The Athabasca Basin unconformably overlies the Archean to Paleoproterozoic basement rocks known as the Western Churchill Structural Province (WSP). The WSP is divided into two lithostructural subprovinces: the Rae Subprovince to the west and the Hearne Subprovince to the east (Annesley et al., 2005) (Fig. 1A). The Trans-Hudson Orogeny (THO), which represents a major Paleoproterozoic belt, is linked to the collision between the Rae, Hearne, and Superior cratons. The beginning of this orogeny is characterized by the amalgamation between 1920 and 1840 Ma of the Rae and Hearne cratons, leading to the formation of the Western Churchill Structural Province (WSP) and to an oceanic arc in the Manikewan Ocean (Ansdell, 2005; Corrigan, 2012; Corrigan et al., 2005, 2009). The gradual closure of the Manikewan Ocean led to the successive accretion of juvenile crustal segments between 1880 and 1860 Ma. The final stage of the THO resulted in the complete closure of the Manikewan Ocean and in the continental collision between the margin of the Hearne craton bordered by the accreted juvenile arcs and the northern edge of the Superior craton between 1830 and 1800 Ma (Corrigan et al., 2009). Finally, the latter stages of THO deformation are marked by the emplacement at ca. 1800 and 1720 Ma of pegmatites and aplites, in relation to sinistral deformation along shear bands in a transpressional context (e.g., Chiarenzelli, 1989; Bickford et al., 1990; Chiarenzelli et al., 1998; Bickford et al., 2005; Schneider et al., 2007).

The western margin of the Hearne Craton comprises two distinct domains of Paleoproterozoic age: the Mudjatik domain and the Wollaston domain (Fig. 1A). The Mudjatik domain is predominantly composed of Archean tonalite-trondhjemite felsic gneiss domes (2640–2580 Ma; Annesley et al., 1997a, b, 1999a, b). In contrast, the Wollaston domain is composed mainly of Paleoproterozoic metasediments attributed to the Wollaston Supergroup (deposited at ca. 2050–1860 Ma; Yeo and Delaney, 2007; Tran et al., 2008), including paragneisses, banded iron formations, quartzites, and marbles. These units are interpreted as evidence of sedimentation initially controlled by a passive-margin setting, which evolved into a back-arc basin and then into a foreland basin in response to tectonic processes related to the collision phase (Annesley et al., 2005; Jeanneret et al., 2016). The two lithotectonic domains are separated by the Wollaston-Mudjatik Transition Zone (WMTZ), within which the present study area is located (Fig. 1A). This zone exhibits an anastomosed structure oriented from northeastern to southwestern, a trait that was inherited from the THO. The deformation phases associated with the various stages of THO formation mentioned above can be observed within this zone.

The basement rocks underwent three tectono-metamorphic events. The first event (M1–D1) is dated from ca. 1840 to 1813 Ma based on U–Pb dating of monazite and zircon from migmatitic metapelites of the Wollaston Supergroup. These rocks host a garnet-cordierite ± sillimanite assemblage that experienced partial melting at peak conditions of 750–825 °C and ≤ 1 GPa (Annesley et al., 1992, 1997a, b, c, 1999a, b; Jeanneret et al., 2017; Toma et al., 2024). The second event (M2–D2), dated from ca. 1813 to 1770 Ma, is recorded within migmatites and pegmatites. It corresponds to an isothermal decompression down to 0.5–0.9 GPa at temperatures of 750–825 °C, within the stability field of cordierite under suprasolidus conditions, as constrained by U–Pb dating of monazite and zircon (Annesley et al., 1992, 1997a, b, c, 1999a, b; Jeanneret et al., 2017; Toma et al., 2024). The third event corresponds to an exhumation and orogenic cooling phase. Although its onset is poorly constrained, this phase is generally considered to extend from ca. 1770 Ma (750 °C, 0.5 GPa) to ca. 1720 Ma (350 °C, 0.2 GPa), as documented by several studies (Annesley and Madore, 1994; Annesley et al., 1992, 1997a, b, c, 2005; Jeanneret et al., 2017; Martz, 2017). This exhumation phase is characterized by pronounced retrograde transformations linked to early-stage hydrothermal activity. These include the breakdown of cordierite in favour of phyllosilicates, biotite chloritization, remobilization and precipitation of sulfides, graphite, and muscovite, circulation of quartz-rich fluids, and dissolution-reprecipitation processes affecting monazite and zircon (Card, 2012, 2014; Card and Noll, 2016).This phase is further constrained by cooling ages derived from biotite Rb–Sr (Worden et al., 1985; Schneider et al., 2007), muscovite K–Ar (Philippe et al., 1993), and muscovite Ar–Ar systems (Alexandre et al., 2009).

