Geochemistry and mineralogy of the shale-hosted vanadium Van Property deposit, Mackenzie Mountains, Northwest Territories, Canada Open Access

The demand for critical metals has significantly increased over the last decade due to advancements in the environmental needs of low-carbon energy technologies [e.g., the development of electric car batteries for mass production, batteries for energy storage (Schulz et al. 2017; Jowitt et al. 2018; Simandl et al. 2023)]. Many governments and agencies have published lists of critical metals (Schulz et al. 2017; European Commission 2020; Natural Resources Canada 2020; Su and Hu 2022; USGS 2022), highlighting both the economic importance and vulnerability in the supply chain of critical metals. Among these critical metals, the element vanadium (V) has yet to be in the spotlight of both research and economy despite its increasing importance, especially in the redox-flux batteries used in solar cells for energy storage (Kelley et al. 2017; Jowitt et al. 2018; Elbokl 2022; Simandl and Paradis 2022).

Most of the global V production is currently mined as a by-product from magmatic vanadiferous titanomagnetite deposits [e.g., Bushveld complex: Maier et al. (2012), Viljoen (2016), Boni et al. (2023); Lac Dore, Quebec, Canada: Fischer (1975), Longridge and Martinez (2020); Panzhihua layered intrusion, China: Zhou et al. (2005)] and to a lesser extent from sedimentary sandstone-hosted deposits that are enriched in other critical elements such as uranium [e.g., Colorado Plateau; Slick Rock district, Colorado: Shawe (2011); Pandora deposit, Utah: Fischer (1968)]. The third important type of V deposit is the shale-hosted V deposits [e.g., Green Giant, Madagascar: Di Cecco et al. (2018); several deposits in Mozambique: Boni et al. (2023) and references therein]; however, no currently active shale deposits are being mined for V, despite the relatively large abundance of black shales and mudstones on and near the surface. The reasons for this are economic (e.g., lower average V content compared to titano-magneto-ferrous deposits and relatively low V prices currently) and geological, such as the lack of exploration in sedimentary basins specifically for V, ignorance of V-enrichment processes in black shales, and beneficiation of black shales at profitable recovery rates for V (Vitolo et al. 2000; Li et al. 2021; Simandl and Paradis 2022). Closing the geological knowledge gap by developing a holistic deposit model that addresses both the mineralogy of shale-hosted V deposits and their formation is important to support exploration efforts and increase chances of success.

This study focuses on the metamorphosed Van Property in the Mackenzie Mountains, Northwest Territories, and utilizes lithogeochemistry and mineral compositions to describe how host lithology and mineralogy impact V enrichment and the role that metamorphism plays in epigenetic V remobilization. Detailed scanning electron microscopy, electron microprobe analyses, and X-ray diffraction focused on V-bearing phases highlight that the enrichment of V occurs syngenetically with metamorphic alteration of the V host phases. The chemical composition of the Van Property is compared to other Paleozoic shale-hosted V deposits, highlighting variations in host lithology, geochemistry, and V-hosting mineralogy among this deposit type. The results of this study contribute to filling knowledge gaps within this yet understudied deposit class and provide a better understanding of the compositional and mineralogical variety among shale-hosted V deposits, which has important implications for V-enrichment processes. Hence, results from this study have direct impacts on our understanding of ore-forming processes in shale-hosted V deposits and their exploration.

The Mackenzie Mountains comprise the western part of the Mackenzie Platform and the eastern part of the Selwyn Mountains; they are part of Ancestral North America, located on the western part of Laurentia (Nelson et al. 2013). There, they form the most northeastern extension of the Cordillera (Gordey and Roots 2011; Ootes et al. 2013; Fig. 1). Stratigraphically, rocks from the Neo-Proterozoic to the Cretaceous are preserved in the Mackenzie Mountains. The Neo-Proterozoic strata include the older Mackenzie Mountains Supergroup (<1.05 and >0.78 Ga; Turner 2010; Gordey and Roots 2011; Long and Turner 2012) and the younger Windermere Supergroup (≈755 to 550 Ma; Narbonne and Aitken 1995; Gordey and Roots 2011; Ootes et al. 2013). The former consists of carbonate and clastic strata of 4 to 6 km thickness formed in deltaic, fluvial, and shelf settings of a large epicratonic basin (Rainbird et al. 1996; Long et al. 2008; Turner and Long 2008). After the breakup of Rodinia, sedimentation on Laurentia’s western margin started as early as 755 Ma, forming the carbonate, shallow to deep-marine clastic, and glacially influenced clastic strata of the Windermere Supergroup, which is 5 to 7 km thick and deposited in a rift-related setting (Narbonne and Aitken 1995; Gordey and Roots 2011; Ootes et al. 2013). Sedimentation continued into the Paleozoic with shallow-water carbonates deposited on the Mackenzie Platform to the northeast and coeval deeper-water marine strata deposited in the Selwyn Basin to the southwest (Fig. 1a and 1b). Although no structurally controlled boundary on the platform margins is evident on the surface, deeper crustal features may have contributed to subsidence and the different depositional environments (Cecile 1982; Gordey and Roots 2011).

Paleozoic strata of the carbonate platform include basal Ediacaran to Cambrian shallow-marine to fluvial quartz sandstone and minor carbonate rocks that are overlain by Cambrian to Devonian formations dominated by dolostone with minor limestone and siliciclastic rocks; open-marine Devonian carbonate rocks complete the succession (e.g., Gordey and Roots 2011; Turner et al. 2011; Ootes et al. 2013). Coeval Paleozoic strata of the Selwyn Basin are dominated by fine-grained deep-water marine sediments. The simplified stratigraphy of the south-central Selwyn Basin is shown in Figure 2. The Vampire, Sekwi, and Rabbitkettle formations, and the Road River Group are all exposed in the study area (Fig. 3); more detailed descriptions of these units are provided below.

During the Late Devonian, depositional regimes changed in the Mackenzie Mountains siliciclastic dominated the strata making up the Lower Devonian to Middle Mississippian Earn Group. Siliceous mudstones and carbonaceous cherts of this unit are present both in the south-central Selwyn Basin and study area (Figs. 2 and 3) and received their detritus from western sources that included elevated fault blocks of older Selwyn Basin strata (Gordey 1988; Gorey and Anderson 1993). Although sedimentation continued during the Early Carboniferous and Permian, the stratigraphic record is fragmented throughout the Mackenzie Mountains and absent in the study area. Cretaceous shale and sandstone, pebble and cobble conglomerate, and localized coal streams are restricted to the Mackenzie Platform in the northern Mackenzie Mountains (Fig. 1b).

