33 0.48 4.81 0.71 1.61 0.05 2.21

26 421 55 19 190 21 37 21 126 N.A. 14 63 32 327

76.97 0.42 9.90 3.67 0.11 1.43 1.52 0.82 3.90 0.03 2.10 100.87

7M-513A 7M-557C D1ff2 D2ff2

VAN WAGONER et al.


terbedded, littoral-facies sedimentary rocks throughout the sequence indicates that the area remained at or near sea level throughout the volcanic activity and that the basin was continually subsiding (Van Wagoner et al., 1994). Analyses of both mafic and felsic volcanic rocks have continental within-plate tectonic affinities (Figs. 13 and 14). Despite negative Nb on primitive mantle–normalized diagrams (Fig. 7), there is no other geochemical indication that the mafic rocks are arc related. Relationship to Other Silurian–Devonian Volcanic Rocks

Figure 9. Chemical variation diagrams for the rhyolitic rocks, showing major element oxides plotted against SiO2. and 12). The preferential decrease of the HREE is most likely controlled by zircon fractionation. Compared to primitive mantle, the felsic rocks are depleted in Nb, Ta, Ti (strongly depleted), Ba, Sr, and P. These characteristics are typical of felsic rocks derived from continental crust. Flows D1ff3 and D4ff4 are also depleted in Hf, consistent with extensive fractionation of zircon. DISCUSSION Tectonic Affinity Interpretation of the tectonic affinity based on tectonic discrimination diagrams must be used in conjunction with a knowledge of the

geology of the area. The stratigraphy and physical volcanology of the Passamaquoddy Bay area most resembles that of continental extensional or hotspot-related environments such as the Yellowstone–Snake River Plain region (Geist and Richards, 1993; Vetter and Shervais, 1992). The sequence lacks the large volumes of reworked volcaniclastic material that are typically interstratified with volcanic rocks in arc terranes. The juxtaposition of silicic proximal and distal volcanic facies suggests that there were several small volcanic centers in the area (Van Wagoner et al., 1994). No caldera complexes have been identified. Mafic flows appear to have come from a single source region to the north of the Passamaquoddy Bay region. The occurrence of in-


Other bimodal Silurian–Devonian volcanic successions occur within the Coastal volcanic belt. These include the Silurian (424 Ma) Cranberry Island series in Maine (Seaman et al., 1995, 1999) and the Eastport, Leighton, Edmunds, and Dennys Formations of Maine (Amos, 1963; Gates and Moench, 1981). The only other detailed geochemical study is from the Cranberry Island series. The Cranberry Island series comprises ;1.8 km of felsic pyroclastic rocks and lava flows overlain by ;0.8 km of mafic tuff and lava flows (Seaman et al., 1999). These rocks are chemically similar to some of the Passamaquoddy Bay volcanic sequence (Figs. 7 and 11). The Cranberry Island basalts are also tholeiitic and have within-plate basalts affinities (Seaman et al., 1999). The mafic flows from Cranberry Island are chemically similar to the flows from our cycle 4 (east), although the Cranberry Island mafic unit is dominated by pyroclastic rocks, and the Passamaquoddy Bay cycle 4 (east) is dominated by flows. The Cranberry Island felsic rocks are similar to some of the Passamaquoddy Bay felsic rocks, but the former trend to lower chondrite-normalized REE contents (Fig. 11). The Eastport Formation of Maine has been correlated with the Passamaquoddy Bay volcanic sequence by Cumming (1967), Hay (1967), Ruitenberg, (1968), and Pickerill and Pajari (1976), on the basis of similarities in lithology and fossil assemblages. The Eastport Formation overlies other bimodal volcanic units in Maine, including the Dennys, Leighton, and Edmunds Formations. Although these units document a significant succession of bimodal volcanic rocks within the Coastal volcanic belt, the exact thickness, age, and geological and chemical characteristics of these formations are not well known. Petrogenesis Mafic Petrogenesis There are currently several favored models to explain the chemistry and petrogenesis of

411

VAN WAGONER et al.

