Introduction
Pegmatites are a fascinating field of study due to their commonly exotic mineral assemblages, gigantic crystal sizes, unique textures and significant economic potential (Linnen et al., Reference Linnen, Van Lichtervelde and Černý2012). They are a major source of many critical elements, such as Li, Cs and Ta and are one of the main sources of world-class gems (London and Kontak, Reference London and Kontak2012, McCaffrey and Jowitt, Reference McCaffrey and Jowitt2023).
The Prof pegmatite is located on Boulder Mountain ∼7 km northwest of Revelstoke in southeastern British Columbia, Canada (Fig. 1). An assessment report by Lane (Reference Lane2017) noted many rare minerals in the pegmatite including elbaite [Na(Li1.5Al1.5)Al6(Si6O18)(BO3)3(OH)3(OH)] and lepidolite (a mineralogical series between trilithionite [KLi1.5Al1.5(Si3AlO10)F2] and polylithionite [KLi2AlSi4O10F2]). The Prof pegmatite is part of a field that hosts other lithium-bearing pegmatites, which includes Grail, Red and Green dykes, and numerous barren pegmatite intrusions (Lane, Reference Lane2017). Regardless of these known occurrences, the extent of the field is unknown. The geochemistry and mineralogy of the neighbouring Mount Begbie pegmatites, 15 km south of Boulder Mountain, has been described by Dixon et al. (Reference Dixon, Cempírek and Groat2014). Mount Begbie hosts an extensive pegmatite group with an abundance of the rare minerals elbaite and beryl, however, these pegmatites are not known to host significant amounts of Li mineralisation.
Regional geology of the Revelstoke area modified after Wheeler and McFeely (Reference Wheeler and McFeely1991). The red square shows the location of Boulder Mountain and the Prof pegmatite. The overview map shows location of Monashee Complex in relation to geological belts of British Columbia (BC), Canada.
Although other pegmatites of the region have been investigated from a regional tectonic perspective, exotic minerals such as those containing lithium have not been noted (e.g. Johnston et al., Reference Johnston, Williams, Brown, Crowley and Carr2000, Kruse, Reference Kruse2007). Johnston et al. (Reference Johnston, Williams, Brown, Crowley and Carr2000) documented the occurrence of tourmaline-bearing simple pegmatites which crystallised contemporaneously with a late deformation event in the Blanket Mountain area (located south of Mount Begbie) at 50.2 ± 0.5 Ma using U–Pb radiometric age determination of monazite. Kruse (Reference Kruse2007) investigated the relationships between multiple pegmatites and the implications of this for the structural evolution of the Thor-Odin Cumulation.
The growing importance of Li in battery technologies makes understanding Li ore-forming processes important for locating future resources. This investigation documents the first published data on the Prof pegmatite occurrence which includes the recent discovery of abundant petalite mineralisation together with detailed mineralogical assemblage descriptions, and mineral compositional evolution analyses during crystallisation.
Geological background
The Shuswap Metamorphic Complex is an extensive exposure of exhumed, former mid-crustal rocks that extends from Washington State in the USA to Quesnel Lake in British Columbia (BC; Okulitch Reference Okulitch1984). The complex is a part of the Omineca Belt of the Canadian Cordillera (Fig. 1; Wheeler and McFeely, Reference Wheeler and McFeely1991, Kruse et al., Reference Kruse and Williams2005). Boulder Mountain is located northwest of Revelstoke, BC, within the Monashee Complex (a division of the larger Shuswap Complex), which comprises two major structural culminations: the Frenchman Cap Dome in the north and the Thor-Odin Dome in the south (Fig. 1). Boulder Mountain is located between these cumulations within the Monashee cover sequence (Norlander et al., Reference Norlander, Whitney, Teyssier and Vanderhaeghe2002).
The Monashee complex comprises varied lithologies including ortho- and paragneisses and migmatites, overlain by cover sequence marble, schists, quartzite, calc-silicate and quartzofeldspathic gneisses (Hinchey et al., Reference Hinchey, Carr, McNeill and Rayner2006). The Monashee basement rocks of the region underwent high-grade metamorphism in the late Palaeocene-early Eocene and have experienced related deformation and anatexis as shown by large leucosome accumulations (Hinchey et al., Reference Hinchey, Carr, McNeill and Rayner2006).
The Monashee complex underwent a period of rapid exhumation and decompression between 60 and 50 Ma, which led to the current domal exposures (Spalla et al., Reference Spalla, Zanoni, Williams and Gosso2011). At the conclusion of the principal exhumation stage in the Thor-Odin culmination, before final exhumation, the temperature and pressure were ∼700–800°C and 9–10 kbar (Norlander et al., Reference Norlander, Whitney, Teyssier and Vanderhaeghe2002, Goergen and Whitney Reference Goergen and Whitney2012). These conditions had decreased to ∼600°C and 2.5–5 kbar in the Eocene. Later-stage Eocene faulting created normal brittle faults that cut some of the pegmatites of the region (Kruse and Williams, Reference Kruse and Williams2005, Hinchey et al., Reference Hinchey, Carr, McNeill and Rayner2006).
There is abundant evidence in the geological record in the form of migmatites, leucosomes, and simple pegmatites, that anatexis, as a direct result of decompression, was a major process in the 60–50 Ma time period (Hinchey, Reference Hinchey2005). Johnston et al. (Reference Johnston, Williams, Brown, Crowley and Carr2000) showed, with zircon, monazite and titanite U–Pb radiometric geochronology, that aplite-pegmatites associated with this event in the Blanket Mountain and Grizzly Flats areas (15 km south of Mount Begbie) crystallised between 50.2 ± 0.5 Ma and 52.2 ± 0.5 Ma. In addition, the youngest S-type granite of the region is the Ladybird suite (Hinchey and Carr, Reference Hinchey and Carr2006), which is considered to have formed by anatexis in the early Tertiary period with a U–Pb radiometric zircon age of 62.1 ± 0.3 Ma for the granites and as young as 55.5 ± 0.3 Ma for associated pegmatites (Carr, Reference Carr1992).
The Boulder Mountain Pegmatite Group
The Boulder Mountain Pegmatite Group is composed of the Prof, Red, Green and Grail pegmatites. These pegmatites were described by Lane et al. (Reference Lane2017), and are known to contain lithium-bearing minerals, such as purple mica and pink tourmaline. Other newly discovered and unmapped pegmatites of the Boulder Mountain Pegmatite Group, such as the Mole pegmatite, are barren and composed of quartz, feldspar, mica and tourmaline. These have variable brittle/ductile contacts with the country rock and the majority have a tourmaline comb structure at the contacts.
The Prof pegmatite is a lens-shaped intrusion exposed for ∼70 m in length and 5 m in width (Fig. 2a). It strikes northeast at 060° and crosscuts the foliation (110°, 60°, south) of the country rock. This pegmatite was described briefly in an assessment report by Lane (Reference Lane2017). The other pegmatites in this Group are in general thinner and more laterally extensive than the Prof pegmatite, with some reaching more than 100 m in length and ∼5 m in maximum width. Lithium-bearing pegmatites in the Boulder Mountain Pegmatite Group generally have a strike of ∼055–065° and vertical dip, whereas the abundant barren pegmatites are more randomly oriented. The Prof pegmatite is emplaced into the Monashee cover sequence within paragneiss of variable composition (mostly biotite, quartz, diopside and plagioclase). The paragneiss also hosts pinching bands of a combination of diopside, garnet, epidote, quartz and actinolite (Fig. S7). The country rock locally contains marble bands and units rich in nodular sillimanite. Significant host-rock tourmalinisation of the country rock has taken place at pegmatite contacts (Fig. 2b), however, there is little evidence in hand sample of other metasomatic halos surrounding the Prof pegmatite.
Outcrop photos of the Prof pegmatite. (a) Prof pegmatite outcrop. Geologist = 171 cm. (b) Country rock raft within the Prof pegmatite showing tourmalinisation at contacts. (c) Comb tourmaline structure marking contact zone between the intermediate sub zones. Graphic subzone to the left, overgrowth zone to the right. Pencil = 14 cm. (d) Cream-white petalite lenses surrounded by fine-grained purple lepidolite and hot pink Ca-chabazite and montmorillonite rim. (e) Central core containing radial pink elbaites with purple lepidolite roses and white feldspar. Notebook = 12 cm height. Lpd – lepidolite, Chb – chabazite, Mnt – montmorillonite, Ptl – petalite, Elb – elbaite, Or – orthoclase, Tur – tourmaline, Ms – muscovite.