The Paleoproterozoic to Mesoproterozoic Athabasca Basin is a large intracratonic basin located in northern Saskatchewan and Alberta (Ramaekers et al., 2007). The sedimentary sequence of this basin, known as the Athabasca Supergroup, was deposited between ca. 1760–1720 Ma (Ramaekers et al., 2007; Jeanneret et al., 2016) and 1541 Ma for the youngest known age (Creaser and Stasiuk, 2007). Stratigraphic, geochronological, and fluid-inclusion P-T data indicate that the basin reached a maximum depth of 2 to 5 km depending on the methods considered (Pagel et al., 1975a, b; Chi et al., 2018), with a preserved thickness of 1.5 km today (Hobson et al., 1969). The stratigraphic sequence is organized into four major clastic sedimentary sequences that record a progressive deepening of the basin (Jefferson et al., 2007). The first sequence, represented by the Fair Point Formation, consists of conglomerates and quartz-pebble arenites infilling the Jackfish Sub-basin, located in the western part of the basin, during the initial sag phase around 1760 Ma (e.g., Ramaekers et al., 2007; Pehrsson et al., 2023). The second sequence, including the Smart/Read and Manitou Falls formations, fills the Cree Sub-basin, situated in the eastern and southeastern parts of the basin, and reflects the development of a shallow half-graben with NW-directed palaeoflow between 1680 and 1650 Ma (Ramaekers et al., 2007; Jeanneret et al., 2016; Pehrsson et al., 2023). The third sequence, represented by the Lazenby Lake and Wolverine Point formations, shows vertical fining and records a major paleoenvironmental shift around 1640 Ma, characterized by abrupt deepening in the western basin, a marine incursion, and a transition to axial palaeocurrents sourced from the north or east (Rainbird et al., 2007; Ramaekers et al., 2007), as indicated by vitric tuff layers dated at 1644 ± 13 Ma (U–Pb zircon; Rainbird et al., 2007). The final sequence, which includes the Locker Lake, Otherside, Douglas, and Carswell formations, reflects detrital and marine sedimentation during continued subsidence and tectonic reorganization associated with the Kuungmi rifting (Pehrsson et al., 2023), notably recorded by organic-rich black shales of the Douglas Formation dated by Re–Os at 1541 ± 13 Ma (Creaser and Stasiuk, 2007). Overall, the basin maintains a largely coherent oval geometry with minimal deformation, except for the Ordovician Carswell impact structure, which locally disrupts the Athabasca Supergroup (Ramaekers et al., 2007; Pehrsson et al., 2023).

The Athabasca Basin unconformity-related uranium deposits are globally renowned for their exceptionally high grades, reaching up to 20 wt % U₃O₈. The principal uranium deposits are concentrated in the eastern part of the Athabasca Basin and are spatially associated with a NE-SW-trending anastomosed transition zone inherited from the Trans-Hudsonian Orogeny (THO), commonly referred to as the Wollaston-Mudjatik Transition Zone. This structural corridor comprises a ca. 50-km-wide network of graphite-rich shear zones, typically described as graphite conductors, as they exhibit high electrical conductivity, a physical property that allows their detection by electromagnetic geophysical methods used for uranium exploration (Fig. 1). The uranium deposits are spatially associated with graphite-rich shear zones, and uranium orebodies show diverse morphologies and structural associations depending on their location: (i) at the unconformity, within the first meters of sandstones and conglomerates; (ii) within the basement, extending several hundred meters below the unconformity along shear zones; or (iii) as perched deposits located tens to hundreds of meters above the unconformity within the sedimentary sequence. The mineralized zones are surrounded by hydrothermal alteration halos that can extend several tens of meters into both the basin and the basement and serve as critical geochemical and mineralogical guides for U exploration. In the basin, alteration halos overlying basement-hosted deposits or surrounding sandstone-hosted mineralization are dominated by illite, with variable amounts of kaolinite, sudoite (tri-dioctahedral Mg-chlorite), and dravite (magnesian tourmaline), depending on the deposit.

Within the basement, three principal alteration types are superimposed, as exemplified by the world-class Cigar Lake deposit (Fig. 1B). The first type corresponds to retrograde metasomatic alteration related to the final exhumation and orogenic cooling phase of the Trans-Hudsonian Orogeny (THO). It is marked by partial to complete biotite chloritization, discrete quartz and muscovite crystallization, and hydrothermal graphite formation. These alteration types and their associated petrographic features have been extensively documented in the vicinity of unconformity-related uranium deposits within the Athabasca Basin (e.g., Carl et al., 1992; Derome et al., 2005; Cloutier et al., 2011; Mercadier et al., 2011; Alexandre et al., 2012; Adlakha and Hattori, 2015; Martz et al., 2019a; Menier et al., 2020; Powell et al., 2022). The second type corresponds to regolithic alteration developed within the upper tens of meters of the basement following its exhumation after the Trans-Hudsonian Orogeny. It is characterized by a laterally correlative weathering profile comprising four zones: (i) a bleached kaolinite-rich and hematite-depleted zone at the unconformity; (ii) a highly oxidized hematite-rich zone; (iii) a red-green transitional zone characterized by hematite and chlorite; and (iv) a thicker green-to-red zone enriched in illite, chlorite, and hematite (e.g., Macdonald, 1980; Halter, 1988). The third type corresponds to post-basin-deposition hydrothermal alteration, which represents the most extensive alteration type. It is proposed to be linked with uranium mineralization and is characterized by the massive replacement of primary minerals by an illite- and sudoite-dominated clay assemblage.

The main accepted model for the formation of these deposits is known as “diagenetic-hydrothermal” and involves the circulation of oxidized brines of marine origin containing NaCl-CaCl₂ at temperatures of around 120–200 °C in the basin and underlying bedrock (Derome et al., 2005; Richard et al., 2011, 2013, 2015). The interaction of these brines with basin and/or basement rocks favors the formation of alteration halos and mobilizes uranium in its oxidized form (U⁶⁺), which is transported through fractures and reactivated graphitic structures. Uranium precipitates within structural and physicochemical traps, where dissolved hexavalent uranium (U⁶⁺) is reduced to tetravalent uranium (U⁴⁺) (Hoeve and Sibbald, 1978; Jefferson et al., 2007).

