The ca. 620 Ma Huntington Mountain pluton intruded the northeastern part of the East Bay Hills Group (Fig. 1). The pluton is dominated by diorite, but also contains granodiorite, leucogranite, monzogranite, and syenogranite. The East Bay Hills Group forms an elongate belt along and inland from the southeastern shore of Bras d’Or Lake. This group is dominated by andesitic to dacitic rocks with lesser quantities of basaltic to rhyolitic tuffs and flows and epiclastic sedimentary rocks. Field relationships suggests that the Huntington Mountain diorite was emplaced first, followed by granodiorite, leucogranite, and syenogranite; the small body of monzogranite may have been emplaced at some time between the granodiorite and syenogranite (Barr et al. 1996). Earlier basaltic, andesitic-dacitic, and dacitic rocks in the East Bay Hills Group have a thermal overprint attributed to pluton emplacement. Minor pyrite, chalcopyrite, and malachite in the basaltic rocks are associated with the Huntington Mountain pluton (Barr et al. 1996). Later basaltic andesite and rhyolite are considered to be coeval with pluton emplacement as all units are dated at ca. 620 Ma (Barr et al. 1990; Keppie et al. 1990; Bevier et al. 1993). All rocks in the study area are characterized by some degree of post-magmatic alteration.

Plagioclase and hornblende are the main primary minerals in the Huntington Mountain diorite and granodiorite; K-feldspar, quartz, and biotite occur in lesser quantities. Primary phases in the leucogranite are dominantly perthitic K-feldspar, quartz, and plagioclase, and in the syenogranite, K-feldspar, quartz, and plagioclase. These rocks have undergone weak to strong secondary alteration. Basaltic to andesitic rocks of the East Bay Hills Group are typically aphanitic and massive, with slight to moderate porphyritic textures. Phenocrysts consist mostly of plagioclase and rarely, clinopyroxene. Plagioclase-rich groundmass is commonly overprinted by secondary phases and in a few localities, all primary features have been obscured. Dacite is typically aphanitic to weakly porphyritic, consisting mainly of plagioclase phenocrysts suspended in a plagioclase-rich groundmass with minor quartz. Dacite typically exhibits moderate to strong alteration. Rhyolite samples are dominated by ash to lapilli crystal tuff; aphanitic to weakly porphyritic flow-banded pyroclastic beds are also present. Rhyolite consists of feldspar microlites, quartz, and minor K-feldspar, and exhibits only weak alteration.

Three main alteration types were observed. Type 1 consists of fine-grained chlorite, epidote, sericite, and Fe-Ti oxides (Figs. 2a, b, c), which are characteristic of propylitic alteration (Meyer and Hemley 1967). This alteration is widespread and varies in intensity from weak to strong (partial to complete replacement of primary phases). Type 2 consists of fine- to coarse-grained, anhedral quartz, sericite, and calcite, with minor Fe-oxides. This moderate to strong alteration is highly localized, and has partially to completely replaced primary minerals and textures (Fig. 2d). Type 2 alteration appears to overprint the propylitic alteration (Fig. 2e). Type 3, which was observed only rarely, is typical of phyllic alteration (Meyer and Hemley 1967). Most primary minerals have been moderately to strongly overprinted by fine-grained euhedral quartz, sericite, and pyrite (Fig. 2f). Type 3 alteration occurs mainly in the contact zone between the Huntington Mountain pluton and East Bay Hills Group, suggesting its development during contact metamorphism. Contact metamorphic features in the East Bay Hills dacite, including hornfelsic texture and pyrite mineralization, were previously reported by Barr et al. (1996). Such alteration is a common product of interaction with magmatic-dominated fluids in the hottest parts of a hydrothermal system (~ 350– 500 °C; Harris and Golding 2002).