The Prof is the most evolved pegmatite known of the Boulder Mountain Pegmatite Group with the Grail, Green and Red zoned pegmatites also containing evolved lithium-bearing minerals. Each of these pegmatites contain a central zone with the presence of purple lepidolite and elbaite, however, the Prof pegmatite is currently the only known pegmatite to contain abundant elbaite ([Na(Li1.5Al1.5)Al6(Si6O18)(BO3)3(OH)3(OH)] ∼2% of the total pegmatite mode) and petalite ([LiAl(Si4O10)] ∼5% of total pegmatite mode).
The Red pegmatite is a 5 m long × 1.5 m wide dyke that pinches to the southwest along a 060° strike dipping vertically. It crosscuts foliation and is surrounded locally by quartz lenses and regions of partial melt accumulation in the country rock. The Red pegmatite contains significant exo-contact tourmaline, and dominantly consists of graphic intergrowths of quartz and K-feldspar (90%) and tourmaline. Muscovite and biotite are present. The Green pegmatite forms a large thin dyke ∼5 m in width that extends for 200 m. The central zone of this pegmatite is ∼1 m thick and hosts trace elbaite and lepidolite together with rose quartz. Major pegmatite minerals are feldspar, quartz and tourmaline. The Grail pegmatite hosts a similar lithology to the Red pegmatite and is more laterally extensive along a 049° oriented strike (vertical dip). The mineralised region is minor, consisting of ∼1 m of a central zone that hosts lepidolite and zoned tourmalines from schorl to elbaite (black to green to pink). Major minerals are K-feldspar, quartz and large schorl crystals (up to 5 cm in size). The Mole is a newly discovered pegmatite located ∼40 m west of the Prof pegmatite. This intrusion strikes north–south, has a subvertical dip and contains mostly quartz and feldspar. A pocket zone hosts schorl sprays up to 30 cm in size. This contribution reports the mineralogy, petrography, geochemistry and paragenesis of the Prof pegmatite.
Methods
During 2021 and 2022, fieldwork involving detailed mapping, describing zonation and textural relationships, and recording structural measurements which revealed the extensive nature of pegmatite-hosted mineralisation on Boulder Mountain. A representative sample set was collected, which forms the basis of this investigation. To ensure this representation, a sample was taken from each zone in the pegmatites together with additional samples of each unique texture and mineralogical feature.
Mineral identification: powder XRD, Raman spectroscopy and SEM
Visual identification and polished thin-section petrography were used to identify minerals from each zone of the pegmatite. To confirm these observations, we used scanning electron microscopy (SEM) and energy dispersive spectroscopy (EDS), powder X-ray diffraction (XRD) and Raman spectroscopy. A Philips XL30 electron microscope with a Bruker Quantax 200 energy-dispersion X-ray microanalysis system was used at the University of British Columbia, Vancouver (UBCV). This highlighted the geochemical variability seen on back-scattered electron (BSE) images of polished thin sections and EDS provided qualitative measurements for mineral compositions. Scanning electron microscopy-mineral liberation analysis (MLA) maps of selected polished thin sections were created using a 50 μm grid of spot analyses. Compilation and mineral identification were completed using the Advanced Mineral Identification and Characterization System (AMICS) software. Minerals of similar elemental proportions were compared to petrographic characteristics for identification.
Raman spectroscopy was conducted with a Horiba XploRA Plus μ instrument at UBCV using a 532 nm laser, 200 μm slit, acquisition time of 30 s and an accumulation of 2. The spectra were compared to the RRUFF online database for identification (https://rruff.info/), which aided in identifying small grains within the polished thin sections. Larger unknown minerals were identified with powder X-ray diffraction employing a Bruker D8 Focus (0-20, LynxEye detector) instrument at UBCV.
Electron probe microanalysis (EPMA)
Mineral compositions were determined using a JEOL JXA-iHP200F field emission electron microprobe in the Department of Earth, Ocean and Atmospheric Sciences at UBCV. Compositional data were acquired for tourmaline, mica, feldspar, garnet and oxides and all X-ray intensities were processed together within the Probe for EPMA software (Probe Software Inc.). Detailed descriptions of the analytical conditions are given in Supplementary items 1 and 2. The instrument was operated at an accelerating voltage of 15 kV for all mineral phases. A beam current of 15 nA was used for the feldspar and mica and 20 nA for oxides, garnet and tourmaline. The beam diameter was 1 μm for garnet, 2 μm for oxides, 5 μm for tourmaline and feldspar and 10 μm for mica to mitigate volatile loss. Oxygen was calculated from cation stoichiometry and included in the matrix correction.
Feldspar data were quantified and corrected using the following standards: corundum (Taylor) for AlKα; wollastonite (Taylor) for CaKα; RbTiPO5 (Astimex) for RbLα; albite (SPI) for SiKα and NaKα; baryte (SPI) for BaLα; pollucite (SPI) for CsLα; celestite (SPI) for SrLα; crocoite (SPI) for PbMα; hematite (SPI) for FeKα; periclase (SPI) for MgKα; and orthoclase (SPI) for KKα. Structural formulae were calculated on the basis of 8 oxygens.
Mica compositions were quantified and corrected using the following standards: corundum (Taylor) for AlKα; wollastonite (Taylor) for CaKα; albite (SPI) for NaKα and SiKα; pollucite (SPI) for CsLα; bustamite (SPI) for MnKα; fluorite (SPI) for FKα; Hematite (SPI) for FeKα; periclase (SPI) for MgKα; rutile (SPI) for TiKα; and orthoclase (SPI) for KKα. Structural formulae were calculated on the basis of 24 (O + OH + F).
Oxide data were quantified and corrected using the following standards: Ta (Astimex) for TaLα; scheelite (Taylor) for WLα; columbite for NbLα and MnKα; cassiterite (Taylor) for SnLα; ScPO4 (NMNH168495) for ScKα; albite (SPI) for SiKα; magnetite (SPI) for FeKα; periclase (SPI) for MgKα; rutile (SPI) for TiKα; and cubic zirconia (SPI) for ZrLα. Structural formulae were calculated on the basis of 6 oxygens for Nb–Ta oxides, 2 oxygens for cassiterites; 4 oxygens for wolframites and 10 oxygens for qitianlingites.
Garnet data were quantified and corrected using the following standards: Nb (Astimex) for NbLα; Ta (Astimex) for TaLα; wollastonite (Taylor) for CaKα; apatite (Taylor) for PKα; albite (SPI) for SiKα; pollucite (SPI) for CsLα; bustamite (SPI) for MnKα; cassiterite (SPI) for SnLα; celestite (SPI) for SrLα; hematite (SPI) for FeKα; pyrope (SPI) for MgKα; AlKα and rutile (SPI) for TiKα. Structural formulae were calculated on the basis of 12 oxygens.
Tourmaline compositions were quantified and corrected using the following standards: albite (Taylor) for SiKα and NaKα; corundum (Taylor) for AlKα; spessartine (Taylor) for MnKα; ScPO4 (NMNH168495) for ScKα; RbTiPO5 (Astimex) for RbLα; pollucite (SPI) for CsLα; chromium oxide (SPI) for CrKa; diopside (SPI) for CaKα and MgKα; fluorite (SPI) for FKα; hematite (SPI) for FeKα; rutile (SPI) for TiKα; orthoclase (SPI) for KKα; tugtupite (SPI) for ClKα; and Cu (JEOL) for CuKα. Structural formulae were calculated on the basis of 31 anions, assuming stoichiometric amounts of H2O as (OH), i.e. OH + F = 4 atoms per formula unit (apfu), B2O3 (3 B apfu) and Li2O (as Li+) (Burns et al., Reference Burns, Macdonald and Hawthorne1994, MacDonald et al., Reference MacDonald, Hawthorne and Grice1993). The amount of Li assigned to the Y site corresponds to the ideal sum of the cations occupying the T + Z + Y sites (15 apfu) minus the sum of the cations actually occupying these sites [Li = 15 – (T + Z + Y) or Li = 15 – (Si + Al + Mg + Fe + Mn + Zn + Ti + Sc + Cr)]; the calculation was iterated to self-consistency. All EPMA compositions and detection limits are attached in the Supplementary materials (Supplementary 2) apart from the tourmaline data, which will form the focus of a subsequent study.
Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry (LA-ICP-MS)
Trace-element concentrations in mica were obtained using an ESL NWR 193 laser with a TwoVol3 sample cell paired with an Agilent 8900 tandem mass spectrometer operated in single spectrometer mode in the Fipke Laboratory for Trace Element Research (FiLTER) at the University of British Columbia, Okanagan, Canada. A circular laser spot with a diameter of 65 μm, a repetition rate of 10 Hz and an approximate fluence of 4 J/cm2 was used to ablate the material. A two-pulse pre-ablation using a 100-μm diameter laser beam was used to clean the sample surface prior to analysis. The main ablations lasted 30 s followed by 25 s of washout/background monitoring. Ablated material was moved out of the sample cell via 0.3 L/min He carrier gas and mixed with the sample Ar (0.85 L/min) through an in-house mixing bulb prior to entering the plasma. Additional N2 was added to the plasma (4 ml/min) to maximise sensitivity (Hu et al., Reference Hu, Gao, Liu, Hu, Chen and Yuan2008). The reference glasses NIST610 and NIST612 (Jochum et al., Reference Jochum, Weis, Stoll, Kuzmin, Yang, Raczek, Jacob, Stracke, Birbaum and Frick2011) were analysed to correct for instrument drift, normalise trace-element concentrations and verify the method. The glasses were sampled at the beginning and end of each experiment and every 10–15 unknown analyses. Trace-element concentrations were normalised to Si values measured at the same location via EPMA.
Results
Host rock
The host-rock contact contains significant tourmaline mineralisation (schorl–dravite) together with disrupted bands of phlogopite and quartz, plagioclase and garnet. Spindly needles of clinochlore are also present. Distinct, discontinuous blebby bands of garnet and diopside are found within the gneiss. The garnet forms massive bands of red–orange, extensively fractured crystals. The diopside has high relief and a strong cleavage. Epidote is also common in these bands and it infills cracks in the garnet. Actinolite is also present as a sea-green mineral in plane polarised light (PPL), which form bands of multiple-aligned stubby crystals. Locally, illite and magnetite are present. The mineral mode of the country rock is 30% phlogopite, 40% quartz, 2% garnet, 16% plagioclase and 2% diopside in modal percent (not including any imposed pegmatite-derived metasomatism).
The Prof pegmatite
The Prof pegmatite can be divided into four zones: (1) the border zone; (2) the intermediate zone; (3) the central zone; and (4) the quartz zone. A mineralisation map of the Prof pegmatite is shown in Fig. 3 and thin-section examples of mineral relationships are shown in Figs 4 and 5. All data from this investigation are in the Supplementary materials (Tables S3–6).
(a) Schematic representation of the Prof pegmatite and mineralogical details (some minerals exaggerated in size to show texture). (b) Zonation of the Prof pegmatite (border zone exaggerated, other zones to scale).
Photomicrographs of thin sections from the Prof pegmatite in cross-polarised light: (a) perthite with plagioclase exsolution (sample PFS1C from graphic subzone); (b) overgrowth sequence of garnet centre coated with secondary mica and tourmaline and quartz intergrowths (sample PFG2A from overgrowth subzone); (c) two generations of mica growth shown by difference in birefringence colours. Generation 1 shows yellow to pink birefringence muscovite, generation 2 shows blue to purple birefringence trilithionite (sample PFL12F from central zone). (d) Elbaite with later-stage quartz vein cutting through crystal (sample P26B from central zone core). Or – orthoclase, Ms – muscovite, Qz – quartz, Elb – elbaite, Gr – garnet, Tur – tourmaline, Ab – albite, Tln - trilithionite.
Thin section example images of alteration and metasomatism in the Prof pegmatite. (a) Orthoclase with metasomatic veins of albite (sample PFL3B). (b) Petalite breaking down into chabazite and montmorillonite. Metasomatic assemblage of lepidolite, quartz and albite cutting through sequence (sample PFL6B). (c) Quartz recrystallisation due to fluid flow (sample PFS1B). (d) Orthoclase with metasomatic veins of sericite (sample PFS1B). Lpd – lepidolite, Chb – chabazite, Mnt – montmorillonite, Ptl – petalite, Ser – sericite.
Border zone
The border zone mineralogy is variable and forms a mm-scale band surrounding the pegmatite (Fig. 3a). The major mineral abundances of the border zone are (by volume) 32% quartz, 30% plagioclase–quartz myrmekite, 20% perthitic K-feldspar, 10% tourmaline (which is found locally in comb structures perpendicular to the country rock contact), and 8% brown mica. These values are estimated using multiple thin sections. Abundant euhedral 1 mm crystals of magnetite and biotite are found at the contact within both the country rock and the pegmatite (Fig. 3a).
The contact between the pegmatite and the country rock is difficult to demarcate in thin section with the pegmatite containing irregular areas of disaggregation of country rock. The border zone hosts tourmaline, which has strong zonation with a pale brown/olive green core encompassed by a darker orange, brown rim. These 0.5 cm crystals host quartz inclusions at the rim. Smaller, 0.1 cm tourmaline crystals (dravite–schorl) with indigo blue cores with small rims of green and brown colouration. Another major mineral present is quartz, which has irregular grain boundaries and equant bands that surround the larger crystals. Large crystals of K-feldspar have perthitic unmixing. Orthoclase crystals are replaced commonly by a myrmekitic plagioclase–quartz intergrowth. The feldspars are microcrystalline in some instances, although they are distinguishable by their mottled dark-brown appearance and association with fine needles of sericite. Magnetite is common at the country rock contact; the crystals are internally zoned, 0.5 cm in size and euhedral. A mica and clay assemblage is also found associated with the magnetite. Crystals lack preferred orientation and have a chaotic crystalline growth mesh. Fine-grained micas replace minerals and minor veins of albite are present.
Intermediate zone, graphic subzone
Due to the similar mineralogy and variable presence of the graphic subzone and the overgrowth subzone (below), they are classified within the same intermediate zone (Fig. 3b). In the graphic subzone macro- (>1 cm) to micro-scale (<0.2 mm) graphic intergrowths of K-feldspar (65% mode of intergrowth) and quartz (35% mode of intergrowth) dominate. Muscovite books are common and comprise up to 8% of the bulk unit locally. Schorl–dravite is also common in the unit and comprises up to 5% of the zone in some areas. Aplite bands run parallel to the intrusion, and frequently occur with a composition of quartz, perthitic K-feldspar, albite and intensely zoned schorl–dravite (Fig. 3b; Supplementary 7–9). The aplites do not have crosscutting textures in thin section. Accessory minerals include garnet (up to 0.2 mm in size) and apatite.
The graphic subzone consists of quartz, albite, perthitic microcline, tourmaline and apatite, and varies in texture from large interlocking crystals to fine-grained aplitic bands. The tourmalines are strongly zoned and vary in colour from deep blue to olive green to brown and orange and range from 0.1 mm equant crystals to 2 cm in the coarse portions of the unit. The tourmalines are generally pristine with little evidence of post-crystallisation alteration with a dravite–schorl composition. Plagioclase feldspar and orthoclase crystals appear internally mottled and vary in size from 0.5 mm to 3 cm. Perthitic K-feldspar contains macroscopic intergrowths of albite and orthoclase (Fig. 4a).
Intermediate zone, overgrowth subzone
This subzone lacks graphic intergrowths, tourmaline is much more abundant. It is dispersed throughout the unit rather than occurring exclusively in aplite bands or comb textures as is common in the graphic zone. A distinctive feature of this zone is the growth of multiple minerals in rims around older phases, which can be seen distinctly in hand sample and thin section (Fig. 4b) together with the development of a black-and-white speckled rock (Fig. 2c). The major mineral assemblage of this unit consists of quartz, perthitic K-feldspar, plagioclase and tourmaline.
The tourmalines of the overgrowth subzone can be much larger than those found in the graphic zone (up to 5 cm in size) and black tourmaline (schorl–dravite) clots can be found rimmed with quartz in spray textures (Fig. 2c; Fig. S8). Garnets are a common accessory phase, attaining 1 cm in size; these are locally rimmed with tourmaline (fluor-elbaite), mica and quartz. Biotite is common in the southern edge of the deposit close to the wall zone. Accessory minerals include muscovite, garnet, apatite and amblygonite. The border zone between the graphic zone and the overgrowth subzone locally has large radial tourmaline sprays surrounding the graphic crystals (Fig. 2c). This zone also contains a complex assemblage of finely intergrown phosphates provisionally identified as graftonite [Fe2+Fe2+\! 2(PO4)2], crandallite [CaAl3(PO4)(PO3OH)(OH)6], and childrenite [Fe2+Al(PO4)(OH)2·H2O]. These minerals are not within the scope of the current investigation and will be the focus of future work.
