Regional depth-controlled hydrothermal metamorphism in the Zymoetz River area, British Columbia1

M.G. Mihalynuk and E.D. Ghent

Abstract: An estimated 6 lem of basic to silicic volcanic flows and clastic rocks of the Early Jurassic Telkwa Formation is exposed in mderately east-dipping fault blocks along the Zymoetz River, British Columbia. Extensive wholesale zeolitic replacement of porous tuff beds suggests widespread hydrothermal activity. Metamorphic grade increases regionally from laumontite - albite facies to prehnite - pumpellyite facies with increasing stratigraphic depth. Telkwa Formation strata and the Imparted metamorphic zonation are cut and tilted by rotational block faulting, and are repeated in each of the upturned blocks. Late Mesozoic to Tertiary plutonism locally thermally overprinted the regional fdcies, particularly in the western part of the area. Fluid-inclusion isochores, combined with calculated mineral equilibria, suggest that metamorphism took place at fluid pressures of 2 kbar (1 kbar = 100 MPa) or less, consistent with estimates of stratigraphic burial. Metamorphic fluids were H20 rich and low in dissolved salts. Maximum temperatures during regional depth-sontrolled hydrothermal metamorphism, based upon the widespread presence of laumontite and the lack of wairakite in the middle to upper parts of the Telkwa Formation, probably did not exceed about 250°C at W,O pressures of 2 kbar. Mineral zones, estimated paleotemperatures, and geothermal gradients are comparable to regional hydrothermal metamorphism in active volcanic settings such as Iceland.

RCsurnC : Les coulees volcaniques de composition basique B silicique et les roshes clastiques, de la Formation de Telkwa du Jurassique prkcoce, d'kpaisseur estimee a 6 km. sont exposees dans des blocs de faille mod6riiment inclinks vers l'est, le long de la rivikre de Zymoetz, Colombie-Britannique. Le retnplacement gCnCralisC et 6tendu par des zeolites des lits de tuf poreux tkrmoigne d'une activitk hydrotherinale trks repandue. Le degr@ de metamoqkisme augmente rkgionalement du faciks a laumontite-albite i celui a prkhnite-pumpellyite avec l'accroissement de la profondeur stratigraphique. Les strates de la Formation de Telkwa et la zonation mktamorphique developpee sont recoupCes et bascul6es en blocs par des failles rotationnelles. et elles sont rkpetees dans chacun des blocs rebroussCs. L'activite plutonique entre le MCsozoi'que tardif et le Tertiaire a crCC localement une surimpression thermique des faciks rkgionaux, particulikrement dam la partie occidentale de la region. Les isochores fondCes sur les inclusions fluides, couplkes aux Cquilibres mineralogiques calcules, suggkrent une activitC m6tamorphique sous une pression des fluides de 2 kbar ( 1 kbar = 100 MPa) ou moins. en accord avec les estimations deduites de l'enfouissement stratigraphique. Les fluides mCtamorpkiques Ctaient riches en W28 et contenaient peu de sels dissous. Les temperatures maximales durant la phase d'altkration metamorphique hydrathermale sous enfouissement rkgional, basees sur la presence repandue de launsontite et l'absence de wairakite dans Bes niveaux compris entre le milieu et le sommet de la Formation de Telkwa, n'ont pas excCdC vraisemblablement les valeurs de plus ou moins 250°C sous des pressions de H20 de 2 kbar. Les zonations mineralogiques, les pal6tempCratures estimCes et les gradients gCothermiques ressemblent B ceux observes dans les contextes volcaniques actuels subissant un mCtamorphisme hydrothermal regional, par exemple en Islande. [Traduit par la rCdaction]

introduction important monitor of the tectonic history and metallogeny in the Stikine terrane.

Early to Middle Jurassic Hazelton Group volcanic rocks Stikine terrane is an accumulation sf sedimentary, vsl- cover much of the Stikine terrane, by far the largest, and canic, and plutonic rocks formed in an arc environment. one of the most precious-metal-rich terranes in the Cana- Several pulses of magmatic-arc construction are recorded, dian Cordillera (Fig. la ) . In north-central British Cslum- beginning in Devonian time and culminating in voluminous bia, Early Jurassic Telkwa Formation strata dominate volcanic outpourings in Early Jurassic time that included the the Hazelton Group. Thus, the Telkwa Formation is an Telkwa Formation. Magmatism waned by the early Middle

Received August 24, 1995. Accepted March 22, 1996.

M.G. Mihalynuk2 British Columbia Geological Survey Branch, Ministry of Energy, Mines and Petroleum Resources, Parliament Buildings, Victoria, BC V8V 1x4, Canada. E.D. Ghent. Department of Geology and Geophysics, The University of Calgary, Calgary, AB T2N lN4, Canada.

' British Columbia Geological Survey Contribution 18. Corresponding author (e-mail : MMihaBynuk@galaxy . gov. bc. ca) .

Can. 1. Earth Sci. 33: 1169 - 1184 (1996). Printed in Canada 1 Imprim6 au Canada

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Jurassic, immediately following accretion of Stikine terrane, along with the more inboard Cache Creek and Quesnel terranes (Mihalynuk et al. 1994; Nixon et a1 1993), to the ancestral margin of North America. A large proportion of sediments shed westward from the resulting tectonic welt collected in the foredeep Bowser Basin (Ricketts et al. 1992). A brief, robust period of early Middle Jurassic calc-alkaline volcanism produced outpourings of mainly subaerial, low- potassium tholeiites in a north-trending belt along the north- western Bowser Basin margin. Locus of this magmatic belt shifts radically eastward across the southern Bowser Basin and extends south from the southeastern margin of the basin (L. Biakow , personal communication, 1996). Between latest Jurassic and latest Cretaceous time the Bowser Basin was shortened by folds and thrusts now preserved in the Skeena Fold Belt (Evenchick 199 lb).

Rocks in the Zymoetz River area were first assigned to the Hazelton Group by Hanson (1926). Later reconnaissance geological mapping by Buffell and Souther (1964), Tipper (1976), and Woodsworth et al. (1985) led to refinements in the subdivision and distribution of these rocks. Metamor- phism within the Telkwa Formation in the region has been studied by Freeman (1986) and Dudley (1983). These respec- tive studies focused on authigenic mineralogy and the effects of bulk-rock permeability on their distribution in the Kleanza Creek area (immediately north of the Zymoetz River valley), and characterization of mainly zeolite minerals and the geo- thermal environment that led to their formation in the Smithers area. Within the Zymoetz River area (Fig. lb), hydrated calcium aluminum silicates are widespread and locally spec- tacularly developed as wholesale replacement of tuff beds in the Telkwa Formation. Such occurrences are consistent with a hydrothermal model of regional extent (e.g., Bird et al. 1984), probably produced within the thick pile of slowly cool- ing Telkwa Formation volcanics. Metamorphic grade generally increases with stratigraphic depth. We refer to this as regional depth-controlled hydrothermal metamorphism (RDCMM) .