Published absolute ages for the precipitation of primary uranium mineralization at the scale of the basin span a broad range, from ca. 1590 to 1200 Ma. These estimates derive from U–Pb analyses of UO₂ and from K–Ar or Ar–Ar geochronology of clay minerals (Cumming and Krstic, 1992; Philippe et al., 1993; McGill et al., 1993; Fayek and Riciputi, 2002; Alexandre et al., 2009; Cloutier et al., 2011; Powell et al., 2022). Building on these age constraints, an alternative hypothesis invoking an earlier episode of uranium mineralization and/or hydrothermal alteration merits consideration. Numerous studies have shown that U–Pb systematics of uraninite and associated alteration minerals in unconformity-related deposits are commonly overprinted by multiple fluid-flow events, resulting in partial Pb loss and a broad dispersion of apparent ages (Kotzer and Kyser, 1990; Fayek and Kyser, 1997; Fayek et al., 2002; Fayek and Riciputi, 2002; Kyser et al., 2015). Consequently, the existence of an older hydrothermal system, potentially active in the ca. 1640–1680 Ma interval, cannot be ruled out and may be cryptic due to subsequent alteration, uranium remobilization, and isotopic resetting during younger tectono-hydrothermal episodes. This interpretation is consistent with evidence for repeated basin-scale fluid circulation and late-stage meteoric fluid ingress documented in the Athabasca Basin, processes that are known to modify both clay mineral K–Ar / Ar–Ar ages and uraninite U–Pb signatures. A similarly complex temporal evolution is recognized in unconformity-related uranium systems of northern Australia, where in situ U–Pb analyses of uraninite suggest initial mineralization or hydrothermal alteration as early as ca. 1680–1640 Ma, followed by multiple younger resetting events extending into the Mesoproterozoic and Paleozoic (Clauer et al., 2015; Skirrow et al., 2016).

Four study areas located in the northeastern segment of the Wollaston-Mudjatik Transition Zone (WMTZ) were selected to evaluate the applicability of in-situ Rb–Sr geochronology to date the main hydrothermal alteration (illite-sudoite) associated with the U mineralisation within the basement rocks. From north to south, these areas comprise the Waterfound (WF) and McClean South (MCS) prospects, and the Cigar Lake (CL) and McArthur River (MAC) deposits. All samples from WF, MCS, and CL analyzed in this study were obtained from drill cores collected during the past five years as part of U exploration programs conducted by Orano Canada Inc., whereas samples from MAC were provided by Cameco Corporation. The sampling strategy at the four study sites was designed to ensure both lateral and vertical representativeness, spanning from unaltered basement zones (Fig. 2A, B) to hydrothermally altered domains characterized by illite-sudoite assemblages associated with the mineralization (Fig. 2C, D, E, F). For MCS, five drill holes were considered from west to east: MCS-03, MCS-05, MCS-34, MCS-35C and MCS-31 (mineralized conductor). For WF, six drill holes were considered from southwest to northeast: WF-93, WF-98, WF-66, WF-77 and WF-75 (mineralized conductor) and WF-100 (outside the La Rocque Conductive Corridor). For CL, two drill holes were considered from west to east: WC-449 (mineralized conductor) and WC-473 (non-mineralized conductor). For MAC, five drill holes were considered from west to east: MC-408, MC-408-01, MC-418, MC-410-01, MC-410-02 (mineralized conductor). Pelitic gneisses and anatectic granitoids represent the two dominant lithologies within the WMTZ, particularly within the graphite-rich shear zones that host mineralization, which justifies their selection for this study. Across all investigated sites, four samples of pelitic gneisses unaffected by illite-sudoite alteration were analyzed, together with five samples of anatectic granitoids and/or associated leucosomes that were likewise preserved from this alteration. Their altered equivalents comprise ten samples of pelitic gneisses affected by illite-sudoite alteration and nineteen samples of anatectic granitoids and/or leucosomes exhibiting the same alteration signature. All samples were specifically selected based on their high muscovite modal abundances and the presence of muscovite grains of sufficient size to enable in situ Rb–Sr analyses by LA-ICP-MS.

Scanning electron microscopy (SEM), electron probe microanalysis (EPMA), and micro-X-ray fluorescence (µ-XRF) analyses were performed at the Service Commun de Microscopies Électronique et de Microanalyse X (SCMEM) in Nancy, France, whereas in situ Rb–Sr isotopic analyses were carried out on the ICP platform at GeoRessources, Université de Lorraine (Nancy, France).

Metamorphic muscovite is exclusively developed along the foliation planes of Paleoproterozoic metapelites (Fig. 2A, B) (e.g., Schneider et al., 2007; Skipton et al., 2016; Jeanneret et al., 2017; Martz, 2017). These rocks are characterized by quartz and sericitized K-feldspar, and are bounded by graphite-rich shear zones containing biotite, muscovite, and pyrite (Fig. 3A). Biotite occurs as lath-shaped crystals, locally plastically deformed showing C-S planes, ranging from 0.5 to 1 mm in length, and oriented parallel to the subvertical S1 foliation. Muscovite is present as euhedral to subhedral crystals, predominantly crystallizing along the rims of metamorphic biotite, with an orientation perpendicular to the S1 foliation, suggesting a crystallization stage postdating biotite growth possibly along the retrograde P-T path (Fig. 3A). Muscovite is also present within pressure shadows around biotite mica-fish, suggesting that muscovite crystallization was contemporaneous with ductile deformation (Fig. 3B). Metamorphic biotites are locally pseudomorphosed by Fe-Mg chlorite, a marker of retrograde metamorphism or fluid-induced alteration (Fig. 3C). The metamorphic muscovite is commonly associated with medium-sized, euhedral to tabular flakes of metamorphic graphite aligned along the foliation. Sulfides such as pyrite and chalcopyrite are frequently associated with this metamorphic graphite (Fig. 3D).