The spessartine–almandine garnet cores are euhedral, ∼5 mm in size, well preserved and have no significant inclusions or alteration. Qitianlingite [Fe2+\!2Nb2W6+O16] and hübnerite [Mn2+(WO4)] are found within cracks in garnet and dispersed elsewhere throughout the overgrowth subzone. The surrounding micas are bladed and <1 mm in size (Fig. 4b). Tourmalines with inclusions of quartz are found surrounding the garnet (Fig. 4b). Radial veins of pure albite are present in thin section. Plagioclase feldspar can be found oriented along the vein or as radial fans that cut into crystals (Fig. 5a). Other crosscutting veins contain trilithionite, albite, quartz and elbaite.
Central zone
This zone features bands of different mineral associations surrounding the most evolved core (Fig. 3a). The core of the Prof pegmatite contains radial sprays of pink elbaite surrounded by aggregates of purple trilithionite and quartz. The trilithionite occurs in a variety of sizes, from 2 cm when associated with the elbaite sprays and down to mm-scale more distal to the central evolved core (Fig. 2e). Brown perthite, which is mottled internally, occurs to the south of the central core and forms ≤20 cm megacrysts. The brown feldspars are cross-cut by multiple veins of pale white albite.
The quartz and fine mica-rich distal tourmaline unit of the central zone also hosts a variety of phosphates, including amblygonite, lithiophilite, triphylite, apatite, and possibly graftonite, crandallite and childrenite. Niobium- and Ta-oxides [columbite-(Mn), columbite-(Fe)] and cassiterite are also present. The zone also hosts a variety of tourmaline including blue, green and pink fluor-elbaite.
Petalite is a major phase in the pegmatite and forms 10–15 cm lens-shaped crystals (Fig. 2d; Fig. S9a). These can be recognised by their strong cleavage and vitreous lustre. The pods are surrounded by a hot-pink-coloured rim consisting of Ca-chabazite and montmorillonite, which aids in field identification (Fig. 2d). The remainder of this unit is infilled with fine-grained purple trilithionite and white muscovite. Coarse books of white mica surround the petalite zone, which mark the transition into the graphic zone to the east. Vugs infilled with other zeolite group and clay minerals are also present.
The petalite forms altered macrocrysts which are cross-cut by fine bands of sericite, visible in thin section surrounded by Ca-chabazite and montmorillonite forming a dark mottled microscale aggregate (Fig. 5b). The petalite is cross-cut by veins, containing quartz, mica, albite and elbaite. The mica crystals are stubby and oriented. Albite is aligned parallel to veins and forms tabular elongate crystals up to 1 cm in size. Microcrystalline quartz interstitially fills the veins. The majority of the petalite area is microcrystalline and contains sericite (Fig. 6d) and microcrystals of spodumene (?) which are smaller than the measurement grid spot size (Fig. 6). Elbaite is found close to this region intergrown with quartz and brown feldspar and is crosscut by multiple veins of pure albite. Veins have either fine-grained quartz aggregates or albite and radial mica at their contacts. Lithiophilite [LiMn2+(PO4)] is surrounded by an intergrowth rim of tourmaline and quartz (Fig. 6c). Phosphates are typically found in complex intergrowths with diverse minerals.
Thin-section mineral-distribution maps of various regions of the Prof deposit. (a) Contact zone (sample PFL7B). (b) Graphic subzone aplite band contact (sample PFS1C). (c) Central zone mica zonation (sample PFL12F). (d) Central zone petalite and lepidolite (sample PFL6A).
Elbaites from the core of the central zone from the elbaite spray are locally preserved pristinely and do not have the intergrown appearance of elbaites from other portions of the pegmatite. In some instances, these elbaite are crosscut by veins of quartz (Fig. 4d).
Quartz core
This is a massive pod-shaped quartz unit ∼1–3 m in size, with large blue–black peripheral tourmaline crystals (Fig. 3a). Trace minerals of this zone include 0.1 mm mica and sub-millimetre euhedral beryl which are inclusions within ≤ 3 cm schorl crystals. The quartz appears to be cut by veins of microcrystalline quartz.
Mineral geochemistry
Potassium feldspar
Potassium feldspar in the Prof pegmatite ranges from milky white in the border and intermediate zones to a deep brown colour in the central zone. They commonly occur as blocky crystals and the majority of the K-feldspar are perthitic. Point analyses of the K-feldspar gave an average composition of Or90Ab10 within perthites at the contact zone, Or95Ab5 within perthites in the intermediate zone, Or93Ab7 in the brown feldspar perthite of the central zone and Or92Ab8 in the distal tourmaline region of the central zone. Rubidium is the most commonly substituting element with up to 0.51 wt.% Rb2O. Strontium and Cs are also common substituting elements in minor quantities, generally <0.05 wt.% oxides. Representative compositions of K-feldspar are given in Table 1 and plotted in Fig. 7.
(a) Or–Ab–An feldspar ternary diagram showing the composition of the feldspars from the Prof pegmatite. (b) Enlarged albite corner of the ternary diagram.
Feldspar compositional data (wt.%; EMPA)of the Prof pegmatite.
n.d. – not determined
Plagioclase feldspar
Plagioclase feldspar occurs in three different varieties at the Prof pegmatite, as: (1) blocky crystals; (2) exsolutions within perthite; and (3) crystals that cross-cut earlier minerals. The blocky plagioclase is the most variable in composition ranging from pure albite, Ab100, to compositions that plot in the oligoclase field in the border zone, Or1Ab88An11. In thin section, blocky plagioclase was distinguished from cross-cutting plagioclase by the textural and optical relationships to other minerals, i.e. the blocky plagioclase is highly mottled and the cross-cutting plagioclase appears pristine. Exsolved plagioclase in perthite hosts albite of variable composition ranging from Or1Ab99 in the brown feldspar zone to Or1Ab93An6 in the intermediate zone graphic subzone. Cross-cutting plagioclase has a narrow-range in composition close to the albite end-member with an average of Or1Ab99. Representative compositions of plagioclase are given in Table 1 and illustrated in Fig. 7.
Mica
The micas of the border zone have the closest compositions to end-member muscovite in relation to the pegmatite as a whole (average mica composition: [K0.95Na0.08]Al2.73Si3.10O10[OH]1.9F0.1). They are silver in appearance and typically <1 cm in size. Back-scattered electron images show that there is no notable compositional zonation within these micas. They have the lowest Li (Li2O 0.1 wt.%) and F (0.11 wt.%) contents of the mica in the Prof pegmatite. The micas of the intermediate zone have a well-defined BSE-image zonation with a muscovite core (e.g. [K0.91Na0.11]Al2.83Si3.03O10[OH]1.80F0.20) and a trilithionite rim (e.g. K1.91Na0.06[Li1.52Al1.23][Al0.58Si3.42O10][F0.73{OH}0.27]2). These two minerals appear very similar in colour with a grey to pale purple appearance in hand sample. The rims commonly have a jagged contact with the muscovite and are associated with cross-cutting veins as observed in thin sections (Fig. 5b). Data in Figs 8 and 9a–c show the geochemical variation between the muscovite core and trilithionite rims. The Li2O contents range from 0.37 wt.% in muscovite to 5.73 wt.% in trilithionite. Fluorine ranges from 0.95 wt.% in the muscovite to 7.92 wt.% in the trilithionite. The micas of the central zone also show similar relationships of muscovite crystallisation with a trilithionite rim. The majority of purple micas in the deposit are classified as trilithionite, however, one thin section (P26B) has polylithionite compositions from the central zone (K0.97Na0.03)(Li1.5Al1.25)(Al0.32Si3.68O10)(F0.87[OH]0.13)2. The muscovites contain Li2O ranging from 0.63 to 2.91 wt.%. The trilithionite–polylithionites contain Li2O ranging from 4.55 to 6.54 wt.%. Finally, the micas of the quartz zone are zinnwaldites (K0.98Li0.96(Fe0.79Mn0.18)Al0.96(Al0.89Si3.11)O10(F0.63, OH0.37)2 average) with elevated Fe contents of up to 14.33 wt.% FeO. Representative mica compositions are presented in Table 2.
Classification scheme for mica (after Tischendorf et al., Reference Tischendorf, Forster, Gottesmann and Rieder2007) showing both magmatic and metasomatic mica. Black circles indicate end-member compositions and grey circles indicate ideal member compositions. Calculated apfu values were halved to fit the classification scheme calculated at 10 O and 2 (OH, F) apfu.