Mid-Cretaceous plutonism is common throughout the Mackenzie Mountains (Gordey and Roots 2011; Ootes et al. 2013). However, plutons intruded in Selwyn Basin strata are typically of Jurassic age and are composed of alkaline, calc-alkaline, and two-mica granite (Woodsworth et al. 1991; Gordey and Roots 2011).

Strata of the Selwyn Basin underwent low-grade metamorphism of sub-greenschist to greenschist facies during the Jurassic in relation to the formation of the Cordillera in the west (Gordey and Roots 2011). In the aureoles of Cretaceous plutons, contact metamorphism is reported throughout the Mackenzie Mountains (Gordey and Anderson 1993). Dominant faults in the Mackenzie Mountains are northwest-southeast trending (Fig. 1b) and are either syn-sedimentary extensional faults developed dominantly during the Paleozoic (Aitken and Cook 1974; Eisbacher 1981; Turner and Long 2008; Pyle and Jones 2009; Gordey et al. 2011) or Cretaceous-Tertiary folds and thrust faults (Morris and Nesbitt 1998; Mair et al. 2006; Pyle and Jones 2009; Gordey and Roots 2011).

Mineralization is common in the Mackenzie Platform and Selwyn Basin strata of the Mackenzie Mountains (Fig. 1b) with over 300 known mineral occurrences (Ootes et al. 2013). Most of these occurrences are clastic-dominated (a.k.a., sedimentary exhalative) Zn-Pb deposits in the Selwyn Basin (e.g., Howard’s Pass) or carbonate-hosted Zn-Pb deposits (e.g., Prairie Creek, Gayna River) in the Mackenzie Platform; these deposits formed during the Paleozoic (Fig. 2; (Ootes et al. 2013). Tungsten-skarn deposits (e.g., Cantung and Mactung deposits) related to Cretaceous intrusions occur in the Selwyn Basin close to the Van Property study area (Fig. 1b).

The Van Property is located in the Flat Lakes area, southwestern Northwest Territories (Fig. 1). Stratigraphically, the Van Property occurs in Paleozoic strata of the Selwyn Basin, which form a large-scale, northwest-trending, upright syncline with near vertical limbs; no large-scale regional faults are observed (Carne and Gish 1999). At the Van Property, Upper Proterozoic to Middle Mississippian strata occur (Fig. 2), which are summarized here after Carne and Gish (1999) and Flavelle (2013).

The oldest unit, Upper Proterozoic to Lower Cambrian Vampire Formation, comprises gray to brown siltstone and green phyllite with local limestone interbeds and is conformably overlain by the Lower Cambrian Sekwi Formation. At the Van Property, the latter only occurs in the northeast (Fig. 2) and consists of dolostone and distinctively orange-weathered limestone. Both lithologies are common throughout the Selwyn Basin, with reported thicknesses of 390 and 700 m for the Vampire and Sekwi formations, respectively (Gordey and Roots 2011). However, the thickness of each formation is not reported from the Van Property. The Upper Cambrian to Lower Ordovician Rabbitkettle Formation sharply overlays the Sekwi Formation and is the most visually distinctive unit in the study area due to the differential weathering of more competent fine-grained dolomite beds with less-weather-resistant, silty limestone lenses. At the Van Property, this unit is 65 to 75 m thick, much less than the reported thickness of 435 to 840 m in other parts of the Selwyn Basin (Gordey and Roots 2011). Carne and Gish (1999) interpreted the sharp transition between the Rabbitkettle Formation and overlying Road River Group as an unconformable or a disconformable contact. None of the afore-described units (Vampire, Sekwi, Rabbitkettle formations) show mineralization, including V, in the study area.

The Lower Ordovician to Devonian Road River Group is widespread in the Selwyn Basin, although its nomenclature is problematic (Cecile 1982; Gordey and Roots 2011; Turner et al. 2011). At the Van Property, the Road River Group is represented by the shale-dominated, recessive Duo Lake Formation (OSD). In the study area, the OSD, which apparently formed in a deep-marine setting below the carbonate compensation depth (Turner et al. 2011), is divided into five members (Fig. 2). The bottom unit, the Transition Member, comprises a thin-bedded limestone and phyllite sequence that becomes more pelitic and carbonaceous up section, grading into dark gray, non-calcareous phyllite and phyllitic siltstone toward the top. This unit is overlain by the Siliceous Mudstone Member consisting of up to 50 m of pyritic, black siliceous mudstone in which pyrite is present as thin laminae, ovoid concretions, and disseminations. In sharp contact above occurs finely to thickly bedded gray to black, graptolite-bearing, recessive carbonaceous limestone and calcareous mudstone of the Calcareous Mudstone Member, which can reach a thickness of several hundred meters and has a sharp contact to the overlying unit. The Lower Cherty Mudstone Member is the primary V-bearing horizon at the Van Property, with reported grades of about 0.6% V₂O₅ over 50 m vertically and 700 m along strike. This unit is characterized by blocky fracturing, gray-black to black weathering, and uniform bedding of 3 to 15 cm occurring in cherty mudstone or siliceous argillite with an elevated siliceous content. Despite the relatively high silica content, the carbon content is higher than in the adjacent and, especially, underlying units. On weathered surfaces, secondary blue zinc and yellow vanadium minerals are reported. The unit is up to 100 m thick at the Van Property. It sharply changes into the last member of the OSD, the Upper Siliceous Mudstone Member, consisting of poorly laminated, well-cleaved, carbonaceous, and siliceous mudstone with abundant limestone concretions and can exceed 600 m thickness. The top of the unit is graptolite-bearing, constraining the age to Lower Devonian. An anomalous V-rich horizon of up to 21.5 m at 0.6% V₂O₅ occurs at the top of the unit. The contact with the overlying Earn Group is gradational.

The Lower Devonian to Middle Mississippian Earn Group, a turbiditic sequence, comprises the Portrait Lake and Prevost formations at the Van Property, of which the latter is the youngest unit preserved in the study area. The Portrait Lake Formation, 70 to 100 m thick, consists of carbonaceous chert and cherty argillite with limestone concretions and does not show cleavage. The Prevost Formation has non-siliceous to siliceous distal turbidites with a basal cross-laminated to parallel-laminated siltstone base that grades upward into a thicker, fine-grained argillite. Barite is common, and cleavage is well developed in the unit.

Although the study area did not undergo excessive deformation, relatively low-grade (sub-)greenschist metamorphism occurred during the Jurassic (e.g., Gordey and Roots 2011). The recent study by McGill et al. (2024) revealed metamorphic temperatures of 345–361 °C for the Van Property. Additionally, cleavage is common in all Paleozoic units at the Van Property. However, the Rabbitkettle Formation has two cleavages and small-scale open folds with east-trending axes, which are absent in the overlying Road River Group. This led to the assumption that the area underwent a period of deformation following the deposition of the Rabbitkettle Formation prior to the onset of Road River Group deposition (Carne and Gish 1999). Hence, more than one deformation event occurred.