This study is a characterization of RDCMM in the Telkwa Formation. It is based on two summers of field mapping and sampling in the well-exposed and accessible Zymoetz River area (gravel roads extend up both sides of the river). Subse- quent petrographic, X-ray diffraction (XRD) , X-ray fluores- cence spectroscopy (XRF) , microprobe, and fluid-inclusion investigations (Mihalynuk 1987) were undertaken on selected samples. Of 127 samples collected for their authigenic miner- alogy, only nine were found to contain suitably large fluid inclusions ( > 5 pm diameter) for microthermometric analysis. Generalized geology, microprobe, and fluid-inclusion results and their bearing on phase equilibria and characterization of RDCMM in the Telkwa Formation are presented here.

Geology of the Zymoetz River area

Arguably, the most complete section of Telkwa Formation stratigraphy is exposed in the Zymoetz River area (Figs. lb, lc). Here the true thickness of Telkwa Formation strata can be reconstructed from repetitions within a disrupted, moderately east-dipping homocline exposed for 25 km across strike. This homocline is dissected by the west-Wowing Z ymoetz River.

Stratigraphy The oldest strata in the area are Paleozoic Stikine arc vol- canic~ mantled by calcareous sediments. Most conspicuous are Permian limestones atop which is developed a regionally correlated, 15 - 100 m thick, polymictic conglomerate. Based upon field evidence presented elsewhere (Mihalynuk 1987) the conglomerate is, in part, of Late Triassic age. This con- glomerate grades into the Telkwa Formation lower member tuffite and intercalated lapilli ash tuff of andesitic composi- tion, together about 480 m thick. Approximately 1600 m of overlying, dominantly pyroxene and feldspar-phyric breccia and flow rocks mark the transition to a dominantly subaerial depositional environment. A locally abrupt, but conformable contact with overlying maroon feldspar porphyry tuff marks the base of the 2580-3500 m thick middle member. A diversity of textures and colours and rapid facies changes in these mainly feldspar-phyric pyroclastics and Wows would make correlation from one place to another nearly impossible if not for several widespread marker units. The most impor- tant marker, a couplet of coarse-bladed, crowded feldspar porphyry Wows and tuff and overlying mauve ignimbrite, occurs near the midpoint of the middle member. In grada- tional(?) contact with the middle member is a +600 m thick- ness of mainly maroon-coloured units that form the upper member. From bottom to top, these are well-bedded tuffite, zeolitized lapilli ash tuff, quartz-phyric lapilli ash tuff, partly welded feldspar crystal-rich grey dacite ignimbrite, zeoli- tized lithic ash tuff, and finely banded vitric ash tuff or ash Wow.

Upper member rocks are overlain, probably disconform- ably, by monotonous red pyroclastic lapilli ash tuff of the Red Tuff Member (Tipper 1976; Tipper and Richards 1976) of the Hazelton Group. These are in turn unconformably overlain by nearshore marine strata of the Ashman Forma- tion (Tipper 1976), part of the Bowser Lake Group. Thick- ness of these strata was controlled by the paleogeography of the area in the late Early Jurassic to Cretaceous, at which time the study area was situated on the northern flank of the mainly emergent Stikine Arch, approximately cospatial with the northward-prograding, southern shore of the Bowser Basin (Eisbacher 198 1, 1985; Richards and Jeletsky 1975; Tipper and Richards 1976). Paleo-shoreline position is inter- preted from the shoreline facies displayed within the Bajocian to Oxfordian Ashman Formation. This thin veneer of shore- line strata, less than 1QO m thick, directly overlies the sub- jacent Red Tuff Member (see Fig. 1) without any intervening Lower to Middle Jurassic Smithers Formation (Tsarcian to Bajocian; Tipper 1976), representing a hiatus of more than 7 Ma (based upon the time scale of Harland et al. 1998). Thickness of the Red Tuff Member is variable, in most places about 2QO - 300 m (Tipper and Richards 1976). In the study area, about 5OU m (Fig. I c) may be exposed, represent- ing one of the thickest sections of Red Tuff Member. No evidence of Cretaceous or Tertiary deposition is preserved in the region. Maximum geologically inferred thickness of strata overlying the Telkwa Formation can be attributed to the Ashman Formation and Red Tuff Member, comprising at most 600 m of section. Thus, the development of meta- morphic mineral assemblages was controlled dominantly by the depth at which they formed within the - 6 km thickness of preserved Telkwa Formation strata.

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Fig. 2. Distribution of metamorphic minerals as a function of stratigraphic depth and distance from intrusive contacts. The data set is 98 points based on 127 samples. Biotite, actinolite, and laumontite fields are shaded.

Domain 1 2 Metamorphic minerals RDCHM laumontite 0

purnpellyite prehnite + @ epidote rn uu actindite A A

biotite * i&

garnet

1 4 25 rusive contact (km)

Structure Thrust faults in the western Zymoetz River area interleave

ermian limestone and lower Telkwa Formation strata in domain I . However, in the eastern domains (2 -4), well- bedded Telkwa strata show absolutely no evidence of struc- tural thickening by thrust faulting or folding. The homoclinal Telkwa section is affected only by broad, low-amplitude warping and high-angle faults with kilometre-scale offsets. These normal faults appear to cut a Jurassic pluton (Evenchick 1991a) in the northern map area, but, together with thrust

faults, are cut by Cretaceous('?) to Tertiary granite bodies. Based on analogy with exposed normal fauits having the same north-south strike, these major bounding faults are believed to have normal displacements with a strong dextral component. Conservative offset estimates would range from 2 to -4 km if all of the offset shown in Fig. IIc is due to dip-slip motion.

Three major east-tilted fault blocks of Telkwa Formation repeat the stratigraphic succession. They are referred to as domains 2-4 (Fig. 2). Consistent with this interpretation is

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repetition of metamorphic facies within the fault blocks from near greenschist facies at the upturned bases to laumontite- albite facies at the tops. Clearly, WBCHM predated faulting.

Although direct age data are lacking. RDCHM in the Zymoetz River area is consistent with that of active to waning volcanic belts studied elsewhere (e.g., Bird et al. 1984). Large-scale block faults probably formed as a result of regional extension within this part of Stikinia. This structural style persists into the Whitesail Lake area, 125 km to the south- east (Diakow and Drobe 1989). in marked contrast with the Skeena Fold Belt, where mapped 150 -258 Bun1 alongstrike to the north. Thrust faulting apparently does involve Permian and lowest Telkwa strata in the western Zymoetz River map area (Fig. I), and folds affect Middle and Upper Jurassic sediments to the immediate north and east. The lack of con- tractional structures in the Telkwa stratigra and worthy of further investigation.

ineralogy and regional distribution sf mineral assemblages

The spatial distribution of critical metamorphic minerals with respect to relative stratigraphic depth, distance from intrusive contacts, and the different geologic domains is graphically summarized in Fig. 2. Eaumontite is largely a product of the waxing and waning stages of RDCHM. As such, it occurs as replacement of glassy rock matrix and of primary igneous calc-silicate minerals. These occurrences are limited to the middle and upper members, generally more than 1 km from intrusive contacts (Fig. 2). However, veins s f laumontite can occur almost anywhere in the stratigraphy, reflecting the presence of late-stage hydrothermal fluids restricted to fractures. For example, laumontite vein systems formed late during cooling of the volcanic pile commonly crosscut higher grade RDCHM mineral assemblages. Laumon- tite may also be found replacing feldspars and as crosscutting veins within Cretaceous-Tertiary intrusions, but is mainly restricted to fracture zones within these bodies. Prehnite and epidote occurrences overlap those of laumontite produced by WBCHM, although they are more abundant with increas- ing stratigraphic depth. Actinolite and biotite are limited to even greater stratigraphic depths or closer proximity to intrusive contacts.