Compositional variation of K/Rb vs. Rb (a), Li (b) and Cs (c) in micas from the Prof pegmatite with data point colours showing pegmatite zone location and shape reflecting origin determined by thin-section relationships and mineral compositions.
Mica compositional data (wt.%; EMPA) of the Prof pegmatite.
n.d. – not determined
Tourmaline
Tourmalines of diverse texture, colour and composition are abundant in the Prof pegmatite and occur in every zone. A summary of the tourmaline textural variations and compositions is presented below
All tourmaline identified in the Prof pegmatite are of the alkali-group including dravite, schorl and elbaite (in some instances fluor-elbaite and fluor-schorl; Table 3). The tourmaline of the border zone are typically zoned and found dispersed throughout this zone as fine-grained disseminated crystals or as comb-textured prisms orientated perpendicular to the contact. Tourmaline crystals in the intermediate zone have strong zonation. In the aplitic regions, the tourmaline cores are dravite with high contents of Mg and Fe at the Y site. These correspond with a black colour in hand sample and a deep blue hue in thin section. These cores are rimmed progressively by a dull green coloured zone corresponding to an increase in Fe content of the Y site and transition to classsifcation as schorl (Y site Fe > [Mg or Li+Al]). These are locally coated by a rim of a F-rich tourmaline, which can either be fluor-elbaite (containing Li and Al as the dominant substitution at the Y site) or fluor-schorl.
Tourmaline compositional data (wt.%; EMPA) of the Prof pegmatite.
n.d. – not determined
A dispersed area of multi-coloured tourmalines exists in the outer regions of the central zone, close to the contact with the intermediate zones. This region is termed the distal tourmaline subzone, and hosts an impressive array of tourmaline colours (green, pink and blue), all of which are fluor-elbaite. Tourmaline in the central core zone are radial, elongate and pink in colour. They are fluor-elbaite and commonly have elevated Mn substitution at the Y site (up to 0.32 apfu). They are relatively homogenous and some are locally included within quartz. The central core zone contains a radial spray of pale pink fluor-elbaite. The quartz zone hosts large blue–black fluor-schorl tourmaline at the periphery of the zone. Representative compositions can be found in Table 3 and relationships between tourmaline Mn and F content are ilustrated in Fig. 10.
Variation of F vs. Mn in tourmaline from the Prof pegmatite with point colours showing pegmatite zone location and shape reflecting the EPMA point position within the mineral.
Nb–Ta oxide minerals
Nb–Ta oxides exist as disseminated grains (Fig. 11) or inclusions in cassiterite and are found dominantly in the central zone and the intermediate zone in minor amounts. Compositions show a transition from columbite-(Fe) to columbite-(Mn) (Table 4). The majority of these data plot within the columbite-(Mn) quadrant. There is dominance of Nb over Ta with the Ta/(Ta+Nb) ratio ranging from to 0.07 to 0.44 apfu (Fig. 12). There is also dominance of Mn over Fe with the Mn/(Mn+Fe) ratios ranging from to 0.15 to 1 apfu. Common trace elements include Ti, Zr and Sn. The compositional variation between the oxide minerals can be attributed to the exchange of Ta for Nb and Mn for Fe producing a linear correlation between these elements. Zonation within individual grains is typically concentric as seen in Fig. 11.
Element-distribution maps (EPMA) of Nb–Ta oxides from sample PFL2A, central zone. (a) BSE image with corresponding point data shown in Table 4, (b) Nb intensity map, (c) Ta intensity map and (d) ratio of Nb to Ta map showing Nb5+ – Ta5+ substitution.
Compositional variation of Nb–Ta oxides; arrows indicate evolutionary trends.
Columbite–tantalite compositional data (wt.%; EMPA) from the Central Distal Tur Zone (PFL2A) from EPMA of the Prof Pegmatite.
n.d. – not determined
Cassiterite
Cassiterite (SnO2) is a common accessory mineral found in the intermediate and central zones of the Prof pegmatite. It contains a suite of inclusions, including Nb–Ta oxides, as observed in compositional transects across inclusions as described by Neiva (Reference Neiva1996). Point data can be seen in Table 5. The overall cassiterite compositions range from pure 100 wt.% SnO2 to 90.05 wt.% SnO2 (with substitution of 6.31 wt.% Ta2O5 and 2.43 wt.% Nb2O5).
Nb-W-oxide compositional data (wt.%; EMPA) of the Prof Pegmatite.
n.d. – not determined
The cassiterites are highly compositionally zoned with elevated Mn, Ta and Nb contents. These relationships can also be seen in Fig. 13 where the core of the cassiterite hosts low Sn and elevated Mn, Nb and Ta contents. The core is coated by a complex, chaotic and oscillatory zonation pattern highlighted by variable contents of Sn, Ta and Mn. The bright red zones at the outer rim of the crystal map (Fig. 13b–d) are qitianlingite.
Element-distribution maps from EPMA of a zoned cassiterite crystal (PFL2A) showing Sn substitution for (b) Mn, (c) Ta and (d) Nb – microprobe point 243 in Table 5.
Garnet
Garnet is found in the intermediate zone within the overgrowth subzone, has an average composition of Sps64Alm36 and commonly contains elevated P values of up to 0.57 wt.% P2O5.
Other minerals
Tungsten-bearing oxides are present in the Prof pegmatite, and include minerals of the wolframite family (ferberite [Fe2+WO4] and hübnerite [Mn2+WO4]) with the latter dominating the assemblage. The wolframite ranges from Hbr86Feb14 to Hbr98Feb2. The hübnerites contain elevated levels of Nb up to 3.58 wt.% Nb2O5 and Ta (up to 0.94 wt.% Ta2O5) with Sc contents of up to 0.76 wt.% Sc2O3. Another rare W-bearing mineral, with Mn>Fe, is qitianlingite. The ideal formula of qitianlingite is Fe2+\!2Nb2W6+O10 and its average composition observed in the Prof pegmatite is Mn1.23Fe0.72Nb1.86Ta0.15WO10.
The qitianlingite ([Fe,Mn]2[Nb,Ta]2WO10) contains up to 1.76 wt.% TiO2. Point data for these minerals are presented in Table 5. Multiple microscopic scale phosphates were identified in the Prof pegmatite together with zeolites and clay minerals. Descriptions of these minerals are not the scope of the present investigation.
Discussion
The Prof pegmatite hosts abundant large crystals of petalite in high proportions in the central zone. Thus, it can be classified as a rare element class, lithium subclass, complex type, petalite subtype pegmatite in the Černý and Ercit (Reference Černý and Ercit2005) classification system, and as a Group one pegmatite, using the Wise et al. (Reference Wise, Müller and Simmons2022) classification scheme. Several other petalite subtype pegmatites can be found in Canada, most notably the Tanco pegmatite in Manitoba, in which petalite pervasively recrystallised to spodumene and quartz intergrowths (“SQUI”, Černý, Reference Černý2005). Other petalite subtype pegmatites include the Big Whopper, Big Mac and Marko’s pegmatites in the Separation Rapids area of northwestern Ontario. The Big Whopper and Big Mac pegmatites contain abundant petalite with minor spodumene, lepidolite, feldspars and beryl. They also host a variable suite of oxides, including Ta-bearing wodginite (Tindle and Breaks, Reference Tindle and Breaks2000). The dominant mineralogy of Marko’s pegmatite is beryl, muscovite, quartz, albite, petalite and a variety of Ta–Nb oxides (Tindle and Breaks, Reference Tindle and Breaks1998). The Prof is the first petalite subtype pegmatite identified in British Columbia, Canada.
Paragenesis
The zones of the Prof pegmatite are interpreted to have crystallised in the following order: fine crystalline border zone; intermediate zone; quartz; and central zones. A paragenetic sequence for crystallisation of the Prof pegmatite is given in Fig. 14. The fine border zone is considered to represent a chilled margin showing a finer grained crystalline assemblage representative of the bulk intermediate zone, typical of most pegmatites (London, Reference London2008). The mineral relationships and compositional data show that quartz, magnetite, biotite and dravite then schorl grew contemporaneously as the primary minerals (Fig. 6a), followed by an interstitial infill of orthoclase and plagioclase.
Paragenetic sequence of mineral crystallisation in the Prof pegmatite divided into zones representing magmatic versus metasomatic versus hydrothermal events. Brd – border zone, Int – intermediate zone, Qz – quartz zone, Cent – central zone, Alb – albitisation.