The δ¹⁸O_WR results for the Huntington Mountain pluton range from –1.5 to +7.1‰ (Table 1a, Fig. 3a, 4): diorite and granodiorite, –1.5 to +4.5‰; syenogranite, –1.2 to +6.2‰; leucogranite, +3.8 to +7.1‰ (except for one sample, +0.2‰). Two leucogranite samples have δ²H_WR values of –60 and –56‰, whereas values are lower for other plutonic rock types (–78 to –61‰). The East Bay Hills Group has a wider range of δ¹⁸O_WR values (–3.8 to +9.5‰) than the plutonic suite: rhyolite-andesite and dacite, –3.8 to +6.5‰; basaltic rocks, –1.9 to +9.5‰; rhyolite, +4.6 to +6.9‰ (Table 1b, Fig. 3b). The δ²H_WR values of the volcanic rocks range from –93 to –70‰. In Figure 4 no clear pattern is observed between δ 18 O_WR values and the intrusive sequence and/or rock types in this complex.

Primary quartz (qtz) in the plutonic rocks has δ¹⁸O_qtz values ranging from +0.5 to +7.0‰, with leucogranite having higher values (≥ +4.8‰) than other rock types. The pattern for alkali feldspar (fs) is similar (leucogranite δ¹⁸O_fs , +4.4 to +6.0‰; other rock types, +1.4 to +4.6‰). Hornblende (hbl) has δ¹⁸O_hbl values of +1.4 to +5.5‰. The δ¹⁸O values for phenocrysts from volcanic samples also vary widely: quartz, –1.5 to +8.2‰; alkali feldspar, –0.3 to +6.5‰.

Chlorite (chl) δ¹⁸O_chl values vary widely: granodiorite, syenogranite and andesite, –6.3 to –4.0‰; diorite and leucogranite (–5.2 to +3.3‰). The δ²H_chl values for all rock types range from –77 to –69‰. Disseminated secondary quartz from a diorite sample affected by phyllic alteration has a δ¹⁸O_qtz value of +3.3‰, whereas values for basaltic to andesitic rocks and leucogranite affected by quartz-sericitecalcite alteration are much higher (+11.1 to +15.1‰). Disseminated calcite (cal) from plutonic rocks has δ¹⁸O_cal values of +5.4 to +8.4‰, and from volcanic rocks, +5.8 to +11.7‰. Disseminated calcite from the plutonic rocks has δ¹³C_cal values of –3.8 to +0.1‰; values for the volcanic rocks are lower (–7.2 to –4.5‰).

Rare vein quartz associated with phyllic alteration of diorite has a δ¹⁸O_qtz value of –2.1‰. Veins are more abundant in the volcanic rocks, with δ¹⁸O_qtz values ranging from +6.1 to +10.5‰. Vein calcite δ¹⁸O_cal and δ¹³C_cal values range from +6.0 to +14.3‰, and –7.4 to –3.2‰, respectively. As for the disseminated calcite, the vein calcite samples most enriched in ¹⁸O are also the most depleted of ¹³C.

Most Huntington Mountain area samples have δ¹⁸O_WR values that are lower than typical for “normal” (+6 to +10 ‰, Taylor 1974) igneous rocks, thus following the pattern of ¹⁸O-depletion reported for the Avalon terrane by Potter et al. (2008a, b). Most commonly, igneous rocks with abnormally low δ¹⁸O values have experienced post-crystallization alteration (see Taylor 1974), but crystallization from a low- ¹⁸O magma (see Taylor 1986; Bindeman and Valley 2000) cannot be ruled out a priori. In the latter case, high-temperature oxygen isotopic equilibrium should produce primary compositions in which (i) δ¹⁸O_qtz > δ¹⁸O_fs > δ¹⁸O_hbl , and (ii) values of δ¹⁸O qtz-fs and δ¹⁸O fs-hbl (Δ a - b = δ _a – δ _b) are small and positive (~1 to 2 ‰, Taylor and Epstein 1962). Hydrothermally altered igneous rocks, by comparison, typically lack consistent high-temperature oxygen-isotope equilibrium between coexisting primary phases (Criss and Taylor 1986; Liu 2000). While some samples from the study area have δ¹⁸O mineral1-mineral2 values within the range of magmatic systems, others – most particularly diorite, granodiorite, and syenodiorite – display oxygen isotopic disequilibrium, as manifested by negative values for Δ 18 O qtz-fs and δ¹⁸O fs-hbl . Such reversals are a hallmark of post-crystallization, hydrothermal alteration.