Locally, the Telkwa Formation RBCHM is overprinted by greenschist- to amphibolite-facies contact-metamorphic assemblages in the thermal aureoles sf several small plutons and a large easterly extending apophysis of the Coast Plutonic Belt, named the Williams Creek apophysis (Fig. 1 b ) . Contact metamorphism of the Telkwa Formation is most intense adja- cent to the Williams Creek apophysis. The zones of over- printing are typically very narrow, less than 1 - 1.5 km wide. suggesting that the earlier low-grade metamorphism decreased the bulk permeability of these rocks and reduced the extent of the thermal effects of the intrusions (e.g . , Freeman 1986). Contact metamorphism of the pre-Telkwa rocks has locally produced garnet - biotite and wollastonite-bearing mineral assemblages, which can be used to set P - T limits on meta- morphism near the base of the Telkwa Formation.

Mineralogy: WDCHM phases Minerd abbreviations used in the following sections are those of Kretz (1983), except for wairakite (Wa).

Actindite (Act) Actinolite is most abundant within the lower division litholo- gies, primarily as an alteration of pyroxene or in the ground- mass (Fig. 2). It occurs as thin radiating blades in the rock matrix or in optical continuity with host pyroxene. Commonly associated minerals are epidote, chlorite, titanite, prehnite, f pumpellyite, f biotite.

Albite (Ab) Albite commonly replaces plagioclase phenocrysts. Electron microprobe analyses yield compositions of (Table 1). Freeman (1986) reports analyses from partly albltized feld- spars as An3,,- 19.6 in Telkwa Formation rocks, and Dudley (1983) reports compositions of and in unaltered andesite and basalt flows and Ar41,4-6,5 in replace- ment albite. Variability in the degree of albitization is not uncommon in low-grade rocks (Boles 1982).

Culcitu (Cal) Calcite is common, but is rarely present in strongly zeoli- tized strata. It typically occurs as veinlets, or to a lesser degree as a replacement of groundmass feldspar or pyroxene. It is generally one of the last authigenic phases to form, commonly along with quartz. Rare but conspicuous occur- rences as concretion-like masses are found within upper member tuffs.

Chlsr i t~ (ChB) Chlorite is present in all but the most siliceous volcanic units. It commonly occurs as vesicular infillings and groundmass alteration. It locally replaces primary olivine (along with serpentine), pyroxene, biotite, and, to a lesser degree, feld- spars. It also occurs as veinlets and as coatings on fracture surfaces, rarely occurring in masses up to 2 m thick of almost pure chlorite.

Epihte (Ep) Epidote occurs as pleochroic yellow, granular grains and aggregates of grains as an alteration of rock matrix or primary Ca-Al-bearing phases such as plagioclase and pyroxene. In amygdales or veins, it typically has a bladed or acicular habit.

bumontite (Lmt) Laumontite typically occurs as matrix alteration or as alter- ation of plagioclase and pyroxene within porous tuffs. It massively replaces up to 90 awl. % of tuffaceous matrix material. Less intense laumontite alteration extends over strike lengths of 18 km along the ridge east of Clore River (Figs. 1, 2). Original vitric shard outlines are locally pre- served within blocky secondary laumontite overgrowths . Less commcmly. it occurs as amygdales in vesicular flows of the middle and upper members.

Electron microprobe analyses by Dudley (1 983), Freeman (1986), and the present authors indicate that laumontite from several different areas is near end-member composition and can be treated as a pure phase in thermodynamic calculations.

Prehrzit~ (Prh) Prehnite is found in veins and amygdales, and as alteration of plagioclase, pyroxene, and aphanitic groundmass in mafic to intermediate rocks. As void fillings it is coarsely crys- talline and reniform. It is most common in mafic rocks. Associated minerals include quartz - calcite - Cu(meta1) or epidote - pumpellyite -chlorite -quartz -calcite (see also

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Table 1. Representative electron microprobe analyses of minerals frorn the Zymoetz River area, B.C.

Sample: 48-4 Grt 48-4 Bt 101-6 Ep 102-1 Prh 102-5 Pl

Si02 Ti02

A1203

FeO MnO M I P Ca0 Na,O

K20 Sum

Number of atoms Si ~ l i v

Ti Al" Fe Mn MI2 Ca Na K Charge -

Notes: Ali analyses in weight percent. nd, not determined. "Fe as Fe,O,. "Number sf Fe3+ atoms.

Table 2). Microprobe analyses of prehnite indicate strong zoning in Fe-A1, ranging from less than 4 to greater than $ wt. % Fe203 across a 4.2 rnm wide vein (Table 1).

Bumpellyite (Pmp) Pumpellyite is the least abundant authigenic Ca - A1 silicate, and it occurs as small, low-birefringence, green pleochroic blades or felted masses. It is typically a replacement mineral with prehnite or epidote, commonly after plagioclase or pyroxene. With rare exceptions, its occurrence is restricted to the lower member and lower middle members.

Quartz (Qtd Quartz is ubiquitous as either fluid-inclusion-bearing, termi- nated crystals in veins and amygdales, or as polycrystalline to cryptocrystalline groundmass alteration. In some units it is clearly a product of devitrification.

8 t h - minerals Adularia is rare within the study area. It occurs as sharply terminated crystals growing from vein walls, commonly associated with prehnite - epidote - quartz - calcite. Titanite is ubiquitous within the rock matrix as rounded, amorphous masses with difhse boundaries ( < 0.1 mm diameter, 1 - 4%). It is less common as void infillings or as alteration of pyroxene phenocrysts.

Mineralogy: phases specific to contact metamorphism

Biotite (Bt) Biotite commonly forms as a product of greenschist-facies metamorphism of basic volcanic rocks and muddy epiclastics

or tuffaceous pelites adjacent to major intrusions (e.g., Williams Creek apophysis), or more rarely, in fault zones. Microprobe analyses of biotite coexisting with garnet are presented in Table 1. Fine-grained stilpnomelane also occurs locally.

Diopside (Bi) Diopside only occurs near igneous contacts in rnetacarbonates and it forms subsequent grains or clusters of grains in asso- ciation with calcite - tremolite - grossular _+ wollastonite & epidote quartz. Together with minor quartz and fassaite (XRD identification), it also forms 6 cm thick, parallel black bands near the Williams Creek apophysis.

Epidote (Ep) In contact-metamorphic zones adjacent to the Permian carbo- nate, metre-thick layers contain up to 70% epidote. A repre- sentative microprobe analysis of epidote (kL6) is presented in Table 1 .