A total of six samples were collected in the summer of 2005 at the Van Property by geologists of the Northwest Territories Geological Survey to characterize the composition and mineralogy of the prospect, with special attention paid to V-enriched horizons (Fig. 3; Table 1). Samples are from three stratigraphic levels listed from oldest to youngest: (1) Lower Cherty Mudstone Member, OSD (hf05-33 to -36; Figs. 3, 4a, and 4b); (2) Upper Siliceous Mudstone Member, OSD (hf05-37; Figs. 3 and 4c); and (3) Prevost Formation, Earn Group (hf05-38; Figs. 3 and 4d).

Various methods were applied to constrain the composition and mineralogy of the V-bearing siliceous argillite and siliceous mudstone at the Van Property. Details on whole-rock lithogeochemistry, X-ray diffraction (XRD), scanning electron microscopy (SEM), and electron microprobe analysis (EMPA) can be found in the Online Materials¹.

The siliceous argillite samples of the OSD are very fine-grained (<20 μm) with varying degrees of thickness and carbon content; the latter is especially prominent in samples with a sooty appearance. Some samples are cross-cut by millimeter- to centimeter-thick calciteveinlets (Fig. 4a). Yellow (e.g., secondary V minerals; Carne and Gish 1999) to brown weathering (i.e., gossan made of apparently secondary goethite) is common (Figs. 4b and 4c). The sample from the Prevost Formation (hf05-38) is gray, fine-grained with millimeter-thick barite interbeds, yellow weathered surfaces, and prominent cleavage (Fig. 4d).

The samples from the OSD were examined in more detail due to the V enrichment reported for these strata in Carne and Gish (1999) and Flavelle (2013), and the V concentration was determined by whole-rock analysis (see below). Their major mineral is quartz with minor to trace amounts of carbonaceous matter and traces of muscovite, illite, apatite, rutile, calcite, dolomite, and sulfides (i.e., pyrite, sphalerite; Fig. 5). Fine, aligned quartz schlieren with trace amounts of silicates alternate with black carbonaceous material (Fig. 5a), whereas carbonates, sulfides, and rutile are commonly disseminated in the quartz-rich matrix with no preferred orientation (Figs. 5b–5h).

Quartz (Figs. 5c and 5e) is the dominant phase and occurs as very fine-grained (<10 μm), sub-angular to sub-rounded grains of which the long axis is commonly aligned parallel to bedding. Convex-concave grain boundaries and triple junctions of ≈120° between adjacent grains (Fig. 5h) are indicative of pressure solution and recrystallization at (sub-)greenschist facies metamorphism (e.g., Becker 1995; Wenk et al. 2022). Carbonaceous matter (Figs. 5b–5h) is the second most abundant phase in the OSD samples. It is very fine-grained (<20 μm), flaky with amoeboid shapes, and is commonly found in occurrences between quartz grains and adjacent to illite, rutile, and sulfides. Although X-ray diffraction identified graphite, based on reflectance and the assumed P-T conditions [i.e., (sub-)greenschist facies], additional carbonaceous phases could not be unambiguously identified but are assumed to be present. Illite (Figs. 5c–5e) occurs as prismatic single grains within the quartz matrix and is commonly at an obtuse angle to the long axis of quartz grains; instead, it shows alignment parallel to cleavage than to bedding. Grains are up to 30 μm long, can show internal cleavage planes, can be kinked due to deformation, and rarely have minute sulfide inclusions. Illite is in contact with carbonaceous matter on at least one surface and can occur spatially close to sulfides and rutile. Rutile (Figs. 5c and 5e–5h) is amoeboid, <10 μm in size, and occurs most commonly in direct contact with either pyrite or sphalerite, forming rutile-sulfide aggregates; rutile can have minute, round sulfide inclusions. Rutile-sulfide aggregates have no preferred orientation, are randomly disseminated, and occur adjacent to carbonaceous matter. Disseminated rutile occurs rarely as isolated grains within the quartz-rich matrix. Pyrite and sphalerite are of similar abundance and occur as disseminated grains within the quartz-rich matrix with no preferred orientation. Pyrite (Figs. 5c–5f) is commonly amoeboid in shape, less commonly prismatic, and rarely framboidal. Sphalerite (Figs. 5c, 5g, and 5h) has amoeboid shapes, is intergrown with rutile, and forms in direct contact with carbonaceous matter. Other trace phases such as muscovite, calcite, dolomite, and apatite (Figs. 5c and 5e–5g) occur as very small (<15 μm), disseminated, subhedral to anhedral, single grains in the quartz-rich matrix. Calcite and dolomite are also observed in close contact with rutile-sulfide aggregates.

Deformation is most prominently developed in recrystallized quartz grains and kinked illite (Figs. 5e and 5h). Pyrite and sphalerite can show cataclastic cracks formed under brittle conditions.

Results from major, minor, and trace element composition of siliceous argillite (OSD) and siliceous mudstone (Prevost Formation) at the Van Property are summarized in Tables 2 and 3, and the complete data set is found in the Online Materials¹.

The compositions of the analyzed samples differ between those from the OSD (hf05-33 to -37) and the younger Prevost Formation (hf05-38). Although samples from the OSD are from different stratigraphic members (i.e., hf05-37: Upper Siliceous Mudstone Member; hf05-33 to -36: Lower Cherty Mudstone Member), their major to trace element compositions are similar (Online Materials¹).

Using the chemical classification of Herron (1988) to distinguish between shales and sands by utilizing log ratios of SiO₂/Al₂O₃ and Fe₂O₃t/K₂O, samples from the OSD cannot be defined as (black) shales (Fig. 6a). Instead, they are referred to as siliceous argillite here due to their relatively high SiO₂ content and log(SiO₂/Al₂O₃) ratio (Fig. 6a). This is also in accordance with the nomenclature used in the previous study at the Van Property by Carne and Gish (1999). Additionally, these siliceous argillites vary in CaO concentration (Herron 1988), with most samples being non-calcareous (CaO < 2.8 wt%) and only one sample classifying as calcareous (CaO = 2.8–10.7 wt%, Fig. 6b). The sample from the Prevost Formation classifies as non-calcareous siliceous mudstone (Fig. 6). In the OSD siliceous argillites, the Al₂O₃ (2.32 ± 0.52 wt%), K₂O (0.68 ± 0.14 wt%), Na₂O (0.03 ± 0.01 wt%), TiO₂ (0.12 ± 0.03 wt%), and BaO (0.05 ± 0.02 wt%) contents are lower relative to the younger siliceous mudstone (Figs. 7a–7b; Table 2). In contrast, SiO₂ (80.6 ± 9.17 wt%), Fe₂O₃t (1.03 ± 0.25 wt%), total organic carbon (TOC; 7.9 ± 3.2 wt%), and S (0.82 ± 0.49 wt%) contents are higher relative to the siliceous mudstone from the younger Prevost Formation (Figs. 7c–7d; Table 2).