The intermediate zone of banded aplites (composed of quartz, perthitic K-feldspar, plagioclase (aforementioned blocky plagioclase) and zoned tourmalines) of the graphic subzone are interpreted to have formed simultaneously, as shown by the equigranular interlocking nature of the minerals. These relationships can be seen in Fig. 6b, with aplitic equigranular relationships to the left of the figure. In other non-aplitic regions of this zone, the first generation of mica, quartz, orthoclase, amblygonite, apatite and plagioclase crystallised contemporaneously, as shown by planar contacts in thin section. The overgrowth subzone is represented by initial garnet and Nb–Ta oxide crystallisation followed by coating of these crystals with a combination of quartz, mica and tourmaline. The graphic subzone quartz and feldspar crystallised before the overgrowth subzone crystallised providing a nucleation surface for which comb tourmaline of the overgrowth subzone could grow; separating the two subzones. Fig. 3).
The central zone is interpreted to be the last region of the pegmatite to have crystallised due to the high abundances of exotic minerals. Petalite and orthoclase are interpreted to have crystallised simultaneously as indicated by their euhedral habit. Zoned tourmaline, phosphates and Nb–Ta oxides also crystallised early in the paragenetic sequence, as shown by their euhedral morphology and planar and equant textural relationships. Primary elbaite crystallised together with plagioclase forming large euhedral crystals. The beginning of mica crystallisation occurred at the same time as elbaite crystallisation and with the formation of magmatic muscovite. Mica continued to crystallise after the conclusion of plagioclase and elbaite crystallisation, as shown by planar and cross-cutting replacement relationships. There were two generations of mica crystallisation. The first grew contemporaneously with plagioclase. A late-stage albitisation event formed secondary albite (aforementioned cross-cutting plagioclase) and produced overgrowths of secondary Li-rich trilithionite mica on the first phase of muscovite crystallisation (Figs 4c; 6c). Elbaites are zoned as observed in BSE images, exhibiting both oscillatory magmatic and zonation-free metasomatic origins as determined by thin-section textural relationships.
Petalite pods in the central zone are surrounded by a pink-coloured weathering rim composed of Ca-chabazite and montmorillonite, as determined by powder XRD analysis, and result from the breakdown of petalite. The abundance of petalite in the pegmatite is an important factor to consider when deciphering the evolution of the Prof pegmatite. To crystallise petalite, significant lithium enrichment of the melt is required (Maneta et al., Reference Maneta, Baker and Minarik2015). The presence of petalite indicates crystallisation at relatively low pressures (∼1.5–3 kbar) (London, Reference London1984). The absence of spodumene and quartz intergrowths in the pegmatite indicates that there was not a re-equilibration of the Li-aluminosilicate minerals as the pegmatite cooled or that crystallisation occurred at low temperatures within the stability field of petalite. Needle-like bands of very fine-grained spodumene (?) cut the petalite indicating that some crystallisation occurred in the spodumene stability field, this being potentially due to deformation-related or secondary recrystallisation events. The petalite is also cross-cut by a Li-rich metasomatic assemblage, which contains quartz, trilithionite, albite and elbaite. Complex intergrown phosphates, such as childrenite, resulted from the breakdown of primary phosphate phases, such as triphylite [LiFe2+(PO4)] (Moore, Reference Moore1973; Simmons et al., Reference Simmons, Webber, Falster, Roda-Robles and Dallaire2022).
The quartz zone is composed of massive quartz, which has been recrystallised in a patchy manner. Fluid-flow pathways can be identified by flow recrystallisation structures shown in micro-crystalline equant quartz (similar to those shown in Fig. 5c). Due to the presence of less-evolved schorl tourmaline of magmatic origin, this zone is interpreted to have crystallised prior to the central zone. This tourmaline is believed to be geochemically linked to some degree of country rock contamination. The determination of geochemical origins of the tourmalines were based on similar textural features and mineralogical associations identified in thin section and BSE image zonation trends.
Three major metasomatic events have been identified within the Prof pegmatite. The first was an albitisation event in which albite replaced orthoclase (Fig. 5a). A second event gradually transitioned from pure albitisation to late-stage albitisation, that produced veins hosting Li- and F-rich trilithionite to polylithionite and elbaite, in addition to quartz and albite (Fig. 5b). The fluid infiltrated crystal boundaries causing a flow-like recrystallisation texture, as illustrated in Fig. 5b. The final alteration event was mass recrystallisation of quartz (Fig. 5c), and pervasive sericitisation of feldspars (particularly in the intermediate zone; Fig. 5d). The source of the K for this event could be from orthoclase breakdown related to episodes of albitisation.
Geochemical evolution and fractional trends
Various minerals in pegmatites are known to record the geochemical evolution and fractionation of melt as indicated by the incorporation of incompatible elements. K-feldspars are commonly used as geochemical monitors of melt evolution and fractionation. These can incorporate Li, Rb, Be, Sr and Ba (Roda-Robles et al., Reference Roda-Robles, Pesquera, Gil-Crespo and Torres-Ruiz2012). When compositions are plotted against the K/Rb ratio, the resultant trends can be considered as an indicator of melt fractionation. Increasing incompatible element contents coupled with a decreasing K/Rb ratio indicates increasing fractionation (Černý and Burt, Reference Černý, Burt and Bailey1984; Wise et al., Reference Wise1995; Selway et al., Reference Selway, Breaks and Tindle2005). Micas in pegmatites incorporate the incompatible elements Li, Cs, Rb, Ta and Nb into their structure, and can also be used as chemical evolution indicators (Fleet et al., Reference Fleet, Deer, Howie and Zussman2003). In common with feldspars, when compositions are plotted against K/Rb ratio, they indicate the evolution of the melt. Mica and feldspar are useful minerals as they are typically major components of pegmatites and can record incompatible element variability throughout the entire pegmatite crystallisation sequence (e.g. Marchal et al., Reference Marchal, Simmons, Falster, Webber and Roda-Robles2014; Garate-Olave et al., Reference Garate-Olave, Roda-Robles, Gil-Crespo and Pesquera2018). Tourmaline commonly shows a progressive evolutionary trend from Fe- and Mg-bearing schorl and dravite to Li-bearing elbaite. The variation in tourmaline Fe, Li, Al, Mn and F corresponds to pegmatite evolution (Selway et al., Reference Selway, Novák, Černý and Hawthorne1999). The Ta, Nb, Fe and Mn contents of Nb–Ta oxides record geochemical evolution information (Van Lichtervelde et al., Reference Van Lichtervelde, Salvi, Beziat and Linnen2007) with an increasing evolutionary trend indicated by an increase in Mn/(Mn+Fe) followed by an increase in Ta/(Ta+Nb).
The feldspars of the Prof pegmatite contain limited elemental substitution of incompatible elements and are almost entirely composed of Si, Al, Na, K and Ca. The low concentrations of incompatible elements in feldspar result in poorly-defined geochemical trends, Hence, feldspar is a poor geochemical proxy for melt evolution in the Prof pegmatite. The feldspar compositions could be a result of chemical re-equilibration during the formation of microcline and perthite as suggested by Teertstra et al. (Reference Teertstra, Černý and Hawthorne1998). The geochemical evolution of the Prof pegmatite can be best assessed by the trace-element behaviour of mica (Fig. 9), tourmaline (Fig. 10) and Nb–Ta oxides (Fig. 12). From petrographic study, BSE evidence, textural relationships and interpretation and mineral composition in the Prof pegmatite, two main compositional groups of mica were identified to have crystallised after the more primitive border-zone muscovites. The first group to crystallise is white to pale-purple muscovite of relatively consistent composition (Fig. 8), with an average Li2O content of 0.68 wt.% and F content of 1.79 wt.%. These are micas that have contemporaneous growth with blocky feldspars, quartz and other rock-forming pegmatite minerals. Additionally, it is common to see these primary magmatic mica altered by secondary metasomatic processes as shown by cross-cutting relationships and corroded rims of secondary minerals.
The secondary purple micas of the trilithionite–polylithionite group are metasomatic in origin due to the contemporaneous crystallisation relationships with cross-cutting veins of metasomatic assemblages, such as the albitisation in the deposit. This group have compositions between the trilithionite and polylithionite end-member compositions and have more compositional spread than the muscovite. The trilithionite has an average Li2O content of 5.35 wt.% and F content of 6.96 wt.%. The secondary micas are commonly found crosscutting previous generations of minerals and forming rims around minerals crystallised previously. Due to the association of trilithionite, and in some cases polylithionite, with the ablitisation metasomatic event, it is concluded that the metasomatic event that crystallised the lepidolites was enriched in F and Li. These geochemical associations correspond with similar elevations of F and Li in fluor-elbaite rims (Table 3), which were determined to also be secondary metasomatic and associated with this same albitisation phase. These relationships have been noted in the neighbouring Mount Begbie pegmatites by Dixon et al. (Reference Dixon, Cempírek and Groat2014). There is no apparent connecting trend between the two mica groups in Fig. 8, and when coupled with thin-section evidence of magmatic vs. metasomatic origins, this indicates two separate crystallisation events. Alternatively, a miscibility gap has been noted in numerous studies between the dioctahedral muscovite and triocta-hedral polylithionite micas, which could explain the groupings (Foster, Reference Foster1960; Monier and Robert, Reference Monier and Robert1986; Roda-Robles et al., Reference Roda-Robles, Pesquera, Gil-Crespo, Torres-Ruiz and De Parseval2006).