Gurnet (Grt) Occurrences of garnet are restricted to within approximately 1.4 km of intrusive contacts (Fig. 2). Sparse idioblastic Fe -Mn garnets occur within tuffaceous pelites together with biotite, quartz, and albite (Table 1). Ca-garnet is common in impure Permian carbonates or conglomerate containing car- bonate clasts. Within the sub-Telkwa conglomerates, a 5 m thick stratabound zone is dominated by massive grossular (similar wollastonite- and epidote-rich zones occur nearer to, and farther frorn, the intrusive contact, respectively). Metamorphosed tuf'aceous partings within the carbonates con-

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'Table 2. Representative paragenetic sequences.

Sample Sequence

Permian voleanics 25-2

Basal conglomerate 113

Lower member 115 116 117 136

Middle member 3 6 40- 1 68- 1 75- 1 83- 1 85-2 93 180-2 110 111-2 111-3 133-1

Upper membera 48-2

Act - Ep -Chl (m) -- Ep (v) -- Lmt (v)

Ep-Grt (m) - Qtz -Cal (v) -- Cal (m)

Chl- Act (m) -- Qtz -Cal (v) Chl -Ep (v) -- Ep (m) Ep -Act (v) -- Qtz (v) Ep -Prh (m) -- Prh (v) and (or) Prh-Ep (v) - Prh - Cal

(a) Prh -- Cal / Chl -- Prh-Qtz -- Cal- Ep I Prh -- Qtz (a) Cal -- CRL / Qtz -- Cal (a) Prh -- Chl / Chl -- Cal -- Prh / Chl -- Prh -- Cal (a) Prh -- Cal (a) Ep -+ Ch1 -- Qtz -- Chl / Ep - Ms-Qtz / Ep -- Ep-Qtz-Cal / Ep - Ep-Ms -- Qtz-Cal-Ms (a) Chl -- Qtz / Ep -- Chl / Ep -- Ms I Chl (a) Ep -- Chl -Qtz (a) Chl -- Cal-Qtz Qtz-Cal (v) -- Ep (m) Cal -Prh (v) -- Cal (v) / (a) Prh-Cal -- Chl -- Qtz (a) Chl -- Cal / (a) Ep -- Chl -- Cal / Qtz (v) -- Ep (m) (a) Prh -- Ep -- Prh f Qtz / Chl - Prh-Qtz -- Prh / Prh -- Chl - Qtz / Chl / Prh / Ep

(a) Chl -- Qtz - Cal (v) / Cal and (or) Qtz 1 Lmt

Notes: Arrows point toward succeeding phase(§), and solidi separate sequences. (a), indicates all phases in sequence as amygdale infill; (m), denotes matrix alteration; (v), denotes vein or vug infill.

"Most samples (e.g,, 89-1, 89-3, 91-1, 94-1, 101-1, 105-1, 127-2, 127-3, 128-2, 129, 130, and 131-1) display Lmt only or Lnlt k Qtz k Cal as matrix alteration (m).

- - - -

tain the assemblage epidste - grossular - quartz - calcite f diopside.

Wollastunite (Wo) This mineral occurs in Permian limestone within the contact aureole of the Williams Creek apophysis (Figs. 1, 2). It occurs with diopside - calcite - quartz f epidote & grossular as stratabound compact fibrous masses in impure carbonates.

Paragenesis of minerals Paragenetic sequences within individual hand samples and thin sections (defined on the basis of crosscutting, over- growth, and replacement textures) are presented in Table 2. Some minerals tend to crystallize early (e.g., epidote, actino- lite), whereas others crystallize late (e.g . , quartz, calcite). Laumontite does not generally occur intesgsown with other phases; more typically it occurs alone.

Microthermobarornetry

Quartz and calcite are the primary hosts of inclusions suitable for microthermobarometric analysis. Although petrogenetic observations indicate that calcite and quartz were typically among the last minerds crystallized, they are commonly intergrown with the highest grade RDCHM minerals (i.e.. prehnite. epidote, pumpellyite; 28-2b, 48-4,751, and 101-7; Fig. 3). En such cases the inclusions were probably trapped

near peak RBCHM temperatures (T,,,). Techniques and some details of the observational criteria are presented in the Appendix.

Fluid coanyositiorz The fluid inclusions in the Telkwa rocks are normally two- phase vapour and liquid (V + E) and H20-rich. Salts were determined to be minor but persistent components (about 2.5 wt. % equivalent on average, Table 3). Fluid composi- tions range between 0.9 and 7.4 wt. % equiv. NaC1, except for sample 48-4 (inclusion set 2-I), which, based on tenuous freezing measurements, was calculated as 15 wt. % equiv. NaCl (Table 3). Sample 48-4 is from a quartz-epidote segregation in metavolcanics within the contact aureole s f the Williams Creek apophysis (Fig. 1). Such a high dissolved salts content is not uncommon in a contact-metamorphic environment (Roedder 1984). A separate CU2-rich phase was not directly observed in any sample, but small volumes were indicated in rare inclusions by a darkening of the inclusion at low temperatures ( - 56"C), presumably due to melting of a C0,-rich phase.

Primary versus secondary inclusions Discrimination between primary (p-) , pseudosecondary (ps-) , and secondary (s-) inclusions and recognition of boiling versus necking down were made according to the criteria of Roedder (1984) and Bodnar et al. (1985). Unequivocal evi-

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Fig. 3. Fluid-inclusion isochore ranges from samples representative of the Telkwa Formation lower, middle, and upper menmbers, and contact rnetanmolphisrn of Permian sediments. Cp W28 is the critical point of H,O.

Table 3. Microthemobarometric results.

380 400 508 Temperature ( "C)

Inchsion Host %,I, NaCB TI No. Typea mineral bOc> (wt. % equiv.) ( "c )

Calcite CMclte Calcite Quartz Quartz Quartz Quartz

Quartz Quartz Quartz

Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz

168 -t 10 182.5 +0.2 182.2k0.3 182.4k0.5 175-7kO.5 175.6k0.2 180.Ok0.1 180.250.2 182.3 50.2 181.1 k0.2 176.9k0.2 B88.6f 0.2 179.5 k0.2

Leaked 178.5k0.5

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Table 3 (conchded) .

Inclusion Host 91n NaCl TI No. Typeu mineral ("C) (wt. 76 eeqiv.) ("C)

ph pb sh s sh s

P

s9 ps s9 ps

ph P? so S

P9 Ps P P P P S

s, ps? s, ps?

P P P? P" P P

p?h

P P P P

P P s F

P P P P P

Quartz Quartz Quartz Quartz Quartz Quartz

Quartz

Quartz Quartz Quartz Quartz Quartz Quartz

Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz Quartz

Calcite Calcite Calcite Calcite Calcite

Quartz Quartz Calcite Calcite Calcite Calcite Calcite Calcite Calcite

Leaked 346f 1.5 336f 1.0

269.5 (necked) 329f 0.5 353 k 4

Leaked 152.2k0.2 l65.6k 0.2 174.0k0.2 l89.9kO.l 189.8k0.3

230.4k0.2 229.3 k0.2 225.4f0.2 220.4k0.2 221.4f0.2 107.9k0.5 225.4k0.2 239.2f 0.3

Decrepitated 221.4k0.2 l94.6f 0.2 232.1 f 0.2 239.6kCB.2 223.3 f 0.2

"Inclusions homogenize in the fluid phase. except for sample 88-MM-48-4, inclusion 5-2, which homogenizes in the vapour phase. p, primary; s, secondary; ps, pseudosecondary.

b~nclusion may contain minor CO,. 'a, = 21.6"C. dClath~ate probably present but not visible due to negligible refractive index contrast between

solid and Biquid phases.