The leucogranite and rhyolite exhibit the least alteration and, except for one sample, have relatively high δ¹⁸O_WR values (≥ +4‰); samples of mafic to intermediate composition, in which mineral phases are more susceptible to alteration, generally have lower δ¹⁸O_WR values (Fig. 3, Table 1a, b). The few samples retaining mineralogical and textural characteristics of phyllic alteration have among the lowest δ¹⁸O_WR (<2‰) and δ¹⁸O_qtz values (<3‰, Fig. 5, Table 1a, b), despite their potential association with magmatic fluids during contact metamorphism. The isotopic results suggest that these samples were overprinted by low- ¹⁸O hydrothermal fluids.

Samples affected by moderate to strong propylitic alteration display the largest range of δ¹⁸O_WR values, with the majority falling between –2 and +4‰ (Fig. 5); primary quartz affected by propylitic alteration likewise has a wide range of δ¹⁸O_qtz values (–2 to +7‰, Table 1a, b). Samples moderately to strongly affected by the later quartz-sericitecalcite alteration have higher δ¹⁸O_WR values (+4 to +10‰, Fig. 5), as does associated vein quartz (δ¹⁸O_qtz = +6 to +11‰) (Table 1a, b). The δ²H_WR values (–93 to –56‰) fall close to or within the normal range for igneous rocks (–85 to –50‰, Sheppard 1986). However, interaction with midto low-latitude meteoric water or seawater can also produce such compositions (Longstaffe 1982). With the exception of hornblende, hydrous minerals in the Huntington Mountain area are of secondary origin (chlorite > epidote > sericite). Where comparisons are possible, the δ²H_chl values are very similar to the δ²H_WR values (Table 1a, b), indicating that chlorite – and fluids that formed it – control the bulk hydrogen-isotope composition of most samples.

The oxygen isotopic results for minerals formed and/ or re-equilibrated during both the propylitic and quartz-sericite-calcite alteration stages have been used to estimate fluid δ¹⁸O values (Fig. 6). Over the 300–450 °C temperature range typical for propylitic alteration (Ferry 1985; Criss and Taylor 1986), δ¹⁸O_H2O values range from –9 to –2‰ at 300 °C, to –5 to +1‰ at 450 °C (Fig. 6a). Coexisting propylitic quartz and chlorite for syenogranite and granodiorite samples (CBI-6-6-55 and CBI-8-5-22) yield virtually identical δ¹⁸O_H2O values of –6 to –5‰ at 300 °C as does coexisting secondary feldspar and chlorite for andesite sample CBI-6-6-52 (Fig. 6a). Hence we use 300 °C as the propylitic alteration temperature in the discussion that follows. These results suggest that meteoric water (δ¹⁸O_H2O < 0‰) comprised a significant fraction of the propylitic alteration fluid. Some estimates based on brachiopod carbonate suggest δ¹⁸O values of –8 to –6‰ for late Precambrian-early Cambrian seawater (Veizer et al. 1986). However, evidence from ophiolite, greenstone, and massive sulphide deposits overwhelmingly suggests open ocean δ 18 O values that varied no more than ±2‰ from its present composition since at least 3 Ga (Muehlenbachs 1986; Johns et al. 2006). Accordingly, we accept a seawater δ¹⁸O value of 0‰ in the discussion that follows.