The trace element concentrations between the different stratigraphic horizons at the Van Property also vary (Figs. 7e–7h and 8; Table 3). Samples from the OSD show enrichments in some metals exceeding twice the concentration of SDO-1, the reference black shale of the U.S. Geological Survey (Huyck 1991; Jochum et al. 2005), defining the investigated samples as metalliferous following Huyck (1991). In contrast, the siliceous mudstone of the Prevost Formation is non-metalliferous. In particular, V and Zn are strongly enriched in the OSD with average concentrations of 2943 ± 978 ppm and 3665 ± 3637 ppm, respectively, compared to 675 ppm V and 19 ppm Zn in the Prevost Formation. Furthermore, the siliceous argillites of the OSD have higher contents of U (8.15 ± 2.81 ppm vs. 3.29 ppm, Fig. 7f), Ni (160 ± 108 ppm vs. 3 ppm, Fig. 7g), Mo (41.7 ± 19.7 ppm vs. 2 ppm, Fig. 7h), Cu (103 ± 108 ppm vs. 2 ppm, Fig. 8), and As (45.2 ± 22 ppm vs. 2 ppm). In contrast, concentrations of Cr, W, and Co are similar between lithologies of both stratigraphic levels (Table 3). The (non-)calcareous siliceous argillites from the OSD show weak-to-strong positive correlations of V with various elements, including Al₂O₃ (R² = 0.74, Fig. 8a), K₂O (R² = 0.62, Fig. 8b), TiO₂ (R² = 0.53, Fig. 8c), S (R² = 0.89, Fig. 8d), TOC (R² = 0.85, Fig. 8e), U (R² = 0.58, Fig. 8f), Ni (R² = 0.92, Fig. 8g), Mo (R² = 0.54, Fig. 8h), Zn (R² = 0.45, Fig. 8i), and Cu (R² = 0.28, Fig. 8j).

Rare earth elements + Y (REE+Y) have relatively low total concentrations in both the OSD and Prevost Formation with sums of 83.7 ± 32.7 ppm and 110 ppm, respectively (Table 3). In samples from the OSD, the REE+Y pattern normalized to Post-Archean Australian Shale (PAAS; McLennan 1989) shows a strong Ce anomaly [Ce/Ce* = 0.49 ± 0.1 and 0.44 ± 0.1 after Lawrence et al. (2006) and Nozaki (2008), respectively] and no Eu anomaly [Eu/Eu* = 1.05 ± 0.1 after McLennan (1989); Fig. 9; Table 3]. The REE pattern increases slightly from light to heavy REE and Y. The REE pattern of the Prevost Formation sample is characterized by neither a significant Ce [Ce/Ce* = 0.93 and 0.9 after Lawrence et al. (2006) and Nozaki (2008), respectively] nor Eu anomaly [Eu/Eu* = 0.93 after McLennan (1989)], with a relatively flat pattern that shows a weak decrease from light to heavy REEs and Y (Fig. 9).

Results from XRD are provided in Table 4. Although samples from both the OSD and Prevost Formation are dominated by quartz (OSD: 82.1%, Prevost: 73.4%), their mineralogy varies (Fig. 10), consistent with their differing geochemistry. In the OSD samples, the second most common phase is carbonaceous matter (identified as graphite), followed by muscovite. However, based on reflectance studies using reflected light microscopy at the University of Manitoba, kerogen or bitumen likely are present as well. Hydrocarbon speciation done by SGS Canada Inc. did not reveal the composition of carbonaceous matter. The results from both methods were inconclusive. Additional mineral phases that occur commonly in the siliceous argillite, albeit at trace (<2%) levels, are illite, pyrite, sphalerite, calcite, dolomite, fluor-apatite, and rutile (Fig. 5). Most illite is of the 2M₁ polytype based on XRD spectra with the presence of characteristic peaks at d = 4.48 Å (021) and 2.42 Å (117). The c parameter of unit cells of the best-fitting illite also is ∼20 Å, confirming the presence of the 2M polytype [personal communication A. Zhang (2023)]. Biotite, gypsum, and goethite are not observed microscopically but by XRD, totaling <2% on average. The most prominent differences between siliceous argillite (OSD) and siliceous mudstone (Prevost Formation) are the enrichment in muscovite and the occurrence of barite, gibbsite, and phlogopite in the latter. Moreover, sulfides are absent, and graphite is a trace phase in the Prevost Formation.

Mineral phases with V as a major component reported from other shale-hosted V deposits [e.g., roscoelite and mannardite (Fu et al. 2023; Yang et al. 2023)] and Jim and Janice Creeks near the study location, NWT (Carne and Gish 1999), were not identified by XRD or SEM; V is only present in minor or trace concentrations in some identified phases. Secondary V-bearing phases (e.g., corvusite) or V-bearing sulfides (e.g., sulvanite), as reported by Carne and Gish (1999), were not identified via XRD or SEM in the studied samples. Common minerals hosting minor V in shale deposits are illite, oxides, and potentially carbonaceous matter (Fester et al. 1927; Fischer et al. 1968; Peacor et al. 2000; Di Cecco et al. 2018; Lu et al. 2021; Veselovský et al. 2021). To further constrain the V host(s) in siliceous argillite at the Van Property, SEM and EMPA were applied to the V-rich samples of the OSD.

All phases of interest (i.e., illite, rutile, carbonaceous matter) are very fine-grained (<30 μm), and detailed elemental maps done both via energy-dispersive spectrometry (EDS) on SEM and wavelength-dispersive spectrometry (WDS) on EMPA were utilized to determine the spatial distribution of V (Fig. 11). The most prominent V counts are observed both in rutile and some illite (Fig. 11) that have been in close contact with carbonaceous matter; sulfides occur with V-bearing rutile as well but do not show any V counts (Figs. 11a and 11h). In general, V signals recorded by EDS are strongest in rutile (Figs. 11b and 11i); the presence of V in this oxide was confirmed by observing spectra on the Kβ peak that do not interfere with Ti emission lines. Vanadium signals were recorded in carbonaceous matter (Figs. 11b and 11i) by EDS; however, the signal is weak and may represent background. It was not possible to confirm the occurrence of V in this phase, most likely due to relatively low abundance and hardness contrast to neighboring phases impacting the polishing of carbonaceous matter and hence its analyses by X-rays. Since neither mapping method allowed for quantitative assessment of V in illite and rutile, EMPA was utilized on selected grains in siliceous argillite.