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Can. J. Earth Sei. VoI. 33, 1996

dence for p-inclusions is relatively rare and we found no compelling evidence of boiling in any sample. Variable liquid/ vapour ratio (E/V) may in general be attributed to necking of fluid inclusions. Good evidence of necking exists in contact- metamorphosed sample 48-4, where secondary, variable &/V inclusions within the same growth plane yield widely dis- parate homogenization temperatures (T,) of 269.5 - 387 "C . Although few conclusive examples of necked inclusion pairs exist in this study, it is likely that many secondary, variable L/V inclusions formed in this manner.

Samples displaying only one pulse of mineral growth with minor subsequent fracturing (i.e., matrix alteration or vesicle infilling) typically yield very narrow p-inclusion Th ranges (e.g., 75-1, 101-7, 28-2b, 111-2, 113-2b). In general, Th does not vary a great deal within individual samples, even within different host minerals (e.g., 75-1). Sparse and com- monly tiny inclusions in 48-4 yielded a range of Th consis- tent with primary inclusions trapped during intrusion of the adjacent pluton. The very highest Th vdues are most likely to be representative of fluids during peak contact metamorphism at that locality.

Th and composition of secondary inclusions may vary greatly from those trapped during RBCHM. ]Inclusions from highly fractured rocks, for example, 95 and 132-5, show a fairly large range of primary inclusion Th as well as a great disparity between the Th of p- and s-inclusions due to multiple fracture and fill cycles. A good example is 11 1-2, in which primary H20-dominated inclusions have a mean 'F, of 226.4 "C with standard deviation of 6.2 "C, whereas a seeon- dary inclusion may have a Th = 108OC. Although averaged Th vdues have little geological significance, they do demon- strate that, in general, temperatures decreased after entrap- ment of primary inclusions. With few exceptions, the average Th of primary inclusions generally exceeds the average Th of secondary inclusions (Table 3). Two exceptions (75- 1, 1 1 1-2) display Th of s-inclusions that is roughly equivalent to Ti, of p-inclusions, probably indicating thermal stability over the period of time between the two trapping events. To deter- mine whether the general trend to lower Th values in s-inclusions is largely a function of the trapping temperature (TJ, it is necessary to consider total pressure (P,).

Pressure of entrapment Interpretation of the P- T significance of homogenization temperatures (Th) requires an estimate of either the pressure or the trapping temperature (T,) by an independent method. Typically, pressure is estimated from stratigraphic depth, but the relationship between the lithostatic pressure due to the rock load (P,) and the hydrostatic pressure (Ph) must also be specified.

A limit may be placed on the minimum depth of inclusion trapping through the lack sf evidence for boiling in d l samples studied. This suggests Ph >40 bar (1 bar = 100 kPa) at 208 - 250 "C using observed fluid compositions (Haas B 97 1 ) , which strongly suggests that, if the alteration took place under hydrostatic conditions, the depth would be greater than about 450 m. The relative values of lithostatic pressure (P,) and fluid pressure (Pf) are difficult to estimate in volcanic terranes because of numerous fractures and other discontinui- ties in permeability that permit Pf to vary between hydrostatic and lithostatic values over a distance of a few centimetres or less. Estimates of permeability in hydrothermally metamor-

phosed rocks are difficult to quantify (e.g., Manning and Bird 1995). In rocks that have undergone episodes of successive fracturing and resealing of fractures, averaged permeabilities have little significance.

The relationship between Pf, Ph, and P, has been the subject of numerous papers (e.g., Wood and UPalther 1986). Fluid pressures measured in boreholes at geothermal fields (e.g., Wairakei, Valles, Caldera, Yellowstone) indicate near or slightly less than hydrostatic conditions (e.g., Goff et al. 1985). These estimates probably do not apply to hydrothermal alteration of the Telkwa Formation, since these drill holes are targeted at the most productive and, therefore, the most permeable zones. In addition, modern geothermal fields are anomalously hot, near-surface upwellings located within other- wise geothermally unspectacular basins or volcanic fields.

Zones having permeabilities (k) of 10-I" -10-l7 down to depths of 1.7-3.3 in crystalline rocks are documented in a survey of 27 sites by Brace (B 980). In situ measured values of permeability in crystalline rocks vary by a factor of 105; laboratory measurements provide only a minimum estimate of k (Brace 1980). Henley and Ellis (1983) noted that one crack (8.04 mrn thick per cubic metre of rock can account for permeabilities of 18-14- m2. To a large extent this explains why laumontite produced by RBCHM is restricted, whereas late laumontite veins are widespread.

Inclusion densities Consistent inclusion densities in most samples (Table 3; Fig. 3) suggest relatively stable Pf during RDCHM of the cooling volcanic pile. Th measurements indicate fluid denslties aver- aging 0.89 g/cm3 assuming an average rock density of 2600 kg/m? One sample (132-5) shows a range of densities, but this sample is from a fault zone with multiple phases of calcite veining. Estimation of pressures over the range of isochores at fixed temperature in this sample nearly covers the range of Pf = P, to Pf = Ph, which is reasonable for a repeatedly fractured and sealed zone.

Evidence for a regime of Pf approaching P, at the time of inclusion trapping is suggested by the uniformity of the iso- chores from both veins and amygdales (45-1). Amygdales approximate small "closed" systems within the flow matrix (they are not connected by veins) and should provide a monitor on relative P - T changes in nearby veins. Pressures within the veins and within the amygdales of 75-1 were not significantly different. Since fluid pressures vary from poten- tially minimum values in fissures to maximum values in the rock matrix, the isochores from 75-1 would suggest that the Pf in veins is equal or nearly equal to that of the matrix (prehnite is also found in 75-1 as alteration of matrix feld- spars). Zoning in vein minerals, for example, Fe3+ - A1 in prehnite described above, could be due ts variation in several intensive variables, like P , T, activity of H20 or the hgacity of oxygen (9;,,). The occurrence of high-variance assemblages does not permit such variables to be constrained.

Based upon the fluid-inclusion measurements, excluding the contact aureole (48-4), the maximum Pf (= P,) at 250°C is about 2 kbar (Fig. 3).

Phase equilibria in the system NCMASW: application to the Zymsetz River area

The system Na20-CaO-MgO-A1203-Si02-I420 (NCMASH) has been the subject of numerous experimental, calorimetric.