Quartz-sericite-calcite alteration is generally known to occur at temperatures ranging from 200 to 300 °C (Ferry 1985; Criss and Taylor 1986). For the Huntington Mountain area, calculated δ¹⁸O_H2O values at 200 °C, with two exceptions, are ≤0‰ (range: –4 to +0‰, Fig. 6b). At 300 °C, calculated δ¹⁸O_H2O values are >0‰, with most ranging from 0 to +5‰, Fig. 6b). Vein calcite from volcanic samples CBI-8-5-30 and CBI 6-6-112B yields anomalously high δ¹⁸O_H2O values; they also have the lowest δ¹³C values measured in this study. Oxygen isotopic disequilibrium in these veins is indicated by negative δ¹⁸O qtz-cal values, and in sample CBI-6-6-112B, by the occurrence of calcite as a vug-filling in quartz. We conclude that calcite in these veins formed from fluids unrelated to quartz-sericite-calcite alteration. Two other vein samples from volcanic rocks (CBI 6-6-17 and CBI 6-6-11), however, have textures characteristic of quartz-calcite co-precipitation. Assuming isotopic equilibrium, temperatures of 220 to 230 °C and δ¹⁸O_H2O values of –1 to +1‰ are obtained for these samples using the quartz-calcite oxygen-isotope geothermometer. The apparent trend towards higher δ¹⁸O_H2O values from propylitic (–6‰) to quartz-sericite-calcite (0‰) alteration is suggestive of either ¹⁸O-enrichment of meteoric water during water-rock interaction and/or an increasing contribution of seawater (~0‰) as the system cooled and waned. Combined consideration of the hydrogen- and oxygen-isotope compositions of altered whole-rock samples and associated chlorite allows evaluation of these two possibilities. Figure 7 compares propylitic fluid hydrogen-and oxygen-isotope compositions at 300 °C for chlorite and whole-rock samples to results calculated for a range of molar oxygen water/rock (w/r) ratios in a simple hydrothermal system (following Ohmoto and Rye 1974; see caption to Figure 7 for details). The oxygen- and hydrogen-isotope compositions predicted for meteoric water at ~560 Ma, based on Avalonia paleolatitude estimates (Murphy et al. 2004) closely match the values calculated using the Ohmoto and Rye (1974) model (see caption to Figure 7 for details). Fluid compositions for chlorite range from –5.3‰ δ¹⁸O_H20 and –37‰ δ²H_H2O , which cluster near the Global Meteoric Water Line (GMWL; Craig 1961) at very high w/r ratios, to +3.3‰ δ¹⁸O_H20 and ~ –33‰ δ² H at very low w/r ratios. Many whole-rock samples follow the same trend. This shift away from the GMWL towards higher δ¹⁸O_H2O values at lower w/r ratios is typical of many modern geothermal systems. Meteoric-dominated hydrothermal fluids become enriched in ¹⁸O as w/r ratios decrease whereas δ²H_H2O values are largely unchanged, except at very low w/r ratios, because of the low initial molar hydrogen content of the rock (Craig 1963; Taylor 1974; Longstaffe 1989; Cole 1994). Addition of seawater as the hydrothermal fluid evolved would not produce such a pattern. One sample (CBI-6-6-11), which is dominated by the lower temperature quartz-sericite-calcite alteration, plots just beyond the end of the trend calculated for propylitic alteration at very low w/r ratios (<0.001). This positioning is consistent with a waning flux of hydrothermal fluid.

Seawater may have been important earlier in the alteration history at higher w/r ratios. A handful of samples, including some with sericite as the main hydrous phase, have δ²H_H2O values plotting well above the meteoric-hydrothermal fluid w/r trend-line for 300 °C (Fig. 7). Such compositions could indicate alteration of these samples at >300 °C. Calculated δ²H_H2O values at 450 °C for these samples would plot closer to the trend-line, albeit at higher δ¹⁸O_H2O values and lower w/r ratios. However, the main alteration phases in two of these samples (CBI-6-6-55 and CBI-8-5-22) are in oxygen isotopic equilibrium at 300 °C. An alternate explanation is that the hydrothermal fluid affecting these samples comprised a mixture of seawater and (evolved) meteoric water (Fig. 7). Based on petrographic, isotopic and fluid-inclusion thermometric data, Potter et al. (2012) have suggested that seawater was a significant component of the hydrothermal fluid associated with early vein assemblag formation during the Mira terrane alteration.