Early hydrothermal muscovite occurs within anatectic granitoids or within the leucosomes of metapelitic gneisses (Fig. 2C, D, E, F) (e.g., Schneider et al., 2007; Skipton et al., 2016; Martz, 2017; Obin, 2025). These lithologies are primarily composed of quartz, K-feldspar of the microcline type, often altered to sericite forming very fine microlamellar aggregates with a shimmering whitish-grey appearance, plagioclase, and biotite laths, which are partially to completely pseudomorphosed by Fe-Mg chlorite (Fig. 3E, F). Muscovite in these rocks occurs either as subhedral crystals or as fibrous aggregates that may have partially to completely replaced chloritized biotite or have precipitated along biotite grain boundaries as fine-grained crystals (Fig. 3E, F, G). It may also have locally replaced K-feldspar (Fig. 3F). Muscovite is frequently associated with hydrothermal graphite, which occurs as large flakes (up to 0.5 cm) precipitating interlayered with muscovite within late fractures or forming radiating patterns in granitic leucosomes (Fig. 3H). These lithologies may also be locally enriched in tourmaline, occurring as euhedral to subhedral crystals with strong relief and second-order yellow to green interference colors. Accessory phases include zircon, monazite, and apatite, present as subhedral crystals or as veinlets crosscutting the aforementioned minerals.

This hydrothermal alteration affects both metamorphic muscovite (Fig. 4A–C) and early hydrothermal muscovite (Fig. 4D–I). This Mg-K-type alteration pervasively affects the two lithologies that constitute the graphite-rich shear zones, whether through the mylonitic foliation planes of pelitic gneisses (Fig. 4A–C), the grain boundaries of magmatic to early hydrothermal minerals in anatectic granitoids and leucosomes (Fig. 4D–I), or through fracture networks indiscriminately crosscutting the two lithologies. At the microscopic scale, this alteration is expressed by the extensive replacement of primary and retro-metamorphic minerals by a clay assemblage predominantly composed of micrometric illite, spatially associated with sudoite (a magnesium-rich trioctahedral chlorite). However, this alteration is not uniformly distributed within the different lithologies, and the illite/sudoite ratio varies between lithologies and within the same lithology (Fig. 4D, G). In anatectic granitoids and leucosomes, the alteration is marked by the partial replacement of K-feldspar and the complete replacement of plagioclase, which are more prone to alteration, by this clay assemblage (Fig. 4D–I). Fractures crosscutting quartz and feldspars are commonly filled with illite and/or sudoite crystals (Fig. 4A). The cleavage planes of muscovite are partially to completely disrupted, and their birefringence colors approach those of illite, reflecting the onset of muscovite illitization, particularly visible along grain margins (Fig. 4E, F, H, I). Under cross-polarized light, illite occurs as lath-shaped crystals (Fig. 4E, F, H, I). Sudoite appears as needle-like crystals with low relief, generally < 2 µ m in length (Fig. 4B, C, H, I). Figure 5 highlights the effect of illite-sudoite alteration within a single sample (MC-410-01, 677.9 m), comparing a relatively preserved zone showing the onset of alteration (Fig. 5A) with a fully altered zone (Fig. 5B). Backscattered electron (BSE) images reveal that in the least altered area, muscovite retains its characteristic cleavage planes (Fig. 5A), whereas in the most altered area, muscovite has completely lost its original morphology and is replaced by sudoite ± illite. Mixtures of these two phases can be distinguished through grayscale contrast variations (Fig. 5B). Elemental X-ray maps obtained by EPMA further emphasize these mineralogical and chemical transformations (Fig. 5C–H). In the least altered zone, muscovite exhibits variable K (Kα) intensities, with maximum values in the most preserved domains and progressively decreasing intensities in areas showing incipient alteration (Fig. 5C). The Rb (Kα) signal follows a similar pattern, reaching its highest values in the K-rich zones (Fig. 5E). This decrease in K and Rb intensities correlates with the occurrence of Mg (Kα)-rich areas associated with sudoite formation along fracture zones (Fig. 5G). In the most altered zone, muscovite relics (or “ghosts”) are identified by their relatively high K (Kα) and Rb (Kα) intensities (Fig. 5D, F), which are spatially anti-correlated with the Mg (Kα) signal. The latter reaches its maximum intensity within the sudoite matrix that surrounds these muscovite remnants (Fig. 5H).

Rb–Sr isotopic analyses were performed on metamorphic and early hydrothermal muscovite sampled from both unaltered and illite-sudoite-altered zones, and from the two investigated lithologies (Table S1). To assess the impact of the illite-sudoite alteration on the isotope system of basement minerals, a representative sample of a pelitic gneiss (MC-410-01, at a depth of 677.9 m) containing both well-preserved (Fig. 6A, B) and partially preserved to altered muscovite grains (Fig. 6C, D) was selected for detailed in situ analysis by LA-ICP-MS/MS. The complete Rb–Sr dataset obtained from this sample defines a Rb–Sr regression corresponding to a date of 1761.9 ± 30.6 Ma (n= 34, MSWD = 5.8, p≈0, 2σ internal; ±41.6 Ma including external reproducibility), with ⁸⁷Rb / ⁸⁶Sr ratios ranging from 19.62 to 670.44 and a corresponding ⁸⁷Sr / ⁸⁶Sr intercept of 0.6524 ± 0.0096 (Fig. 6E). The elevated MSWD value and the null probability associated with the reduced χ2 test indicate significant excess scatter in the isotopic data, consistent with heterogeneous isotopic behavior at the grain scale. Given this dispersion, Rb–Sr single-spot dates were calculated for each laser ablation point following the approach of Rösel and Zack (2022), allowing for a direct comparison between isotopic data, petrographic observations, and elemental mapping. The resulting single-spot dates display a systematic relationship with K (Kα) intensity maps, showing a decrease in apparent dates with decreasing K intensity, consistent with the progressive alteration of muscovite (Fig. 6B, D). Radial plots illustrating the distribution of single-spot dates and isotope ratios (Galbraith, 1988, 1990; Rösel and Zack, 2022) reveal two statistically distinguishable populations with partial overlap: a dominant population at 1785.2 ± 11.2 Ma (76.8 %), corresponding predominantly to the least altered muscovite domains, and a younger population at 1646.9 ± 20.4 Ma (23.2 %), associated with increasingly altered domains (Fig. 6F). When plotted separately on Rb–Sr regression diagrams, analyses from the least altered muscovite domains yield a statistically robust regression corresponding to a date of 1789.1 ± 37 Ma (n= 23, MSWD = 0.51, p= 0.9689, 2σ internal; ± 47 Ma including external reproducibility), with a ⁸⁷Sr / ⁸⁶Sr intercept of 0.6709 ± 0.0115 (Fig. 6G). In contrast, analyses from the more altered and mixed domains define a less statistically reliable regression at 1659.5 ± 55.4 Ma (n= 11, MSWD = 4.1, p= 2.6×10-5, internal; ± 61 Ma including external reproducibility), associated with a ⁸⁷Sr / ⁸⁶Sr intercept of 0.6316±0.018 (Fig. 6H). The elevated MSWD values obtained for the altered domains again indicate substantial isotopic scatter. On the basis of these observations and the coexistence of mixed isotopic populations within a single sample, the analytical approach applied to the entire dataset, supported by petrographic and microstructural evidence, consisted of distinguishing the most preserved metamorphic and early hydrothermal muscovite from those more strongly altered by an illite-sudoite clay assemblage, in order to constrain the crystallization dates of these two muscovite types and the potential timing(s) of alteration (Fig. 7A–E).