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Fig. 4. Pressure-ternperamre dlagram for important phases in the Na,O - @a0 - AB,O, - SiO, - H,O (NCASH) system calculated using $TAX. The activity of H,O (a,;,) = 1.0.

iws + Qtz +

Stb = Lmt + Qtz

[41 -' Wui =

Lrnt + Qtz + H20

[91- Anl + Qtz =

Ab + H2C

Grs + Czo + Qtz + H2Q

and theoretical studies (Berman 1988; de Capitani and Liou 1996). It should be emphasized that an equilibrium model may not strictly apply to all of these low-grade metamorphic rocks. Reactions producing hydrated Ca -A1 silicates from igneous calcic plagioclase and volcanic glass are not revers- ible. At low temperatures the probability of metastable crys- tallization and persistence of metastable phases are significant. In addition, low-grade metamorphism is characterized by the lack of reaction assemblages, which allows one to set only broad limits on PHz0 - P, - T.

The PHzO - Ph - T stability limits of laumontite have been calculated using the thermodynamic data from de Capitani and Liou (1996). Computed equilibria in the Na20 - CaO - A120p - SiQ2 - H20 (NCASH) system are plotted in Fig. 4. The - Ps-T stability field of laumontite is limited by the following equilibria:

[2] Lmt - Lws + 2Qtz + 2W20

[3] Stb = Lmt + 3Qtz + 3H20

141 Hul = Lmt + 3Qtz + 2H20

Capitalized abbreviations refer to phases and lower case abbreviations refer to phase cornponerm (e .g . , Prh vs. prh). Based u p n equilibrium [I], the maximum temperature at which laumontite is stable is about 235°C at 1 kbar and 270°C at 2 kbar with = Ps (Fig. 4). Based upon equi- librium [23, the m i m u m pressure at which laumontite is stable is about 3 kbar for PwJo = P,. Based upon equilibria

[3] and [4), the minimum temperature would be about 185 "C at 2 kbar, and the minimum pressure would be about 2 kbar at 160°C and about 1 kbar at 22S0C, all with PHzo = Psa For the case of hydrostatic conditions, PHlO = I/p P,. equi- librium [I] lies at 155°C at P, = 2 kbar and 130°C at P, = 1 kbar, so laumontite would not be stable above these temperatures. Equilibrium / 3 ] lies at 76°C at P, = 2 kbar and 80°C at P, .= 1 kbar, setting a lower temperature limit for the stability of laumontite. There is no evidence that laumontite crystallized from either heulandite or stilbite, so estimates based upon these equilibria should be viewed with caution.

Laumontite is common as a RBCHM phase in the upper parts of the middle member ad throughout the upper member, and its distribution is limited by the distance to intrusive contacts (Fig. 2). Laumontite occurs within 1 km of the Late Cretaceous dioritic stock in domain 3. 'This dual control on the laumontite distribution is shown by Fig. 2. Subhorizontal boundaries in the mineral distribution diagram suggest that the mineral stability is larlgely a function of depth within the stratigraphic pile, whereas vertical boundaries suggest control by the thermal influence of intrusive bodies.

The P-T stability of prehnite has been the subject of discussion for several years (Liou 1971b; Helgeson et al, 1978: Kerrick and Ghent 1979; Chatterjee et al. 1984). The calculated equilibrium

lies at 397 " C at 4 kbar and 444" C at 1 kbar (Fig. 4). If zoisite is used instead s f clinozoisite, the equilibrium lies at 3'71 "C at 4 kbar and 420°C at 1 kbar. If this equilibrium is applied to natural occurrences, it is necessary to account for solution of the phase components prh, grs, and czo. In addition, the U H ~ O may not be equal to 1.0. For the case of hydrostatic conditions, equilibrium 151 lies at 216°C at P, = 4 kbar and 240°C at P, = 1 kbar. Consequently, equilibrium [5j does not provide a realistic picture of the P- T stability of prehnite in metamorphic rocks. In the study area, where volcanics are not affected by the thermal metamorphic aureole of younger intrusives, prehnite persists down to the Permian carbonates. Since garnet is lacking in these racks, the abundance of actinolite (Act) suggests that prehnite may have reacted with chlorite and quartz to produce actinolite and epidote + pumpellyite. Consider the equilibrium

[6] 2Lnst -b Prh = 2Czo +- SQtz + $H20

Laumontite is essentially end-member composition. Using average mineral analyses, prehnite has XA, -- Qa72[Al/(A1 + Fe3+)], and epidote has XAl = O.'is2[Al/(Al + Fe3 +)I. We calculated activities for these phase components using ideal- solution models for prehnite (a,, = (X,&(XAI) and for clinozoisite (a,, = 1 - 3xFe) (Digel and Ghent 1994). Using these phase component activities, equilibrium 161 lies at about 229°C at 4 kbar and 24Q°C at 2 kbar.

The thermodynamic properties of Mg-purnpellyite have been estimated by a number of researchers (e.g . , Powell et al. 1993). The equilibrium

[7] Mg-Bmp + Qtz = Czo -k Chl + Prh -k N20

in the CaO - MgO - A1203 - Si02 - H30 (CMASH) system is near 250°C at 2 kbar, and the equilibrium

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1 180 Can. J. Earth Sci. VoB. 33, 1996

Fig. 5. Geothermal gradient limits based on the intersection s f P,, and P, pressure estimates from stratigraphic depth with fluid-inclusion isochores.

82- 1 9% es

111-2 11 3-2b

Found within sample O in over- or underlying units

Upper P Limit

minimum geothermal gradient approximately 85"@/km

Lower P Limit

maximum geothermal gradient approximately 16QWkrn

200 380

TEMPERATURE (T)

181 Brh + Chl + Qtz = Czo + Act + W 2 0

in the CMASH system is near 345°C at 2 kbar (Powell et al. 1993). - - - - - According to Powell et al. prehnite + chlorite is not stable d a t i v e tFpumpeHyite-+- e p a e - + - a ~ t i n o l i t e + H2D at pressures greater than about 4 kbar for PHl0 = Ps. We have not documented the stable coexistence of pumpellyite + epidote + actinolite in the Zymoetz River area.

Albite + quartz were apparently stable relative to anal- cime in the Zymoetz River area. The equilibrium

191 An1 + Qtz = Alb 4- M 2 0

lies at about 200°C at 1 kbar (Liou 197 la). We used the heat capacity and entropy of analcime calculated by Berman and Brown (1985) and adjusted the enthalpy of analcime in equilibrium [9] to force it through 205°C at 1 kbar (m?(298) analcime = - 3 307 050 Jlmol) . The minimum temperature at which albite + quartz is stable is ktween about 200°C for PHZ0 = Ps = 1 kbar and about 80°C for PHs = 1/3 PS = I kbar, that is, under hydrostatic conditions (Fig. 4).

Comparison of phase equilibria in NCMASH with nuid-inclusion results

Fluid-inclusion isochores are generally, but not always, con- sistent with the computed phase equilibria of the meta- morphic minerals in the NCMASH system associated with the quartz or calcite fluid-inclusion hosts.