Electron microprobe analyses were utilized to quantify the V content in both illite and rutile identified as V-bearing by both EDS and WDS mapping. The results for both phases are provided in Table 5 and the Online Materials¹.

Based on V-content, two illite compositions are identified: (1) low-V illite with <2 wt% V₂O₃ (Figs. 11a–11c and 11h–11j) and (2) high-V illite with >7 wt% V₂O₃ (Figs. 11e and 11f) of which the former is more common. Both variations are, however, similar in texture (Fig. 11), with the most prominent difference being that low-V illite commonly has cleavage developed and can be kinked (Figs. 11a–11c and 11h–11j). Both low-V and high-V illite are in contact with carbonaceous matter (Fig. 11). The composition of illite based on interlayer occupancy [i.e., illite-smectite with K = 0.5–0.69 atoms per formula unit (apfu) after Aja 1989 and Rosenberg 2002; illite end-member with K = 0.88 apfu after Meunier and Velde (1989); Środoń et al. (1992); Yates and Rosenberg (1997); Aja (2020)] was further constrained using the ternary plot after Newman and Brown (1987) and Aja (2020) (Fig. 12). This plot employs the silica content in the tetrahedral position (4Si), the sum of interlayer cations (M⁺), and the sum of mafic elements on the octahedral site (3R²) to distinguish between different clay end-members (Fig. 12a). Low-V illite is close to end-member composition of illite (K = 0.88 apfu; Fig. 12b) with a calculated average (n = 23) formula $K0.82(Al1.69Fe0.14Mg0.18Ti0.024+V0.023+)2.04vi(Si3.30Al0.70)4.00ivO10(OH)2·(H2O)$ (Table 5; Online Materials¹). In contrast, high-V illite is much closer to the illite-smectite composition (K = 0.50–0.69 apfu) and has strong deficits on the interlayer site (Fig. 12b) with an average (n = 7) formula of $K0.51(Al1.29Fe0.01Mg0.24Ti0.024+V0.563+)2.11vi(Si3.38Al0.62)4.00ivO10(OH)2·(H2O)$ (Table 5; Online Materials¹). Besides V content, low-V and high-V illite also have significantly different concentrations of Al₂O₃, FeO, K₂O, and Na₂O; low-V illite is enriched in all these elements relative to high-V illite (Fig. 13; Table 5). Vanadium also shows different element associations in each type of illite. In low-V illite, V is <0.17 apfu and is negatively correlated with Al³⁺ in the octahedral site; other correlations of V with elements in different sites are not observed (Figs. 13a–13e). In contrast, in high-V illite, the V correlates with several elements (Figs. 13a–13c and 13e), including Si and Al in the tetrahedral site, Al on the octahedral site, and Ti. When the size of the illite grains permitted, analyses close to the rim and core were performed, but no significant differences in composition were observed (Figs. 11 and 13; Online Materials¹).

The small size of rutile grains made it difficult to avoid analyzing neighboring phases. Of the nine spots that were initially analyzed, only six analyses with totals of >97.3 wt% were used. Although these totals are relatively low for oxides and rutile, analyses showed trace concentrations of SiO₂ (1.00 ± 0.25 wt%) and CaO (0.35 ± 0.07 wt%), most likely due to overlaps with adjacent quartz and/or calcite, the composition of the analyzed rutile grains was stoichiometrically homogeneous with an average (n = 6) formula of (Ti_2.93V_0.03)_2.91. Deficiencies in Ti⁴⁺ were attributed to the minute grain size and contamination with SiO₂ and CaO from neighboring grains. After TiO₂ (93.8 ± 2.22 wt%), V₂O₃ (1.66 ± 1.75 wt%) was the second most abundant element in rutile. Similar to illite, V concentrations in rutile were variable, with some spots having less than 0.80 wt% V₂O₃ (Figs. 11b and 11d) and other spots with ≥3.48 wt% V₂O₃ (Figs. 11i and 11k).

Shale-hosted V deposits are one of four major sub-classes of V deposits [i.e., vanadiferous titanomagnetite deposits, sandstone-hosted deposits, shale-hosted deposits, and vanadate deposits, e.g., Fischer and Stewart (1961); Fischer (1968, 1975); Fischer et al. (1968); Kelley et al. (2017)]. None of the approximately 20 globally known shale-hosted V deposits (Kelley et al. 2017; Elbokl 2022; Fu et al. 2023; Yang et al. 2023) are currently in production. Reasons include current limits on the effective beneficiation of V due to the fine-grained nature of host lithology and potential occurrence of V with organic matter (Vitolo et al. 2000; Li et al. 2021; Simandl and Paradis 2022) as well as a poorly developed deposit model. The deposit model does not account for variations in host lithology, V-bearing mineralogy, type and maturity of carbonaceous matter (i.e., kerogen, bitumen, or graphite), metamorphic grade, or what these characteristics imply for V enrichment process(es).

To highlight variations in shale-hosted V deposits and contribute to a more holistic deposit model, the geochemistry of the Van Property is compared to other shale-hosted V deposits (Figs. 6 to 9), which were selected based on V grade (>0.15 V₂O₅ wt%) and their reported major, minor, and trace element composition (Tables 2 and 3). The selected deposits are, from youngest to oldest: (1) Mecca Quarry Shale from Indiana and Illinois, U.S.A. (Coveney and Martin 1983; Peacor et al. 2000); (2) historical data from the Duo Lake Formation at the Van Property, Flat Lakes, NWT, Canada (Carne and Gish 1999); (3) Moss prospect, Yukon, Canada (Gadd et al. 2019); (4) Storsjøn-Myrvikken in Jämtland, Sweden (Leventhal 1991); and (5) Zhongcun deposit, southern China (Fu et al. 2023). All these deposits are of Paleozoic age (early Cambrian to Pennsylvanian) and vary in host lithology, metamorphic grade, TOC and S concentrations, V-hosting minerals, and the reported process(es) responsible for V enrichment (Table 6 with listed references). Some of these features (e.g., age) are common for most known shale-hosted V deposits. In contrast, other characteristics (e.g., host lithology, content of TOC and S content, mineralogy) are much more variable than summarized in Kelley et al. (2017).