The weighted mean of the Rb–Sr data obtained from all preserved metamorphic muscovite grains (four samples) defines a composite Rb–Sr regression at 1780.8 ± 14.6 Ma (n = 57, MSWD = 0.92, p = 0.63, 2σ internal; ± 32 Ma including external reproducibility), associated with a ⁸⁷Sr / ⁸⁶Sr intercept of 0.6704 ± 0.0075 and overall ⁸⁷Rb / ⁸⁶Sr ratios ranging from 19 to 671 (Fig. 7A). Rb–Sr single-spot dates for preserved metamorphic muscovite were calculated using a fixed ⁸⁷Sr / ⁸⁶Sr reference value of 0.6704, corresponding to the intercept of the composite regression. This value is used here as an operational reference parameter; owing to its low magnitude, reasonable variations in this parameter exert only a minor influence on the calculated dates, as previously demonstrated by Rösel and Zack (2022). The resulting single-spot dates range from 1760 ± 30 to 1914 ± 88 Ma and define a single statistically coherent population with a modal value of 1771.2 Ma (Fig. 7E). Matrix-corrected elemental concentrations (using an applied correction factor of 0.532038) range from 33 to 145 ppm (average: 81 ppm) for ⁸⁷Rb, from 0.19 to 2.46 ppm (average: 1.03 ppm) for ⁸⁶Sr, and from 1.28 to 4.50 ppm (average: 2.73 ppm) for ⁸⁷Sr.

The study of basement lithologies along a gradient from non-hydrothermally altered to strongly hydrothermally altered zones beneath unconformity-type uranium deposits reveals significant mineralogical, geochemical, and isotopic modifications. Across the four investigated sites, this alteration, macroscopically characterized by an apple-green color affecting the foliation planes of metapelitic gneisses, leucosomes, and granitoids, as well as filling fractures crosscutting the host lithologies (Figs. 2C, D, E, F and 4), manifests microscopically as the progressive replacement of primary and retro-metamorphic minerals (plagioclase, K-feldspars, muscovite, biotite, Fe-Mg chlorite) by an assemblage of illite and sudoite. Textural evidence suggests that these two minerals crystallized synchronously or nearly synchronously (Pacquet and Weber, 1993; Percival and Kodama, 1989). This alteration induces pronounced mineralogical transformations in both metamorphic and early hydrothermal muscovite, involving a partial to complete breakdown of their crystal structure. This process is expressed either as illitization initiated along the cleavage planes of muscovite, synchronous or quasi-synchronous with sudoite precipitation, or as massive sudoite crystallization enveloping muscovite relics (Figs. 4 and 5B). Illite-sudoite alteration constitutes the most pervasive and intense hydrothermal alteration affecting both the Athabasca Basin sandstones and the underlying basement in the vicinity of unconformity-related uranium mineralization, and has long been recognized as a robust exploration vector for U deposits (e.g., Carl et al., 1992; Derome et al., 2005; Cloutier et al., 2009; Ng et al., 2013; Martz, 2017; Kaczowka et al., 2021; Powell et al., 2022). This alteration is associated with a loss of Al, Na, K and Sr and a gain in Mg, Li and B (Martz et al., 2019a). The loss of K, Na, Al and Sr is easily explained by the K-feldspars and plagioclases replacement by illite plus sudoite. Some of these elements remain in the system and are in situ transferred to the newly-formed clays, whereas the rest leave the system. This process results from the circulation of marine-derived basinal brines, generated during early to middle diagenesis of the sedimentary basin (Hoeve and Sibbald, 1978; Kotzer and Kyser, 1995; Kyser et al., 2000).