Samples 28-2b, 75-1, and 101-4 all contain prehnite in apparent textural equilibrium with quartz hosting fluid inclu- sions. The temperature range from the isochores is 1.5 kbar is 280 - 345 "C (Fig. 5). This temperature range is compatible

with observations made of the temperature of formation of prehnite in hydrothermal systems; the lower temperature (280°C, sample 28-2b) is below that observed in geothermal systems such as Cerro Paieto and Wairakei (Bird et al. 1984; W e y andXl_lis_ 19&3)-. Boles and Coombs (1 977), however, calculate that the Iower cf$afliiation temperature of preh-- nite in the Taringattmra Mills burial metamorphic sequence in New Zealand is approximately 98°C.

Microthermometry s f some samples yields temperatures that are higher than those computed from phase equilibria of laumontite (e.g., 345 vs. 270°C at 1.5 kbar, sample 75-1). Samples immediately overlying and underlying 101-7 and 95-1 contain WDCHM-formed laumontite, yet both sanlples fall outside the laumontite stability field (Fig. 4). Clearly, use of the stability field for laumontite calculated for hydro- static conditions yields an even larger disagreement. A low P - T extension of sample 10 1-7 isochores (Fig. 3) intersects the laumontite stability field, but the low P-T extension of 75-1 does not. Laumontite may have grown during the waning stages of metamorphism from fluids with a tcmpera- ture lower than that of the trapped fluid inclusions in quartz and calcite of samples 101-7 and 75-1. Growth of laumontite rather than prehnite in some rocks was possibly controlled by fluid composition and not P - T, for example, by the aCaL, /

ratio of the fluid species (Fig. 6) , but this is not likely the case for sample 95- I . In sample 75- I , prehnite exists within veins and amygdales (commonly connected by vein- lets) and as alteration of groundmass feldspars, and the fluid inclusions suggest relatively high temperatures. Fluid- inclusion and phase-equilibria constraints imposed by 75-1, and adjacent Iithologies containing laurnontite, might be satisfied by very narrow thermal inversion zones. Thermal

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Fig. 6. Topology sf the aqueous laurnontite-p~hnite equilibrium.

1 Calculated at

Qtz

Activity of aqueous SiO,

reversals with depth in geothermal systems are well docu- mented in the Cerro Prieto field (Bird et al. 1984). Such zones are < 500 - 1000 m thick, with temperature gradients of >200°C over a 500 m interval (see Fig. 13A in Bird et al. 1984).

Fluid chemistry and bulk-rock composition controls Zen ( 196 1) suggested that Ca-zeolites such as laumontite would be unstable relative to clay-carbonate assemblages in contact with aqueous fluids containing small amounts of C02, and numerous authors have subsequently applied this concept to low-grade metamorphism (e .g . , Thompson 197 1). Except for the crystallization of late-stage calcite, which was probably not in equilibrium with earlier formed hydrous Ca- A1 silicates, there is little evidence for appreciable C Q in the Zymoetz River metamorphic fluids. Fluid-inclusion studies support the interpretation that the fluids contained little C02. Another equilibrium that indicates fluid compo- sition control over the relative stabilities of prehnite and laumontite is

A cation versus as,, plot (Fig. 6) indicates that not only is prehnite favoured by? high aCai+, but also it is stable with respect to laumontite at lower values of aSio, Alteration of samples 75-1 and 101-7 and overlying and underlying strata may represent the combined effects of (i) stratabound, hydro-

thermal fluids not in thermal equilibrium wlth the surround- ing rocks: (ii) bulk-rock compositions most compatible with prehnitization; and (iii) reduced permeability resulting in the preservation of early formed laumontite.

P - T - X relations of wollastonite and garnet - biotite gesthermometry

Garnet and biotite occur within pelitic partings and wol- lastonite within the carbonate of contact-metamorphosed Permian strata immediately underlying the Telkwa Forma- tion. They provide a pressure estimate independent of that based on the reconstructed stratigraphic thickness, thereby helping to confirm the pressure limits of RDCHM in Telkwa Formation strata.

Wollastonite in the contact aureole was probably formed by the reaction

I l l ] Cal + Qtz = Wo + C 0 2

We estimated fluid composition using fluid inclusions associated with wollastonite in sample 17. Fluid inclusions are typically < 10 pm, two-phase, and liquid dominated. In no case was a third phase observed. A thin shell of C 0 2 around a water vapour bubble could easily have been missed in inclusions this small; however, crushing results confirmed our petrograhpic observations. The vapour bubbles of decrepi- tating inclusions did not expand, but shrank and disappeared within 5 - 10 s. Such inclusions cannot strictly represent the fluid species at the time of wollastonite formation, since some C 0 2 must have been present, but they indicate the water-rich nature of the fluid phase. For an estimated fluid composition of Xco9 = 0.0 1 -0.05, calculated minimum temperatures of wollastonite stability at 2 kbar would be 450 - 550°C using the database of Berman (1988) and the programs of Brown et al. (1988).

Pelitic partings within the rocks beneath the Permian carbonates contain garnet-biotite. This mineral pair can be used as a geothermometer through application of the exchange equilibrium

[12] Phl + Alm = Ann + Pp

e.g., Ferry and Spear (1978). In this computation we made the approximation that all Fe is ferrous. Microprobe analyses of garnet-biotite pairs from below the Permian carbonates are summarized in Table 2, where garnets have a very high spessartine component, up to 34 mol%. Using the Berman (1988) database with additional data for almandine and annite from Berman (1998) and the solution model of McMullin et al. (1991) for biotite, we calculated an equilibrium temper- ature of 540°C at 2 kbar. This result is in reasonable agree- ment with the stability limits of wollastonite discussed above.

Fluid-inclusion densities in quartz-epidote segregations were determined from Th in a sample collected at a location (48-4) near the garnet -biotite sample described above. A P - T plot of constant distribution coefficient KD for Fe - Mg exchange in equation 11 21 determined from garnet -biotite analyses intersects the isochore at pressures slightly less than 2 kbar (PHZ0 = Ps) at 540°C. This is consistent with pres- sures estimated from stratigraphic thickness and points to a more or less complete Telkwa Formation stratigraphy at the time of pluton emplacement.

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Estimation of geothermal gradients and comparison with modern geothermal systems

Geothermal gradients may be estimated from (i) laumontite stability within the stratigraphic section (Fig. 2); and (ii) the range of fluid-inclusion isochores (Fig. 3) with respect to stratigraphic depth, assuming inclusion entrapment occurred at T,,,. For the first case, laumontite distribution is a func- tion of both depth and distance from intrusive contacts. The presence of laumontite indicates P - T conditions within the laumontite stability field, less than about 240°C at 1 kbar and 255°C at 1.5 kbar (Fig. 4). Sample 106, at the intersection of the two 'laumontite out" curves in Fig. 2, clearly marks the deepest occurrence of laumontite formed by RBCWM with domain 3 stratigraphy. An estimated thickness of strata removed from atop sample 106 is 4 -5 km. Using the maxi- mum temperatures at which laumontite is stable yields a geothermal gradient in the range of 50-60°C/km. If the stratigraphically lowest occurrence of laumontite in domain 3 (Fig. 2) is at the limits of P-T stability for laumontite, this would indicate a geothermal gradient of approximately $SQC/km. This occurrence may actually fall well within the P-T stability limit of laumontite, leading to a 8 5 " C / h geothermal gradient estimate that is too high. If stratigraphic thickness, amount of overlying strata, and relative posltion of each sample are approximately correct, the geothermal gradients can also be estimated from the range of tempera- tures over which isochores fall between P = P, and P = Ph. Such an analysis of the middle and upper members is shown in Fig. 5. The upper and lower limits of the geothermal gradient range, as calculated by this method, are 160 and 85"C/km, respectively (the range of curves that pass through all ofthe intepekdisochore PyT km&s of Fig. %).The lower limit of this range of gradients (85"CIh) encompasses that estimated from laumontite breakdown in domain 3.