A minor composition group exists of zinnwaldite [K0.98Li0.96(Fe0.79Mn0.18)Al0.96(Al0.89Si3.11)O10(F0.63, [OH]0.37)2 average], which are located at the borders of the quartz zone. These micas have elevated FeO contents of up to 14.33 wt.% (Fig. 8). These minerals are interpreted to be the products of late-stage metasomatism of a previously country-rock contaminated portion of the pegmatite.
The micas in the border zone are representative of the most primitive micas in the system, as reflected in their low Li and F compositions. The intermediate zone contains mica with magmatic cores and trilithionite secondary metasomatic rims (Fig. 9). The presence of more geochemically evolved mica rims complements the rimmed tourmaline texture of the intermediate zone. These rims are interpreted to be evidence of a later metasomatic fluid that could have formed in the central zone from where it moved outwards, subsequently metasomatising the intermediate zone. Textures similar to this are observed and interpreted to be of metasomatic origin in the Wekusko Lake pegmatite field (Benn et al., Reference Benn, Martins and Linnen2022). The central zone of the Prof pegmatite hosts a suite of magmatic and metasomatic mica, which have a slight trend towards polylithionite compositions (Fig. 8). Finally, the presence of zinnwaldite and elevated Fe contents of micas in the quartz zone is comparative to the fluor-schorl rims on the tourmaline. This late-stage increase in the Fe content of the quartz zone might be evidence for country rock or fluid contamination of the pegmatitic forming melt prior to metasomatism. Country rock rafts are commonly found in the Prof pegmatite, which supports this hypothesis. Country rock contamination as a late-stage source of Fe has been noted in other studies e.g Novak et al. (Reference Novak, Kadlec and Gadas2013).
Geochemical evolution diagrams of mica in the Prof pegmatite with the K/Rb ratio vs. the incompatible elements Rb, Li and Cs, are shown in Fig. 9. Micas have a strong fractionation trend with increasing Rb as the melt becomes more evolved (Fig. 9a). These data support thin-section textural observations with metasomatic later-stage micas hosting elevated Rb contents and earlier-crystallised primary micas hosting less Rb. The quartz zone micas are interpreted to have been metasomatic replacements of pre-existing country rock contaminated minerals. This could explain their plotted locations on Fig. 9a. A similar relationship is seen between K/Rb and Li with a gradual increase in Li as micas transition from early magmatic to later metasomatic types. The relationship between K/Rb and Cs in general shows a relatively low concentration of Cs in the pegmatitic melt with a geochemical evolution trend of less Cs in the primary magmatic mica and more in the secondary metasomatic micas. The Cs contents of micas of the quartz zone are outliers with values up to 6508 ppm. The micas of the quartz zone are metasomatic in origin and could represent one of the last metasomatic events in the Prof pegmatite due to their unusually high Cs content compared to the micas of the rest of the deposit. The elevated Cs in these micas could be attributed to Cs not accumulating in sufficient quantities to crystallise pollucite. These elevated Cs contents could also be attributed to crystallisation in a boundary layer as noted by Benn et al. (Reference Benn, Martins and Linnen2022) and London et al. (Reference London2022).
The tourmaline shows a positive correlation between Mn and F (Fig. 10), which is indicative of incompatible element accumulation in the residual melt of the pegmatite (Selway et al., Reference Selway, Novák, Černý and Hawthorne1999). As fractional crystallisation progresses, and fewer compatible elements accumulate, they are forced to enter the tourmaline crystal structure, that is recorded in the progression of core to rim mineral composition (Jolliff et al., Reference Jolliff, Papike and Shearer1986; Selway et al., Reference Selway, Novák, Černý and Hawthorne1999). Figure 10 illustrates progressive geochemical evolution from the intermediate, followed by the quartz zone, and finally, to the central zone. During the tourmaline crystallisation in the intermediate zone, Mn remains low and F values vary. This could be due to garnet preferentially incorporating Mn at approximately the same time in the intermediate zones (Maner et al., Reference Maner, London and Icenhower2019). As the tourmalines become more evolved and the central and quartz zones crystallise, the F content of tourmalines increases at a more gradual rate, and there is a large spread in Mn contents. This might be attributed to the conclusion of crystallisation of garnet in the intermediate zone allowing Mn to accumulate and enter the tourmaline structure. This observation indicates that F is a good elemental indicator of geochemical evolution in this pegmatite, whereas Mn availability is being controlled by garnet (Maner et al., Reference Maner, London and Icenhower2019). Tindle and Breaks (Reference Tindle and Breaks2000) have also discussed the idea of garnet buffering Mn in the pegmatites of the Separation Lake area, Ontario.
The tourmalines are commonly coated by a F-rich rim of fluor-elbaite or fluor-schorl due to late-stage accumulation of F in the system. The presence of tourmaline indicates that the melt was enriched in B throughout the crystallisation history of the pegmatite body. Furthermore, the most evolved tourmaline in the pegmatite are classified as fluor-elbaite, the product of a common evolutionary trend seen in pegmatites (Selway et al., Reference Selway, Černý, Hawthorne and Novák2000a, Henry and Dutrow Reference Henry and Dutrow2011). The fluor-schorl rims of the quartz zone could be due to late-stage metasomatism of the existing schorl formed by previous country rock assimilation in the subsurface of the pegmatite adding Fe to the system (Henry and Dutrow, Reference Henry, Dutrow, Anovitz and Grew1996; Selway et al., Reference Selway, Černý, Hawthorne and Novák2000a). This could indicate late-stage enrichment of F in the melt due to gradual fractional crystallisation. The presence of country rock rafts within the pegmatite and surrounding tourmaline crystallisation indicates that Fe in the pegmatite is at least partially due to country-rock assimilation. It is hypothesised that at least some portion of the Mg required for unevolved tourmaline precipitation is sourced from country-rock assimilation and interactions (Selway et al., Reference Selway, Novák, Černý and Hawthorne2000b; Novák et al., Reference Novák, Prokop, Losos and Macek2017; Dyck and Larson Reference Dyck and Larson2023).
The Nb–Ta oxides show a geochemical evolution path from columbite-(Fe) in inclusions within cassiterites to columbite-(Mn) as isolated crystals, representing a progressively more evolved and fractionated geochemical system (Fig. 12). A disjointed fractionation trend is observed in Nb–Ta oxides within the central core zone and intermediate graphic zone reflecting an increase in Mn and Ta contents. Finally, the central zone distal tourmaline region shows a strong linear fractionation trend reflecting a dominant Mn occupancy and decreasing Nb content. These geochemical relationships reflect a late-stage accumulation and precipitation of oxides in the central zone. These are typical general fractionation trends documented for pegmatites globally (Černý Reference Černý1992; Martins et al., Reference Martins, Lima, Simmons, Falster and Noronha2011; Garate-Olave et al., Reference Garate-Olave, Roda-Robles, Gil-Crespo, Pesquera and Errandonea-Martin2020). The disjointed trend seen in Fig. 12 could be attributed to sampling bias, or alternatively, be due to the relatively high solubility of columbite-(Mn) in felsic melts accompanied by high levels of Li and F in the late-stage residual melt that increases solubility (Bartels et al., Reference Bartels, Holtz and Linnen2010). The variation in Mn/(Mn+Fe) could be linked to tourmaline crystallisation, which is known to cause rapid increases in Mn and Fe in the melt upon crystallisation (Van Lichtervelde et al., Reference Van Lichtervelde, Linnen, Salvi and Beziat2006).
In summary, in the Prof pegmatite, the compositions of the rock-forming minerals generally have an increase in Li, Mn and Ta/Nb and a decrease in Fe as the system evolved (Figs 9, 10, 12 and Tables 2, 3 and 4). Precipitation of petalite and elbaite depleted the melt in Li and B and subsequently, a late-stage F-rich fluid formed, which acted to precipitate rims on the existing tourmaline.