This illite + sudoite assemblage formed at temperatures between 150 and 220 °C, under pressures corresponding approximately to the base of the sedimentary basin (ca. 0.1–0.3 GPa) (Halter, 1988; Kotzer and Kyser, 1995; Cloutier et al., 2009; Martz, 2017). It is associated with important chemical modifications of the host rocks, notably the loss of K, Na, Al, and Sr due to destabilization of primary feldspars and micas, and the concomitant gain in Mg, Li, B, and Ca reflecting the composition of the basinal brine and its prolonged interaction with silicates. The observed B and Mg enrichments are attributed to the intrinsic composition of the basinal fluid (Mercadier et al., 2012; Richard et al., 2011) rather than to a local metamorphic source. Variations in the illite/sudoite ratio reflect changes in thermal gradient (Beaufort et al., 2005; Kotzer and Kyser, 1995; Martz, 2017), in fluid composition (Cloutier et al., 2009), notably the K⁺ / Mg²⁺ ratio, as well as in host rock lithology (Bruneton, 1993; Pacquet and Weber, 1993; Mercadier et al., 2012; Cuney and Kyser, 2015; Martz, 2017) and system pH (Kyser et al., 2000; Kister et al., 2005; Mercadier et al., 2012; Martz, 2017). The illite-sudoite alteration is not related to the recrystallization of new generations of metamorphic or early hydrothermal muscovite but instead results from fluid circulation that induces mineralogical, geochemical, and isotopic modifications. Isotopically, progressive alteration leads to systematic variations in Rb–Sr systematics, as mineralogical transformations modify both elemental concentrations and isotope ratios. The range of ⁸⁷Rb / ⁸⁶Sr ratios broadens from 20–670 to 5–3305, while ⁸⁷Sr / ⁸⁶Sr ratios shift from 1–18 to 0.8–77, with corresponding single-spot dates decreasing from 1701.9–1914.8 to 1357.0–1756.2 Ma (Fig. 9). The behaviour of the Rb–Sr geochronometer in the presence of fluid-rock interaction is best understood in terms of fluid-mediated recrystallization and open-system isotopic exchange, rather than purely thermally-activated volume diffusion (Villa, 1998; Villa, 2022).

The observed shift in Rb–Sr systematics is primarily attributed to the partial to complete loss of radiogenic ⁸⁷Sr from the system in response to fluid circulation associated with illite-sudoite alteration. This process results in partial to complete resetting of the Rb–Sr isotope system (Matheney et al., 1990; Kalt et al., 1994; Evans et al., 1995; Eberlei et al., 2015). Elemental maps further demonstrate that Rb is also mobile; however, the mobility of radiogenic ⁸⁷Sr is significantly greater than that of Rb and non-radiogenic Sr. This contrast in mobility accounts for the observed variations in both Rb / Sr ratios and Sr isotopic compositions (Fig. 5E and F) and could explain the erratic data distribution. The structural destabilization of metamorphic and early hydrothermal muscovite during hydrothermal alteration, facilitated by illitization and/or sudoitization along cleavage planes, may have enhanced diffusion and leaching of radiogenic ⁸⁷Sr. This process likely promoted the transfer of ⁸⁷Sr into clay phases and/or the fluid phase (Fig. 9). In addition, radiogenic ⁸⁷Sr produced in situ occupies the interlayer (I-site) in muscovite, substituting for monovalent cations such as K⁺ or Rb⁺ (Brigatti and Guggenheim, 2002), whereas common Sr (⁸⁶Sr, ⁸⁸Sr) is predominantly hosted in the octahedral (M-site), where divalent cations are stable. This crystallochemical distinction means that preferential removal of radiogenic ⁸⁷Sr from the I-site alone would be expected to drive the regression intercept toward chondritic or higher values, not toward the sub-chondritic values observed here. The anomalously low intercept values (∼ 0.67) are therefore better interpreted as the result of isochron rotation under open-system conditions involving non-conservative, grain-scale redistribution of both Rb and Sr (Brooks et al., 1976; Villa, 1998; Villa, 2022), rather than as a direct consequence of selective I-site ⁸⁷Sr loss. As demonstrated by Villa (1998, 2022), fluid-mediated recrystallization operates orders of magnitude faster than thermally-activated volume diffusion and represents the dominant mechanism of isotopic exchange in micas affected by hydrothermal fluids. In this framework, the structural destabilization of muscovite along cleavage planes during illitization and sudoitization, facilitated by the circulating Ca-Cl₂-dominated brines, promoted fluid-assisted dissolution-reprecipitation and non-conservative redistribution of Rb and Sr at the grain scale, consistent with the hygrochronological framework described by Villa (2022). The transfer of Sr into the fluid phase is independently supported by Sr enrichment documented in fluid inclusions from the Athabasca Basin brines (Mercadier et al., 2012; Richard et al., 2011; Martz, 2017).

Furthermore, the measured isotope ratios ⁸⁷Rb / ⁸⁶Sr and ⁸⁷Sr / ⁸⁶Sr, as well as ⁸⁷Rb and ⁸⁷Sr concentrations, are significantly higher in altered muscovites than in unaltered muscovites. The simultaneous increase in ⁸⁷Rb / ⁸⁶Sr and ⁸⁷Sr / ⁸⁶Sr ratios in muscovites affected by alteration or fluid circulation reflects reopening of the Rb–Sr system and selective mass redistribution of Sr during the hydrothermal event. Glodny and Grauert (2009) and Eberlei et al. (2015) demonstrate that such concurrent increases do not require mica recrystallization but rather result from a combination of dynamic fluid-rock interaction processes. These include: (i) a net loss of Sr through accelerated diffusion along defects (microfractures, cleavage planes, subgrains, dislocations), favoring the preferential loss of radiogenic Sr from transient interlayer sites or defect-related positions in muscovite; (ii) a decrease in the molar fraction of Sr, which automatically increases the ⁸⁷Rb / ⁸⁶Sr ratio since Rb is relatively immobile; and (iii) a minor incorporation of Sr, possibly more radiogenic, from the matrix reservoir or from the fluid (e.g., derived from feldspar alteration) into altered muscovite, which can slightly increase the measured ⁸⁷Sr / ⁸⁶Sr ratio.