Within the selected deposits, host lithology is not restricted to black shales, marls, or mudstones (Kelley et al. 2017) and rather includes silica-rich, very fine-grained sedimentary rocks with a minor clay component such as the siliceous argillites observed at both the Van Property and Zhongcun deposits (Fig. 6a). Therefore, deposits show similar or higher than average V grades of the black shales and mudstones from the Mecca Quarry Shale, Moss prospect, and Storsjøn-Myrvikken (Fig. 7e; Tables 2–3). Hence, clay minerals, whose abundance exceeds 55% in shales (Shaw and Weaver 1965), are not solely responsible for scavenging and retaining V, although authigenic illite is commonly referred to as the dominant V host in shale-hosted V deposits (Coveney and Martin 1983; Meunier 1994; Peacor et al. 2000). Therefore, it seems, phases other than illite can contribute to the V budget of shale-hosted V deposits. Recent studies from Fu et al. (2023) and Yang et al. (2023) showed that V is also hosted in hydrothermally formed, V-rich micas (roscoelite, mannardite), anatase, and secondary goethite.

Most shale-hosted V deposits have average TOC contents exceeding 5 wt% (Kelley et al. 2017), with the exception of Moss (2.14 ± 1.69 wt%; Fig. 7c; Table 2). Although there is a positive linear correlation between TOC and V (Fig. 8e), suggesting organic matter plays a critical role in V enrichment (Lewan and Maynard 1982; Breit and Wanty 1991; Leventhal 1991), a direct relationship between high-TOC content and V enrichment is ambiguous because organic-rich shales such as the reference Devonian Ohio shale SDO-1 (Fig. 8e; Table 2) has relatively low metal, especially V, concentrations. Lu et al. (2021) showed that dissolved organic matter, V availability in seawater, and the redox state of bottom waters control V enrichment over time since V concentrations varied in stratigraphic horizons of the lower Cambrian Nuititang Formation, South China. Hence, the occurrence of shale-hosted V deposits may not only be a function of TOC content but also how much V was dissolved in seawater at any given time in the geologic record.

The relationship between reduced S and V in shale-hosted V deposits is even less straightforward (Fig. 8d); S exceeds 1 wt% in most shale-hosted V deposits (Kelley et al. 2017). However, both the Van Property and Zhongcun deposits have S < 1 wt% (Table 2). It has long been assumed that H₂S is an effective reductant in shales converting oxidized, mobile V⁵⁺ to reduced, less mobile V⁴⁺ or V³⁺ (Breit and Wanty 1991; Wanty and Goldhaber 1992). However, the lack of correlation between S and the occurrence of V enrichment in deposits such as the Van Property with relatively low S indicates: (1) S does not play a critical role in V enrichment in shale-hosted deposits with S < 1 wt%, and (2) S is most likely introduced into shale-hosted deposits independently from V, for instance by an oxidized brine similar to fluids forming sedimentary exhalative deposits (e.g., Emsbo et al. 2016; Goodfellow and Lydon 2007). Although bacteria in mudstones and shales are common S producers via bacterial sulfate reduction of seawater (e.g., Canfield 2001) forming framboidal pyrite, this S source is interpreted to be minimal at the Van Property. Here, prismatic pyrite and sphalerite are not authigenic (Fig. 14) and show evidence that they were formed by a hydrothermal fluid that encountered the siliceous argillite of the OSD after diagenesis since sulfides have random orientation, occur with rutile (also secondary; Fig. 14), and overgrow the quartz-rich matrix (Figs. 5 and 11).

The positive correlations of V with other metals such as Mo, Ni, Zn, and Cu (Figs. 8e and 8g–8j) emphasize that V enrichment could be related to several processes: (1) syngenetic scavenging from seawater and enrichment in authigenic material coevally with Mo (e.g., Algeo and Maynard 2004; Tribovillard et al. 2012) and (2) (epigenetic) enrichment involving hydrothermal brines that also add Zn and/or Cu (e.g., Goodfellow and Lydon 2007; Hitzman et al. 2012; Brown 2014; Emsbo et al. 2016). The enrichment of V during sedimentation, diagenesis, or metamorphism in the selected deposits (Table 6 and cited references therein) is a function of host lithology, geologic setting, and most likely the metamorphic grade. McGill et al. (2024) showed that the depositional regime in which the V-enriched siliceous argillites formed was both anoxic and euxinic. The potential impact, especially of the latter, has so far not been sufficiently addressed in literature. It is, however, strongly assumed that metamorphism can contribute to V enrichment by liberation of V from carbonaceous matter due to the destruction of geoporphyrins in crude oils with increasing temperature (Lewan and Maynard 1982; Lewan 1984; Filby 1994; Greenwood et al. 2013). The role of metamorphism on V-bearing phases at the Van Property is discussed below in more detail.

The mineralogy of V-hosting phases has been studied in organic-rich sedimentary rocks. Many workers identified inorganic phases as hosts, which are clays illite, mannardite, and roscoelite, and/or oxides rutile or anatase, karelianite, eskolaite, schreyerite, oxyvanite, unknown V-Ti-bearing oxides, and rare sulfides such sulvanite (Coveney and Martin 1983; Carne and Gish 1999; Peacor et al. 2000; Di Cecco et al. 2018; Lu et al. 2021; Veselovský et al. 2021; Fu et al. 2023; Yang et al. 2023). Additionally, organic matter has long been known to hold substantial amounts of V since crude oil and the kerogen and bitumen fractions of sedimentary rocks can be enriched in V (Fester et al. 1927; Hodgson 1954; Krauskopf 1955; Fischer et al. 1968; Vine and Tourtelot 1970; Mercer et al. 1992, 1993; Filby 1994).