Studies of geothermal systems have documented depth- controlled hydrothermal alteration zones, e . g . , Eiou et al. (1 987). For example, near Reydarfjordur , eastern Iceland, studies of a 3.5 km crustal seetion indicate a low-temperature zeolite zone down to 1.6 km with a maximum paleotempera- ture of 100- 120°C. It is succeeded downward by a high- temperature zeolite zone to a depth of 2.8 km. This zone is dominated by laumontite, chlorite, calcite, and quartz with minor epidote and prehnite in its lower part, in which the maximum tempemture was 230°C. From 2.8 to 3.5 km depth, epidote, quartz, prehnite, chlorite, albite, and quartz are abundant and minor amounts of wairakite, pumpellyite, and actinolite are present. The temperature of this deeper zone is estimated to be less than 300°C. Locally, contact metamorphism superimposed on the preexisting assemblages has produced wollastonite and andradite. Using an estimated geothermal gradient of $8 -90"C/km, the onset of the green- schist facies would occur at about 4 km depth and the amphib- olite facies at about 6 h crustal depth (Liou et al. 1987, pp. 99 - 101).

We do not imply that the geologic setting of the Zymoetz River rocks is the same as that in Iceland, but regional depth- controlled hydrothermal metamorphism has produced similar successions of metamorphic minerals in both areas. Paleo- geothermal gradient estimates for the Zymoetz fiver meta- morphism are not constrained by stratigraphy and mineral

equilibria alone. In this study from fluid-inclusion data.

Acknowledgments

Critical reviews by FLE.

Can. 4. Earth Sci. V d . 33, 1996

we provide additional constraints

Beiersdorfer and R . Springer - -

improved the final manuscript. Their contribution is appreci- ated. Funding for field studies was largely supported by Natural Sciences and Engineering Research Council of Canada grants to E.B. Ghent (1985). Logistical field support was provided by the Geological Survey of Canada (1984), and financial assistance was provided through teaching and research assistantships from the University of Calgary and the Galloway Memorial scholarship.

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Berman, R.G. 1990. Mixing propeorties of Ca- Mg- Fe - Mn garnets. American Mineralogist, 75: 328 - 344.

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Chatterjee, N.J., Johannes, W., and Leistner, #. 1984. The system CaO - Al,O, - Si0, -H,O: new phase equilibria data, some calculated phase relations, and their petrological applications. Contributions to Mineralogy and Petrology, 88: 1 - 83.

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Evenchick, C.A. 1991b. Geometry, evolution, and tectonic frame- work of the Skeena Fold Belt, north-central British Col~~nbiii . Tectonics, 10: 527 - 546.

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Freeman, K. J. 1986. Hydrothermal metanlorphisnn of Telkwa Formation volcanics near Terrace, British Columbia. M. Sc. thesis, the University of Calgary, Calgary.

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Microthermometric study of fluid inclusions provides the potential P - T conditions of their entrapment. and also potentially yields information about the composition of the fluid from which the host mineral formed (Woedder 1984). Host minerals in this study are limited to either quartz or calcite (Table 3), which are present in virtually all mineral assemblages.

Techniques The fluid-inclusion data summarized in Table 3 were collected using a United States Geological Survey-style heating-cooling stage (Werre et al. 1979; Roedder 1984, pp. 196 - 198) mounted on a Eeitz monocular microscope equipped with long free-working- distance lenses and 25 x ocular, Tenlperatures were monitored with a chrome1 - corntantan thermocouple having a range of 200 - 1067 "C and a resolution of 0. 10°C. Induced thermocouple voltages were digitized by a Doric 410A digital indicator with 0. 10°C resolution. The stage was calibrated at zero point using both synthetic H,O inclusions and doubly distilled water wlthin a capillary tube. A heating rate of < 0. BO0C/min on approach to the melting temperature 4%) was sufficiently slow once the stage was thermally stabiHi~ed (approx. 1 h at near 0°C) to yield melting points of approximately -0.1 and 0.1 "C (&O. I "C), respectively. Moderately high tcmper- ature calibrations using potassium dichrornate and sodium nitrate were also made, dcspite instmmcnt specifications claiming maxi- mum linearity errors of only +(4.4"C. a;, and Th measurcnaents of unknowns were made with heating rates of < 0.2 "Cimin (if possible) when in the vicinity of the Tof interest. Commonly, a more dynamic heagngm-n was ~ d k e to guide - - - bter, - more precise measurements.

- - - -

A eutectic temperature (T,) was rarely determined due%3hFw3tE-

rich. NaCl-poor nature of the bulk of the inclusions studied. This composition range results in a very shallow eintectic curve, and minute volumes of liquid forming at the eutectic are very difficult to detect (Crawford 1% 1). Isochcsres were calculated from micro- thermometric data using the FORTRAN programs WATER and HALWAT of Nicholls and Crawford (1985). WATER generates two sets of isobaric data points from (i) the steam tables equatims (Keenan et al. 1969) fit to empirical data below 1 kbar and extrapa- lated to higher pressures, and ( i i ) the Redlich -Kwong equation OC state (Bolloway 198 1 ) . Following the advice of Nicholls and Crawford (19#5), the steam table values were used due to the relatively low pressure naturc of the metamorphism of' the Telkwa Formation rocks. HALWAT calculates isochores from the data of Potter and Brown (1977) using least squarcs intcrpolatioam below 2 kbar and linear extrapolation algorithms above 2 kbar. The resuhnt isochores are plotted in Fig. 3.

Sources of analytical error The uncertainties listed in Table 3 are subjective estimates of error. These estimates are based on the following. listed in order of importance: (i) the optical quality of thc inclusion, ( ibb the heating rate, and (iik) undetected compcsnents . Often, but not a l w q s . the optical quality, having a direct relationship with inclusion size, also varies with inclusion shape, orientation (especially in calcite), interference from other extraneous inclusions, host mineral(s), and quality of sample polish. Although some authors advocate heating rates of 0.4"Cinain (Macdonald and Spooner 1981), the method of Roedder (1984) and Crawk'ord (198 1) using a maximum heating rate of 0.2"Cimin (slower is better) was found to producc the best results (see Techniques). Due to gas tlow tluctuatiesns, however, such slow heating rates were at times unattainable during ;$, heat- ing runs. The presence of undctested components within the fluid inclusion (for example, small volumes of C 0 2 or pcrhaps CaCB,)

WOUH -dkct-thc ingltcrpmt&icrn ef ~ ~ - 4 1 & a ~ ~ i w e d . - - - - -

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