Textures
The graphic, comb and aplitic textures present in the Prof pegmatite are evidence that the system was undercooled upon crystallisation and that rapid crystallisation of a geochemically evolved, incompatible element-rich and flux-bearing melt formed this pegmatite.
Graphic intergrowths in pegmatites have been studied extensively (e.g. Fenn Reference Fenn1986, Lentz and Fowler, Reference Lentz and Fowler1992, Ikeda et al., Reference Ikeda, Nakano and Nakashima2000, Sirbescu et al., Reference Sirbescu, Schmidt, Veksler, Whittington and Wilke2017,). Graphic intergrowths form via undercooling in pegmatites generally in conjunction with the geochemical interactions of incompatible, fluxing elements in the residual melt (London Reference London2008, Simmons and Webber Reference Simmons and Webber2008, Nabelek et al., Reference Nabelek, Whittington and Sirbescu2010, Baker et al., Reference Garate-Olave, Roda-Robles, Gil-Crespo and Pesquera2018). These conditions lead to the co-precipitation of quartz and feldspar in a graphic texture. The prevalence of graphic intergrowths in the Prof pegmatite is consistent with undercooling during the pegmatite crystallisation. It should be noted, however, that Li enrichment has been experimentally demonstrated to have a significant effect on pegmatite texture, favouring undercooled structures at lesser degrees of undercooling (Maneta and Baker, Reference Maneta and Baker2014).
The tourmaline comb structures within the Prof pegmatite are a common pegmatite feature with well-developed examples in the Tanco pegmatite, Manitoba, Canada (Selway et al., Reference Selway, Novák, Černý and Hawthorne2000b) and at Minas Gerais, Brazil (Webber and Simmons, Reference Webber and Simmons2007). This texture is also an indication of rapid and undercooled melt crystallisation (Baker and Freda, Reference Baker and Freda1999).
Aplitic units are parallel with the country rock border zone at the Prof pegmatite. The crystallisation of aplite-bearing pegmatites has been shown to be associated with periods of rapid crystallisation (Webber et al., Reference Webber, Simmons, Falster and Foord1999) and undercooling (Baker et al., Reference Maneta and Baker2014Reference Maneta and Baker2014). A subtle mineralogical layering is present within the aplites showing variable abundances of tourmaline, quartz and feldspar. This layering can be attributed to a local boundary layer diffusion which causes geochemical variability along the crystallisation front and therefore mineralogy of the aplite (Webber et al., Reference Webber, Simmons, Falster and Foord1999; London Reference London2008, Reference London and Kontak2012). This boundary layer effect hypothesis is also supported by the high Cs content of mica also attributed to this effect (London, Reference London2022).
Relationship to other pegmatites in the region
It is hypothesised that the Prof dyke forms part of a much larger pegmatite field and this relatively small intrusion is one of the most evolved pegmatites in the field based on the similar mineralogy and strike of other intrusions on the mountain (Red, Green and Grail). The intrusive contacts between the pegmatites and country rock at the Prof pegmatite, and in the region, indicate that the pegmatites were emplaced into a variably brittle–ductile host rock. The country rock which the Prof pegmatite has intruded is undoubtedly the Monashee cover sequence calc-silicate gneiss. However, further to the east towards the Red, Green and Grail pegmatites, the host rock for the pegmatite field is more migmatitic in character and is interpreted to represent relatively deeper lithologies. This difference in depth could facilitate ductile emplacement in hotter deeper regions and brittle emplacement within the cover sequence calc-silicate rocks. An alternative hypothesis could also be argued that the difference in rheology of the host rock could affect the emplacement mechanisms, and this is the cause of variable pegmatite-host rock contacts. There are numerous pegmatites within the Boulder Mountain Pegmatite Group, which contains the Prof pegmatite, that have not been mapped and could also contain notable rare-element mineralisation.
The Mount Begbie pegmatite group is located on the mountains adjacent to Boulder Mountain and south of Revelstoke. This group was described by Dixon et al. (Reference Dixon, Cempírek and Groat2014) and contains a variety of barren to lithium mineralised pegmatites. These lithium-bearing pegmatites strike at 310° and dip vertically (Dixon et al., Reference Dixon, Cempírek and Groat2014). The Prof pegmatite strikes northeast at 060°, which is almost orthoganal to those at Begbie. The mineralogy described by Dixon et al. (Reference Dixon, Cempírek and Groat2014) has remarkably similar compositions to those of the Prof pegmatite, with a similar tourmaline assemblage, accessory minerals, and the identification of very minor petalite and qitianlingite in one of the Mount Begbie pegmatites. Although the Mount Begbie lithium-bearing pegmatites contain petalite, it is in very minor abundance and they were classified as lepidolite subtype by Dixon et al. (Reference Dixon, Cempírek and Groat2014).
Regardless of the difference in strike between the pegmatites at Mount Begbie and Boulder Mountain, their similar mineralogy begs the question of whether they are related. Numerous pegmatite dykes were recognised during aerial reconnaissance of the region between Mount Begbie and Boulder Mountain, however they have yet to be explored. This is in part due to difficulty accessing the area. Future work will involve mapping and sampling to determine the true extent (geographical, compositional and geochronological) of the pegmatites in the Revelstoke area.
Typically, geochemical evolution trends of pegmatites can be linked and used as prospective exploration vectors to find more pegmatites. If the Prof pegmatite is linked to the Mount Begbie pegmatites, then this indicates that the geochemical evolution trends of the system might increase to the northwest due to increasing contents of lithium-bearing minerals. As the region has not been mapped in detail, it is unknown whether there are larger, or more geochemically evolved, pegmatites in the region. More research is required to assess the extent and mineralisation potential of the pegmatites on Boulder Mountain.
Conclusions
(1) The Prof pegmatite contains minerals that are indicative of a highly geochemically evolved pegmatite including petalite, elbaite, trilithionite–polylithionite, Ta and Nb oxides and amblygonite. This is the first major lithium-bearing pegmatite identified in British Columbia, Canada.
(2) The significant abundance of macrocrystal petalite in the dyke indicates that it is a petalite-subtype or a Group one pegmatite.
(3) The geochemical evolution of the Prof pegmatite can be traced through the compositions of mica, tourmaline and Nb–Ta oxides. Trends in these data show the fractionation of elements and progressive crystallisation of phases indicative of highly evolved geochemical systems.
(4) The mineralogy generally shows an increase in Li, Mn and Ta/Nb and a decrease in Fe recorded in mica, Nb–Ta oxides and tourmaline as the geochemical system evolved. Precipitation of petalite and elbaite depleted the melt in Li and B and subsequently, a late-stage fluorine-rich fluid was exsolved, which precipitated minerals as rims on the existing tourmaline and mica of the deposit.
(5) The Prof pegmatite forms part of the extensive unmapped Boulder Mountain Pegmatite Group. Other lithium-bearing pegmatites found nearby, which host a similar mineralogy, are the Grail, Green and Red pegmatites. More work is required to assess the true extent of this group, and its rare element enrichment and mineral potential.
(6) Comparisons with other pegmatites of the region show that the Prof pegmatite has a similar mineralogy to the documented pegmatites on Mount Begbie.
(7) If the Boulder and Begbie pegmatite groups are linked, the pegmatites evolve to the northwest, therefore, detailed mapping of the area is required to assess whether there are larger and more-evolved pegmatites in the field.
Supplementary material
The supplementary material for this article can be found at https://doi.org/10.1180/mgm.2024.49.
Acknowledgements
This paper is dedicated to the memory of Dr. Alessandro Guastoni, curator of the Museum of Mineralogy at University of Padova, an ardent pegmatologist and a wonderful person. He will be missed. The authors thank Tiera Naber and Tommi Muilu for invaluable field assistance, and Eric Johnstone of Glacier Helicopters Ltd. for helicopter support. Thanks are extended to Jan Cempírek, David Lentz and three anonymous reviewers who helped improve the manuscript. The authors also thank Anette von der Handt for EPMA assistance, Jacob Kabel for SEM assistance, Anita Lam for powder XRD assistance and Mark Button for LA-ICP-MS assistance. Thanks are also extended to Elizabeth Ye, Maya Saldanha and Pearl Bains for their help and enthusiasm. This research was supported by a GSA Lipman Research Award and a SEG Canada Foundation Research Grant to CMB. The authors acknowledge the support of the Natural Sciences and Engineering Research Council of Canada (NSERC) through a Discovery Grant (funding reference 06434) to LAG.
Competing interests
The authors declare none.
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