The current study at the Van Property identified two inorganic phases hosting relatively high amounts of V: (1) high-V illite with up to 13 wt% V₂O₃, and (2) rutile with up to 4.4 wt% V₂O₅ (Fig. 11; Table 5). Although illite is a common V-host in shale-hosted deposits, two types occur at the Van Property (Figs. 11 and 12), of which only high-V illite is relatively enriched in vanadium (average = 10.26 ± 2.32 wt% V₂O₃, Fig. 13), has a composition close to I/S (i.e., K = 0.51 ± 0.03 apfu, Fig. 12), is apparently the 1M polytype (personal communication A. Zhang 2023), is closely associated with carbonaceous matter, and does not show any indication of deformation (i.e., kinking). This is similar to the study by Lu et al. (2021), who identified V-enriched illite as 1M/1M_d polytype, closely associated with pyrite and organic matter. However, in contrast to the study of Lu et al. (2021), the high-V illite at the Van Property does not show pyrite inclusions and a formation involving dissolved organic matter and illitization during diagenesis cannot be assumed. The correlation of V with several elements on various crystallographic sites (Fig. 13) rather indicates that V occurs in more than one oxidation state in high-V illite. The negative correlation of V³⁺ with Al³⁺ on the octahedral site is common in high-V illite and also observed both in low-V illite at the Van Property and in V-bearing authigenic illite from the Mecca Quarry shale, Indiana (Coveney and Martin 1983; Peacor et al. 2000). Vanadium on the tetrahedral site negatively correlates with Si⁴⁺ and positively with Al³⁺ (Figs. 13a and 13b). This could either be due to coupled substitution (i.e., V⁵⁺ + Al³⁺ ↔ 2Si⁴⁺) or the occurrence of a vanadyl ion VO²⁺ on the tetrahedral site. Coupled substitution with V in the highly oxidized mobile state of V⁵⁺ (Breit and Wanty 1991; Wanty and Goldhaber 1992) has only been reported from synthetic illite in a hydrated environment (Wei et al. 1999) and is not seen as viable option due to the charge imbalance and ionic radii difference between Si⁴⁺ and V⁵⁺ of 0.26 and 0.355 Å, respectively (Shannon 1976), on the tetrahedral site. These charge and size differences would hinder effective substitution according to Goldschmidt’s rules of cation substitution (Goldschmidt 1937). In contrast, the vanadyl ion is a common component in organic-rich shales and mudstones since it is formed by the process of metalation of organic matter (i.e., formation of metallo-organic complexes such as geoporphyrins) and which occurs in unconsolidated sediments prior to diagenesis (Lewan and Maynard 1982). During maturation of carbonaceous matter by diagenesis and/or metamorphism, geoporphyrins demetallate, releasing VO²⁺ (Lewan and Maynard 1982; Lewan 1984; Filby 1994), which then can be introduced into inorganic minerals. Yang et al. (2023) described the occurrence of VO²⁺ and V³⁺ in mannardite that formed in strongly reduced, euxinic shales. Although the origin of V³⁺ and VO²⁺ cannot be constrained in this study at the Van Property, a hydrothermal origin for V³⁺ and VO²⁺ is less likely because the transport of VO²⁺ in hydrothermal fluids is restricted to relatively oxidized and acidic conditions (Lewan 1984; Breit and Wanty 1991; Wanty and Goldhaber 1992). Moreover, the solubilities of V³⁺ and VO²⁺ are limited in hydrothermal fluids (Wanty and Goldhaber 1992), and the amount of V occurring in the siliceous argillite at the Van Property could not be explained. As described below, a relatively Zn-rich, late-stage hydrothermal fluid was circulating through siliceous argillite, albeit the relative timing is unknown as of yet, depositing sphalerite (Figs. 5c, 5g, 5h, 11a, and 11h). It is unlikely that Zn and V were transported coevally since transport conditions (T, redox) differ between both metals. Moreover, the V concentration in sediment-hosted Zn deposits is not associated with hydrothermal activity but rather with scavenging from seawater or metamorphism (Canet et al. 2003; Paradis and Goodfellow 2012).

Rutile can have variable amounts of V that substitute for Ti⁴⁺ and, together with other V-bearing oxides such as karelianite, schreyerite, eskoalite, or oxyvanadite, have been reported in moderately metamorphosed shale deposits such as the Giant Green deposit, Madagascar (Di Cecco et al. 2018) and Chylin, Bohemian massif, Czech Republic (Veselovský et al. 2021). The vanadium content in rutile at the Van Property (1.13 ± 1.19 wt% V₂O₃) is lower than in rutile from the Green Giant deposit (2.88 ± 1.75 wt% V₂O₃; Di Cecco et al. 2018). We assume that the higher metamorphic grade (i.e., greenschist facies and higher) at both Green Giant and Chylin compared to the siliceous argillites at the Van Property resulted in a more complete breakdown of V-bearing clays and carbonaceous matter, resulting in the release of V, formation of rutile, and the subsequent incorporation of V⁴⁺ into metamorphic rutile (Di Cecco et al. 2018). Since the metamorphic grade at the Van Property was lower [i.e., sub-greenschist facies (Carne and Gish 1999; Flavelle 2013)], karelianite, eskoalite, schreyerite, or oxvanite did not form and the V content of rutile is lower. Further studies are, however, needed to confirm this assumption.

Although no spectroscopic or chromatographic analyses were performed to either identify if organic carbonaceous matter is present and to determine if it hosts V, the V concentration of crude oil, bitumen, and kerogen can reach several thousands of parts per million (Mercer et al. 1992; Chirinos et al. 2013). Furthermore, the positive correlation between V and TOC for siliceous argillites from the Van Property (Fig. 8c) indicates a strong role of total organic carbon in vanadium uptake. Although XRD identified graphite (Fig. 10) as the only carbonaceous phase, which is consistent with the degradation of organic matter under metamorphic conditions (Buseck and Huang 1985), Costantinides et al. (1959) and Lewan (1980, 1984) stated that V-bearing geoporphyrins in kerogen and bitumen do not fully demetallate (e.g., lose V-bearing organo-metallic complexes and form inorganic complexes such as VO²⁺) at temperatures in excess of 300–350 °C in crude oils. Peacor et al. (2000) showed that kerogen is an important host of V in the weekly metamorphosed black shales of the Mecca Quarry, Indiana. This finding supports the occurrence of V with carbonaceous matter at siliceous argillite (OSD) at the Van Property.

The model for the enrichment of V in shales or siliceous argillites must consider the age, depositional environment, lithological composition, mineralogy of V-hosted phases, amount of carbonaceous matter, and metamorphic grade. Currently, there are several explanations of how V is enriched in shale-hosted V deposits: (1) oxidized V⁵⁺ is scavenged from ambient seawater by organic matter or clays in anoxic or euxinic basins and incorporated in illite as reduced V³⁺ and/or in organic matter as V⁴⁺ (Breit and Wanty 1991) upon deposition or during late diagenesis, forming the additional V-bearing phases roscoelite, mannardite, or anatase (Fu et al. 2023); (2) V is scavenged directly from ambient seawater by dissolved organic matter of nano-size that then accumulated in shale and was taken up by smectite forming V-bearing smectite-nano-composites. During diagenesis, illitization of smectite and pyrolysis (i.e., maturation of organic matter) of organic matter results in V release and uptake of V³⁺ into illite (Lu et al. 2021); (3) low-T, hydrothermal brines circulated through sediments during deposition and early diagenesis forming V-rich, authigenic illite and V-rich organic matter (Peacor et al. 2000); (4) hydrothermal fluids containing immiscible, V-bearing petroleum droplets circulated through unconsolidated, organic-rich seafloor sediments and V³⁺ was incorporated into illite and organic matter. Diagenesis transferred V-bearing illite into roscoelite (Carne and Gish 1999). Alternatively, a petroleum ore fluid was the preferred transport medium for metals including V in shale-hosted deposits (Emsbo et al. 2009); and (5) secondary, supergene enrichment may further upgrade the V content in shales due to the circulation of meteoric fluids through V-enriched shales, forming V-rich goethite (Fu et al. 2023).