Figure 10. Chemical variation diagrams for the rhyolitic rocks, showing trace elements plotted against SiO2.

continental, within-plate magmas. These include modification of mantle melts by fractional crystallization accompanied by crustal contamination (AFC model—assimilation and fractional crystallization) (e.g., Cox and Hawkesworth, 1985; DePaolo, 1981; Dupuy and Dostal, 1984; Fodor et al., 1990; Marsh, 1989; Thompson et al., 1984), fractional crystallization accompanied by replenishment of the chamber with primitive magma (RTF model— replenished, tapped, and fractionated magma) (Cox, 1988; Hogg et al., 1989; O’Hara and Matthews, 1981), and mantle metasomatism prior to partial melting (Dostal et al., 1989; Pegram, 1990; Sheth, 1999). Mantle metasomatism includes that associated with alkalic magmatism (e.g., Hawkesworth et al., 1983)

412

and contamination of the mantle during a prior subduction event (e.g., Dostal et al., 1989; White and Hofmann, 1982). The model chosen must be consistent with geochemical characteristics as well as the petrogenetic model proposed for the felsic volcanic rocks. The major and trace element geochemistry of the Passamaquoddy Bay volcanic sequence clearly illustrates that much of the variation in the mafic volcanic rocks is due to crystal fractionation. Although there are overlapping trends, there is an overall trend toward more primitive compositions upward in the section. This variation can be explained by periodic replenishment of small magma chambers with more primitive magma, at the onset of each new volcanic cycle. This model of replenish-


ment, tapping, and fractionation (the RTF model) has been applied to the American Bar Flows of the Imnaha Basalt, Columbia River Basalt Group, by Hooper (1988) and to the Tertiary volcanic province of Greenland (Hogg et al., 1989). The RTF model is consistent with the cyclic nature of volcanism and with the change upward in the sequence to more primitive magma compositions. Hogg et al. (1989) suggested that the RTF process is most efficient in small magma reservoirs, often located in a densely fractured continental crust. Small magma reservoirs are indicated in the Passamaquoddy Bay area by the presence of thin, low-volume mafic lava flows and multiple source areas for both the mafic and felsic volcanic rocks (Van Wagoner et al., 1994). The mafic rocks are tholeiitic, and LREEenriched,with steep chondrite-normalized REE patterns ([Ce/Yb]n . 1) characteristic of partial melting of a garnet-bearing mantle source. As the REEs are incompatible in the major crystallizing phases (olivine, clinopyroxene, plagioclase, and iron-titanium oxides), the subparallel REE trends for cycles 1–4 (east) result from variations in crystal fractionation. Compared to the rest of the basalts, the cycle 4 (west) basalt flows are even more LREE-enriched ([Ce/Yb]n 5 6–7.8) with even steeper chondrite-normalized REE slopes and are some of the most primitive rocks in the area, and form some of the thickest individual flows. The strong LREE enrichment shown in the cycle 4 (west) REE patterns is likely due to smaller amounts of partial melting of a mantle source that was compositionally similar to the source for the earlier cycles (as shown by the similarity in La/Ce ratios among all the basalt flows; Fig. 8). Probably the most controversial aspect of continental-tholeiite petrogenesis is whether continental tholeiites represent (1) MORB (mid-oceanic-ridge basalt) or OIT (oceanicisland tholeiite) mantle partial melts that have been contaminated during upwelling through continental crust (e.g., Dessureau et al., 2000; Dupuy and Dostal, 1984; Fodor and Vetter, 1985; Mantovani et al., 1985; Papezik et al., 1988), (2) melting of mantle that has been hydrated by a previous subduction event involving subducted continentally derived sediment (Dostal et al., 1989; Dunn and Stringer, 1990; Hergt et al., 1989; Pegram, 1990), or (3) partial melting of enriched metasomatized mantle (e.g., Doe et al., 1982; Menzies et al., 1991; Menzies and Wass, 1983; Turner and Hawkesworth, 1995). Compared to MORB and OIT, continental tholeiites are generally enriched in the mobile incompatible elements Rb, K, Ba,


Figure 11. (A, C, and E) Chondrite-normalized REE abundances and (B, D, and F) primitive mantle–normalized trace element spider diagrams for the rhyolitic rocks of the Passamaquoddy Bay volcanic sequence; the Cranberry Island series rhyolites (shaded region in C and D) (Seaman et al., 1999) are shown for comparison. Chondrite-normalizing values are from Nakamura (1974), and primitive mantle–normalizing values are from Sun and McDonough (1989). and Th, have a wide range in initial 87Sr/86Sr ratios, are LREE enriched, and are typically depleted in Nb, Ta, and, to a lesser extent, Ti (Dupuy and Dostal, 1984). These characteristics are typical features of the continental crust, leading many workers to invoke crustal contamination as a factor in their origin (Alle`gre et al., 1999; Hansen and Nielsen, 1999; Peate and Hawkesworth, 1996; Thompson et

al., 1984). However, other studies have suggested that because these characteristics are also typical of island-arc rocks, continental tholeiites may represent partial melts of a subduction-metasomatized mantle source (Pegram, 1990). The fact that negative Nb anomalies are found in both continental tholeiites and volcanic arc rocks may indicate a strong genetic link between subduction pro-


cesses and the mechanisms of continentalcrust generation (Pegram, 1990). The mafic rocks of the Passamaquoddy Bay volcanic sequence have typical continentaltholeiite chemistry (Fig. 7). On primitive mantle–normalized diagrams, the mafic rocks display marked enrichment in the mobile incompatible elements (Rb, K, and Ba). There is a marked depletion in Nb for all the rocks