The characterization of well-preserved micro-domains enables the acquisition of statistically robust isotopic ages for both the metamorphic muscovite associated with the foliation of metapelitic gneisses and the early hydrothermal muscovite from anatectic granitoids or occurring in the leucosomes of metapelitic gneisses. These muscovite grains preserved from illite-sudoite alteration define well-correlated Rb–Sr regressions yielding dates of 1780.8 ± 14.6 Ma (2σ internal; ±32 Ma including external reproducibility) and 1753.4 ± 12.4 Ma (2σ internal; ±31 Ma including external reproducibility), respectively (Fig. 7A, B). It should be noted that in open-system Rb–Sr behaviour, fluid-rock interaction can rotate an originally valid isochron through non-conservative redistribution of Rb and Sr at the grain scale, yielding geologically meaningful slopes but mathematically anomalous intercepts that fall below physically feasible initial ⁸⁷Sr / ⁸⁶Sr values (Brooks et al., 1976; Villa, 1998; Villa, 2022). As demonstrated by Villa (1998, 2022), the isotopic perturbations formerly attributed to thermally-driven volume diffusion in micas have since been shown to reflect fluid-mediated recrystallization, which operates orders of magnitude faster than pure Fickian diffusion and represents the dominant mechanism of isotopic exchange in micas affected by hydrothermal fluids. Such rotation does not invalidate the age information recorded by the slope of the regression. However, the corresponding regression intercepts (0.6704 ± 0.0075 and 0.6989 ± 0.0168) fall below the minimum ⁸⁷Sr / ⁸⁶Sr value of the solar system (∼ 0.698) and therefore cannot represent physically meaningful initial isotopic compositions. These intercepts are thus not interpreted as true initial ⁸⁷Sr / ⁸⁶Sr ratios but rather as apparent values resulting from disturbed Rb–Sr systematics. Low MSWD values (<1) and high p-values (p≫0.05) indicate strong internal coherence of the datasets, suggesting that the regressions are statistically robust despite their anomalous intercepts. Such features are characteristic of rotated or disturbed isochrons formed under open-system conditions, in which Rb and Sr are redistributed in a non-conservative manner (Brooks et al., 1976; Faure et al., 2013; Dickin, 2018). Proportional loss of Rb relative to Sr, selective Sr mobility during fluid-rock interaction, or partial isotopic resetting during metamorphism may rotate an originally valid isochron, yielding geologically reasonable slopes but non-physical intercepts (Brooks et al., 1976; DePaolo, 1981; Villa, 1998; Villa, 2022; Faure et al., 2013). The reproducibility of measured ⁸⁷Sr / ⁸⁶Sr ratios for the NIST SRM 610 reference material, together with the consistency of the regression intercept obtained for the La Posta biotite reference material, indicates that the measured ⁸⁷Sr / ⁸⁶Sr ratios and derived regression intercepts are not attributable to analytical bias or standardization issues, thereby effectively excluding an analytical or calibration-related origin. Accordingly, the anomalously low intercept values are interpreted as mathematical artefacts arising from open-system Rb–Sr behavior and isochron rotation under fluid-mediated conditions, rather than as geologically meaningful isotopic signatures or evidence of an analytical artefact. In contrast, the slopes of the regressions may still record geologically significant timing information related to muscovite crystallization or partial isotopic resetting, as has been documented in other disturbed Rb–Sr systems (Brooks et al., 1976; Villa, 1998; Villa, 2022; Zack and Hogmalm, 2016).

The single-spot ages obtained from metamorphic muscovite developed along foliation planes in Paleoproterozoic pelitic gneisses (e.g., Schneider et al., 2007; Skipton et al., 2016; Jeanneret et al., 2017; Martz, 2017) range from 1914 ± 88 to 1760 ± 30 Ma, with a modal value of 1771.2 Ma and an isochron age of 1780.8 ± 14.6 Ma (2σ internal; ±32 Ma including external reproducibility) (Fig. 10). The mean age derived from single-spot analyses, together with the isochron age, falls within the temporal window of the M2-D2 event, which generated the northeast-striking structural trend of the WMTZ and formed under a sinistral transpressional tectonic regime during the late stages of the oblique Hudsonian collision, between ca. 1813 and 1770 Ma (Fig. 10). Thermobarometric estimates indicate that rocks of the study area were re-equilibrated under conditions of ca. 0.5 GPa and 750–825 °C at this time (Annesley et al., 1992, 1997a, b, c, 1999a, b; Jeanneret et al., 2017; Toma et al., 2024). The M2–D2 temporal interval is supported by geochronological constraints derived from other mineral phases using various U–Pb chronometers, including monazite (Pb diffusion at ca. 500–800 °C; e.g., Cherniak et al., 2004; McFarlane and Harrison, 2006), zircon (>900 °C; Cherniak and Watson, 2001), and titanite (500–800 °C; Kohn, 2017). The distribution of Rb–Sr single-spot ages in metamorphic muscovite is shifted toward the lower limit of this time window (Fig. 10) and reflects the neoformation of muscovite along the retrograde P-T path (≤∼ 700 °C), possibly in the presence of fluids and the cooling of muscovite below the closure temperature of the Rb–Sr system, estimated at ca. 450–600 °C (Dodson, 1973). Above this closure temperature, Sr diffusion remains sufficiently rapid to maintain an open isotope system, allowing equilibration with the surrounding matrix or between minerals.