The carbon isotopic compositions of disseminated and vein calcite samples (–7.4 to +0.1‰) are not uniquely diagnostic of a particular source (e.g., seawater, magmatic fluids, organic matter oxidation). Further confounding the situation is that Neoproterozoic marine carbonates at 575–550 Ma ranged widely in δ¹³C (–6 to +4‰) (Des Marais 2001, Deines 2002). The majority of samples analyzed here lie within this range (–5.2 to +0.1‰); only those samples with the very highest δ¹⁸O values, and which are unrelated to propylitic or quartz-sericite-calcite alteration, have lower δ¹³C values.

Potter et al. (2008a) proposed two scenarios to explain the regional low- ¹⁸O character of the Mira terrane: (1) a series of local meteoric water-dominated geothermal systems associated with convergent subduction, associated volcanism, and emplacement of individual plutons, or (2) a single event involving regional infiltration of hydrothermal fluids during transcurrent rifting of Avalonia at the Gondwanan margin at ca. 575–550 Ma. Potter et al. (2008a, b) preferred the second model, given the widespread ¹⁸O-depletion of Avalonia, and absence of such ¹⁸O-depletion for the associated inboard Neoproterozoic peri-Gondwanan terranes of the region (e.g., Bras d’Or terrane).

The observations reported here for the Huntington Mountain pluton and East Bay Hills metavolcanic rocks are not typical of a localized hydrothermal system. The putative concentric pattern of progressively cooler alteration zones (potassic, phyllic, propylitic, argillic; Meyer and Hemley 1967; Sheppard et al. 1971; Taylor 1997) that would have been associated with emplacement at ca. 620 Ma is absent – or at the very least – not preserved at present exposure levels. Nor is there a zonation of oxygen isotopic compositions that would be predicted for intrusion-driven water-rock interaction – magmatic fluids near the intrusive centres followed progressively outwards by ¹⁸O-depletion (propylitic alteration) and then ¹⁸O-enrichment (argillic alteration), as meteoric-hydrothermal waters became dominant in the cooling system through circulatory fluid flow above the brittle-ductile boundary as is evident in the distribution of δ¹⁸O_WR values in Figure 4. The minor variations in alteration assemblages across the study area and the absence of any systematic patterning of δ¹⁸O values are more consistent with the regional model for ¹⁸O-depletion across West Avalonia, in which fluid infiltration into deep extensional basins occurred during transition from a subduction to a transtensional rifting environment (Potter et al. 2008a, b). In this model, initial rifting of West Avalonia resulted in the formation of large-scale extensional fault networks, which provided conduits for deep infiltration of meteoric water (and/or seawater during earlier stages). Continued volcanism during the initial stages of rifting served to heat and circulate these fluids. Direct evidence exists throughout West Avalonia for extensional-related volcanism and associated transtensional faulting, including volcanic-sedimentary sequences of the ca. 575–560 Ma Coastal Belt in the Mira terrane, the ca. 560–550 Ma Coldbrook Group of the Caledonia terrane in New Brunswick, and the ca. 580–570 Ma and younger units of the Avalon terrane in Newfoundland (Barr et al. 1996; O’Brien et al. 1996).

Many economic accumulations of base and precious metals are known to form in epithermal systems, commonly as the product of circulation of metal-laden hydrothermal fluids associated with pluton emplacement. To the best of our knowledge, the Huntington Mountain pluton and East Bay Hills Group are barren of significant mineralization (Barr et al. 1996). The hydrothermal fluids responsible for regional ¹⁸O-depletion may not have been metal-rich, perhaps because of the lack of a magmatic fluid component. The δ²H_H2O and δ¹⁸O_H20 values calculated for the Huntington Mountain alteration assemblages certainly do not reveal any strong association with magmatic water (Fig. 7). Alternatively or additionally, it may be that the processes needed to focus the flow of the hydrothermal fluids and/ or precipitate metals efficiently were not operating in this system (Simmons and Brown 2007). Whether regional movement of hydrothermal fluids of meteoric/seawater origin lead to leaching and/or redistribution of pre-existing metals of economic interest in this area, and hence, West Avalonia more widely is worthy of further enquiry.