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Figure 12. REE ratio diagrams for the felsic rocks, showing the variation in (A) LREE/ HREE and (B) MREE/HREE, plotted against (LREE/MREE). (C and D) REE ratios are also plotted against total REE concentration.

and an extreme depletion in Sr for the most evolved rocks. Some of the samples exhibit a trough at Ti (Fig. 7). These spider-diagram patterns are fairly typical for continental-crust rocks (Thompson et al., 1984). Given the fairly low trace element contents of MORB and OIT, relatively small additions of crustal rocks can result in a pattern that mimics that of the crust (Hergt et al., 1989). On the incompatible element versus incompatible element plots, the mafic flows plot in scattered linear arrays that trend away from mixing lines between typical depleted (normal- or N-type) and enriched (plume- or Ptype) mantle sources (Fig. 15, C and D). The most likely explanation is that crustal contamination is responsible for the arrays, with progressive contamination causing greater displacement from the mixing lines, as was shown by Fodor and Vetter (1985) for continental tholeiites from the Parana´ basin of Brazil. However, identifying such a process is strictly conditional on the choice of contaminant. In addition, there is also the question of whether assimilation takes the form of bulk assimilation of crustal rocks or some form of mixing between the mantle melts and partial

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melts of the continental crust (Thompson et al., 1984). For bulk assimilation to be a viable process, large amounts of heat are required, which could be supplied by the latent heat of crystallization, as in the AFC process of DePaolo (1981). This alternative implies that the most evolved rocks should also be the most contaminated. However, for the mafic rocks of the Passamaquoddy Bay volcanic sequence, there is no obvious correlation between Mg# and the range in incompatible element ratios, especially the critical Nb/La ratio (Fig. 15A). However, if the negative Nb and Ti anomalies were the result of crustal contamination, we would expect to see progressively lower Nb/La ratios with increasing fractionation resulting from such contamination (Kerrich et al., 1999), which is clearly not the case (Fig. 15). Crustal contamination may also have occurred as a result of mixing between mantle melts and partial melts of the continental crust (felsic magma). However, the difficulty arises in trying to determine a suitable contaminant, especially given that the basement rocks for the Avalon terrane have not yet been characterized. The only real clue to the nature of the


continental crust beneath the Passamaquoddy Bay area is the felsic rocks in this study that represent partial melts of the continental crust (see subsequent section). Although mobility of Rb may be a factor, on a plot of Nb/Y versus Rb/Y (not shown), the mafic flows display greater variation in Rb/Y compared to Nb/Y, a trend contrary to mantle melts but similar to syncollisional granites (Pearce et al., 1984) and the felsic rocks of the study area. Cox and Hawkesworth (1985) studied continental tholeiites from the Deccan Traps and showed two sequences with good isotopic evidence for crustal contamination that have trace element variations similar to those of the mafic flows of the Passamaquoddy Bay volcanic sequence. Three other Deccan Traps sequences with good isotopic evidence for mixing of normalmantle melts with those from an enrichedmantle source displayed a trend to higher Nb/ Y compared to Rb/Y. The mafic flows of the Passamaquoddy Bay volcanic sequence are moderately to highly evolved (Figs. 5 and 6; Table 1), show a wide range in some incompatible element ratios, and are associated with large volumes of felsic rocks indicative of crustal anatexis. Thus, the mafic magmas probably resided within or at the base of the continental crust for long periods of time, providing sufficient heat to produce crustal anatexis, and therefore had ample opportunity to become contaminated. However, some of the chemical characteristics of the mafic flows of the Passamaquoddy Bay volcanic sequence are somewhat more difficult to reconcile with a model of crustal contamination. The mafic flows have relatively high concentrations of the HFSEs (high field strength elements), in particular, TiO2 (Fig. 7, Table 1). Previous studies of continental tholeiites have identified rocks characterized by both high and low TiO2 (Dupuy and Dostal, 1984; Fodor et al., 1990; Fodor and Vetter, 1985). In these studies, the rocks containing low TiO2 concentrations were modeled as having undergone some form of contamination, whereas the high-TiO2 rocks in all cases had very low 87Sr/86Sr ratios and were assumed to be uncontaminated (Dupuy and Dostal, 1984; Fodor et al., 1990; Fodor and Vetter, 1985). Therefore, the evidence for crustal contamination for the Passamaquoddy Bay mafic flows is by no means unequivocal. In most instances, negative Nb anomalies are associated with volcanic arc basalts (Pearce, 1996). The Passamaquoddy Bay basalts are clearly not arc related (Figs. 7 and 13). However, metasomatism of the mantle by a previous subduction event, in the Early Silurian or Ordovician (e.g., Fyffe et al., 1999; Keskin et