At the Van Property, processes 2 to 5 are dismissed as the dominant explanations of V enrichment. Fu et al. (2023) described the occurrence of micro-sized pyrite inclusions in illite-organic nanocomposites formed by dissolved organic matter that scavenged V from ambient seawater, were incorporated into smectite upon deposition and then into illite upon subsequent illitization of smectite. Although pyrite inclusions were observed in some illite, they only occur in low-V illite (Figs. 5e and 11h). Formation by syn-genetic and syn-diagenetic hydrothermal fluids that carried V either in a hydrothermal fluid (Peacor et al. 2000) or in immiscible petroleum droplets (Carne and Gish 1999; Emsbo et al. 2009) is also excluded for these reasons: (1) solubility of V in hydrothermal fluids is very low and cannot explain the amount of V present at the Van Property (Breit and Wanty 1991); (2) the amount of V held in the carbonaceous fraction is at least 40% (Table 7) and cannot be explained by minor petroleum droplets in a hydrothermal fluid. Although the source of the carbonaceous matter (e.g., primary: source rock or secondary: trap; Filby 1994, and references therein) cannot be constrained in this study, it is a major V source in which V became enriched either upon deposition and/or diagenesis in the two upper members of the OSD; (3) the occurrence of V is stratigraphically limited to the Lower Cherty Mudstone Member and the upper part of the Upper Siliceous Mudstone Member (Fig. 3). Any petroleum fluid flow also should have impacted the underlying members of the OSD (i.e., Calcareous Mudstone Member, Siliceous Mudstone Member, Transition Member) if there were a permeable pathway for petroleum fluids. However, these units are several tens to hundreds of meters thick and consist of less permeable, fine-grained siliciclastic lithologies (Carne and Gish 1999).

Additionally, only the upper part of the Upper Cherty Mudstone Member shows V enrichment (Carne and Gish 1999), whereas the remaining member has low-V content. This selective enrichment within the same stratigraphic horizon is difficult to explain by a circulating petroleum ore fluid that also enriched the older, conformably underlain Lower Siliceous Mudstone Member; (4) high-V illite and rutile formed during or after low-grade metamorphism and incorporated V either as V³⁺, V⁴⁺, or VO²⁺ in their crystal structure. Vanadium most likely originated from the degradation of carbonaceous matter and demetallation of geoporphyrins. Furthermore, the close spatial occurrence of high-V illite and rutile with randomly oriented pyrite and sphalerite and the strong positive correlation of V with Zn suggest that a low-T, Zn-bearing hydrothermal fluid circulated through the argillite; and (5) although Carne and Gish (1999) described the occurrence of secondary corvusite on weathered surfaces at the Janice and James Creek showings, this was not observed in the studied samples and hence, the role of supergene processes is minimal or even absent.

Based on the lithogeochemical and mineralogical results in this study, the enrichment and incorporation of V in various mineral phases in the siliceous argillite of the OSD at the Van Property occurred via different processes (Fig. 15). First, upon the deposition of quartz-rich, fine-grained sediments and organic matter under anoxic conditions and beneath the calcite compensation depth (Turner et al. 2011) in the Lower Devonian, V was incorporated as a trace component into low-V illite and formed geoporphyrins via metalation in carbonaceous matter during early diagenesis (Figs. 15a and 15b). The likely source of V was V⁵⁺ or V⁴⁺ scavenged either from ambient seawater [illite formation (Breit and Wanty 1991)] or pore water interacting with carbonaceous matter (Lewan 1980; Lewan and Maynard 1982; Filby 1994) where an open system was crucial for effective large-scale V scavenging (Lewan and Maynard 1982). During diagenesis, lithification of argillite occurred, as well as maturation of low-V illite to illite end-member composition (i.e., illite with K = 0.80–0.88 apfu) and maturation of carbonaceous matter (Fig. 15b). Furthermore, during (sub-)greenschist metamorphism, quartz and low-V illite deformed, the carbonaceous material transformed to graphite, and geoporphyrins demetallated (Fig. 15c). Lastly, a low-T, Zn-rich hydrothermal fluid circulated through siliceous argillite forming hydrothermal illite of I/S composition in which V was enriched (Fig. 15c). Due to the two oxidation states of V present in high-V illite (i.e., V³⁺ and V⁴⁺ in VO²⁺), we propose that vanadyl ions formed by demetallation of geoporphyrins were incorporated into a later generation of high-V illite during metamorphism. Further studies on high-V illite using transmission electron microscopy and atom probe tomography will be utilized to confirm this. The timing of the Zn-rich hydrothermal fluid is unknown and may be related to metamorphism during the Jurassic (i.e., a late metamorphic fluid) or intrusive activity that formed the nearby Cantung tungsten deposit in the Cretaceous (Ootes et al. 2013).

This study investigated the enrichment of V at the Van Property, Northwest Territories, using lithogeochemistry, SEM, XRD, and EMPA. The results show that: (1) V enrichment is not restricted to shale; (2) high-V illite, rutile, and assumingly carbonaceous matter are the predominant V hosts; and (3) the genesis is a complex interplay of syngenetic (e.g., V scavenging from seawater), diagenetic (e.g., metalation of V forming organo-complexes), and metamorphic (e.g., release of V due to de-metalation; hydrothermal fluid interaction) processes. These results have implications on both the role of metamorphism regarding V enrichment and exploration for shale-hosted V deposits. The former is not well constrained yet, and this study provides insight into how metamorphism can refine metal enrichment and the relationship of V uptake by both inorganic and organic phases during syngenetic, diagenetic, and metamorphic conditions. Vanadium has an increasing demand, and shale-hosted deposits are a viable yet under-exploited option. Results from this study aid in refining the current deposit model by incorporating variations in host lithogeochemistry and mineralogy and the role of metamorphism and, hence, target potential V prospects in metamorphosed Paleozoic sedimentary basins that have not yet been considered for exploration.

The authors thank the team of Chris Gunning and Adrian Zhang at SGS for XRD analyses and discussions on the results. Ravi Sidhu (Manitoba Institute of Materials, University of Manitoba) and Panseok Yang (Department of Earth Sciences, University of Manitoba) are thanked for their support at SEM and EMPA, respectively. S.M.B. thanks Michael Gadd for providing a detailed data set on the Moss prospect for this study. The authors appreciate the insightful review of an anonymous reviewer and Maria Boni, which improved the manuscript. We thank Associate Editor Denis Fougerouse for handling the paper. The study was funded by the Northwest Territories Geological Survey and an NSERC Discovery grant RGPIN-2022-04118 awarded to S.M.B. This is MERC publication number MERC-2024-02.