Figure 14. Tectonic discrimination diagrams for granitic rocks from Pearce et al. (1984), showing the within-plate affinity of the rhyolitic rocks of the Passamaquoddy Bay volcanic sequence. Abbreviations: ORG—orogenic granites, syn-COLG—syncollision granites, VAG—volcanic arc granites, WPG—within-plate granites.

al., 1998; Murphy et al., 1999), could have caused the observed relative Nb depletion of the mantle with respect to other HFSEs and REEs. Therefore, an alternative model to crustal contamination would involve asthenospheric partial melts interacting with subductionmodified sublithospheric mantle. The timing of the modification of the sublithospheric mantle is undetermined; estimates range from Precambrian to Ordovician for this part of North America (Dostal et al., 1989; Pegram, 1990).

Figure 13. Tectonic discrimination diagrams from (A) Cabanis and Lecolle (1989), (B) Meschede (1986), and (C) Pearce and Norry (1979), showing the within-plate affinity of the mafic rocks of the Passamaquoddy Bay volcanic sequence. E-type MORB—enrichedtype MORB, N-type MORB—normal-type MORB.


Felsic Petrogenesis A number of models have been proposed to explain the generation of felsic magmas in bimodal volcanic suites, including crystal fractionation, combined crystal fractionation and assimilation, liquid immiscibility, fractional partial melting, and crustal anatexis (Doe et al., 1982; Hildreth, 1979; Pedersen et al., 1998). Of these models, fractional crystallization of basaltic parental magmas or crustal

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anatexis due to underplating of the continental crust by mantle-derived melts are generally the most accepted models to explain rhyolitic petrogenesis in bimodal volcanic suites. The geochemical and geological evidence from this study favors the generation of the rhyolitic magmas by crustal anatexis. Evidence includes the large SiO2 compositional gap (6% SiO2, or 10% SiO2 excluding the latestage domes), the large volume of felsic volcanic rocks relative to basaltic rocks, and the highly evolved trace element compositions of the rhyolites. Our preferred petrogenetic model involves injection of mafic magma into the lower crust, producing rhyolitic magmas through partial anatexis. This model has been suggested for a number of bimodal volcanic suites including those of the Yellowstone Plateau (Doe et al., 1982) and the southern British Caledonides (Leat et al., 1986). The petrologic processes involved in producing rhyolitic magmas by crustal anatexis were quantified by Huppert and Sparks (1988). The production of rhyolitic rocks from crustal material may further explain the cyclic nature of the volcanism in the study area, with periodic injection of basaltic magma into the crust provided the heat for crustal melts, producing felsic magma and triggering felsic eruptions (Huppert and Sparks, 1988; Sparks et al., 1977). In cycles 1–3, the stratigraphically lowermost unit is the least evolved member of the cycle. Above the lowermost unit in each of these cycles, the trends are toward more evolved compositions. This pattern suggests that within each cycle, volcanism sampled a continuously evolving or stratified magma chamber. Variations within each unit can be explained by magma-chamber zonation, uneven drawdown, and mixing of layers within the chamber during eruption. The overlapping trends in cycles 1–3 suggest that a new melt or melts were generated from a similar crustal source at the beginning of each cycle. Cycle 4 spans the range of cycles 1–3, but does not show consistent trends with stratigraphic order as a result of the complexity of the processes indicated. Implications for Tectonic Setting The Silurian–Devonian history of the northern Appalachians is the subject of considerable debate (van Staal and de Roo, 1995). Although the Coastal volcanic belt is considered to be at the northern margin of the Avalon terrane, the largest terrane of the Canadian Appalachians (Murphy et al., 1999), the precise terrane association is uncertain, as is the

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Figure 15. Variation of (A) [Nb/La]pm with Mg# and (B) with SiO2. (C) and (D) Mixing lines between blocks representing compositions of plume-type (P-type) and normal-type (N-type) mantle. Note that the basalts of the Passamaquoddy Bay volcanic sequence plot on trends away from the mixing lines.

timing of the collision between Avalon and Laurentia (Seaman et al., 1999). This study and that of other related rocks in Maine (Gates and Moench, 1981; Seaman et al., 1999) support continental extension. The distribution of bimodal volcanic rocks within the Coastal volcanic belt suggests (1) that during the Late Silurian, and possibly into the Devonian, this was a region of significant continental extension and related volcanism and (2) that such extension was not confined to what is now the Coastal volcanic belt, but occurred in several belts in New Brunswick, the Gaspe´ region, and Maine (e.g., Dostal et al., 1989). The precise cause of the extension is unclear. There is evidence for significant sinistral transcurrent faulting associated with docking of terrane(s) during the Late Silurian–Early Devonian (e.g., de Roo and van Staal, 1994). There has been speculation that volcanism was caused by the tectonics of transpressional pull-apart basins (Dostal et al., 1993). However, limited modern analogues associated with volcanism in this type of tectonic setting indicate that the volume, thickness, and chem-


istry of the erupted products are not similar to the Coastal volcanic belt. Because of the apparent age association with terrane convergence during the Acadian orogeny, it has been speculated that a backarc setting provides the required extension within this context (Fyffe et al., 1999). Murphy et al. (1999) further speculated on the presence of a plume-modified subduction zone, although the timing of events is inconsistent with the current age data for bimodal volcanic rocks of the Coastal volcanic belt. Our geochemical data are consistent with a mantle source previously metasomatized by slab-derived fluids. CONCLUSIONS We have shown that the Passamaquoddy Bay volcanic sequence is bimodal (basalt and rhyolite). The volcanic sequence has a considerable thickness (4 km) and has an age and chemistry similar to other bimodal volcanic sequences within the Coastal volcanic belt. The basaltic lavas of the Passamaquoddy Bay volcanic sequence were derived by partial


melting of mantle rocks, possibly enriched in mantle-incompatible elements and modified by a combination of crustal contamination and mantle metasomatism by a subduction-related component. The basalts become more primitive upward in the section and flows become thicker, indicating increasing eruptive rates and shorter residence times in the magma chamber, although overall eruptive volume decreased in later cycles. The rhyolitic magmas were derived by crustal anatexis driven by underplating of the continental crust by mantle-derived basaltic magmas. The variation in incompatible elements within particular felsic units and through time can be explained by crystal fractionation and by tapping of zoned magma chambers. Volcanism occurred in an extensional withinplate tectonic setting, but we are uncertain of the precise nature of the tectonic setting. Neither the duration nor the extent of this volcanism within the northern Appalachians is known, yet both are critical to understanding the region’s tectonic history. More work on geologic mapping, geochemistry, age dating, and isotopic studies are required to resolve these issues. ACKNOWLEDGMENTS This study was funded by the Canada–New Brunswick Mineral Development Agreement (MDA) 1984–1989, a National Science and Engineering Research Council of Canada grant (to Van Wagoner), and contributions from the New Brunswick Department of Natural Resources, LAC Minerals, and Acadia Minerals Ventures. We thank the reviewers, Brendan Murphy, Sheila Seaman, Peter Copland, and Scott Baldridge, whose thoughtful comments helped to improve the quality of the manuscript, and the editorial staff of the GSA Bulletin. This is publication number 255 in the GEMOC ARC National Key Centre (www.es.mq.edu.au/ GEMOC/). REFERENCES CITED Alle`gre, C.J., Birck, J.L., Capmas, F., and Courtillot, V., 1999, Age of the Deccan Traps using 187Re/187Os systematics: Earth and Planetary Science Letters, v. 170, p. 197–204. Amos, D.H., 1963, Petrology and age of plutonic rocks, extreme southeastern Maine: Geological Society of America Bulletin, v. 74, p. 169–194. Bevier, M.L., 1989, Preliminary U-Pb geochronologic results for igneous and metamorphic rocks, New Brunswick: Fourteenth annual review of activities: New Brunswick Minerals and Energy Division, Department of Natural Resources Information Circular 89–2, p. 190–194. Bevier, M.L., and Whalen, J.B., 1990, Tectonic significance of Silurian magmatism in the Canadian Appalachians: Geology, v. 18, p. 411–414. Bird, J.M., and Dewey, J.F., 1970, Lithosphere plate–continental margin tectonics and the evolution of the Appalachian orogen: Geological Society of America Bulletin, v. 81, p. 1031–1060.

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