In contrast, the single-spot ages measured in early hydrothermal muscovite crystallized within anatectic granitoids or within leucosomes of metapelitic gneisses (e.g., Schneider et al., 2007; Skipton et al., 2016; Martz, 2017) range from 1799.5 ± 42.8 to 1701.9 ± 57.6 Ma, with a mode at 1752.1 Ma and an isochron age of 1753.4 ± 12.4 Ma (2σ internal; ±31 Ma including external reproducibility) (Fig. 10). The mean age derived from single-spot analyses, together with the isochron age, falls within the temporal window of the exhumation and orogenic cooling phase of the Trans-Hudsonian orogen, a period characterized by the end of migmatization with the crystallization of leucosomes, the emplacement of leucogranitic intrusions, and the onset of hydrothermal activity associated with the cooling phase. Although its precise duration remains poorly constrained, this cooling interval is typically considered to span from ca. 1770 Ma (750 °C, 0.5 GPa) to ca. 1720 Ma (350 °C, 0.2 GPa), as documented in several studies (Annesley and Madore, 1994; Annesley et al., 1992, 1997a, b, c; Annesley et al., 2005; Jeanneret et al., 2017; Martz, 2017). The end of this period is associated with pronounced retrograde transformations, including the breakdown of cordierite to phyllosilicates, biotite chloritization, sulfide, graphite and muscovite precipitation, quartz-rich fluid circulation, and dissolution-reprecipitation of monazite and zircon (Card, 2012, 2014; Card and Noll, 2016). These processes are further constrained by cooling ages obtained from Rb–Sr biotite (Worden et al., 1985; Schneider et al., 2007), K–Ar muscovite (Philippe et al., 1993), and Ar–Ar muscovite (Alexandre et al., 2009). This timeframe has also been interpreted as reflecting a late thermal event associated with the emplacement of the Kivalliq Igneous Suite (KIS) to the northeast of the Athabasca Basin, with the Wollaston Domain being intruded by the Nueltin Granite of the same suite (ca. 1770–1730 Ma), which may have reset U–Pb metamorphic rutile ages (Adlakha and Hattori, 2021). Additionally, hydrothermal graphite, commonly spatially associated with early hydrothermal muscovite within the Wollaston-Mudjatik Transition shear zones, yields highly precise (<1 %) Re–Os isochron ages of 1731.5 ± 7.4 Ma (2σ; MSWD = 1.3). These ages are interpreted to record graphite formation during the exhumation and orogenic cooling phase of the THO (Martz, 2017) and/or a period of elevated heat flow associated with the KIS emplacement (Adlakha and Hattori, 2021), wherein aqueous fluids mobilized carbon from adjacent wall rocks into semi-brittle to brittle-ductile shear zones (Toma et al., 2022). Overall, the Rb–Sr single-spot ages of early hydrothermal muscovite are consistent with the exhumation and orogenic cooling stage of the THO, during which anatectic granitoids were emplaced and/or magmatic activity of the Kivalliq Igneous Suite occurred, and they further align with the circulation of carbon-bearing aqueous fluids responsible for the formation of hydrothermal graphite. Comparison of the Rb–Sr ages obtained from alteration-free metamorphic and early hydrothermal muscovite with existing literature datasets indicates that the Rb–Sr system constitutes a suitable geochronometer in this geological context, as evidenced by the reproducibility of comparable Rb–Sr ages across four distinct sites within the WMTZ of the THO (Fig. 10).

Variably altered muscovite crystals occurring in illite-sudoite hydrothermal alteration zones in the four study sites along the WMTZ in the eastern Athabasca Basin display highly scattered Rb–Sr single-spot dates and substantial individual uncertainties, ranging from 1756.2 ± 45.8 to 1357.0 ± 64.8 Ma. Despite this variability, muscovite from both altered pelitic and granitic protoliths shows comparable Rb–Sr single spot data distributions. Gaussian Mixture Modelling (GMM) and radial plot analysis define five statistically coherent populations shared by both muscovite generations affected by illite-sudoite alteration. However, their geochronological resolution is intrinsically limited because the analytical uncertainties are large. This reflects the inherent difficulty of precisely dating Paleoproterozoic geological events, as all isotope system tend to yield significant errors at such ancient ages (e.g., Fayek and Riciputi, 2002; Alexandre et al., 2005; Rainbird et al., 2007; Schneider et al., 2007; Davis et al., 2011; Jeanneret et al., 2017; Adlakha and Hattori, 2021; Toma et al., 2022). The identification of five sub-populations does not imply that this is the only possible partitioning, but rather the one favored by the model under the present configuration, given the single-spot age values and their associated uncertainties. Small sub-populations should likewise be interpreted cautiously: although they may correspond to genuine age modes, they may also reflect asymmetries, outliers, or an artificial segmentation of an otherwise continuous distribution. Several studies have demonstrated that Gaussian mixture models can generate artificial components when fitting continuous or asymmetric distributions (McLachlan and Peel, 2000; Bishop and Nasrabadi, 2006; Hastie et al., 2001). When substantial overlap exists between Gaussian components, the model often represents a single underlying population using multiple components (Fraley and Raftery, 2002). Ultimately, the components should be regarded primarily as probabilistic centers, which we attempt to compare with published datasets from the literature (Figs. 8, 10). It should be noted that BIC-based GMM algorithms tend to overestimate the number of populations with increasing sample size (Vermeesch, 2018), and that the present data are equally consistent with a smaller number of broader, partially overlapping distributions. The age components identified here are therefore not interpreted as discrete geological events but as probabilistic centers whose significance is evaluated through their correspondence with independently dated tectono-hydrothermal events documented in the regional record.

A dominant population at ca. 1640 Ma, representing the majority of the dataset and potentially overlooked without prior petrographic constraints, is consistently recorded across four independent sites within the WMTZ of the THO, supporting the regional significance and robustness of this age population. This main cluster is accompanied by subordinate age groups at ca. 1680, ca. 1600, ca. 1550 Ma, and by a minor population near ca. 1467 Ma, which display limited inter-site variability and likely reflect local heterogeneities or superimposed geological processes. These ages fall within the post-Athabasca Basin evolution, following basin deposition between ca. 1760 and 1720 Ma (Ramaekers et al., 2007; Jeanneret et al., 2016) and continuing until ca. 1541 Ma (Creaser and Stasiuk, 2007). The data obtained in strongly illite-sudoite-altered zones of the basement, spatially associated with the uranium mineralization, record extensive mineralogical transformations and major geochemical and isotopic modifications driven by the circulation of basinal brines; i.e., the same fluids responsible for the characteristic illite-sudoite alteration halos observed around all unconformity-related uranium deposits in the eastern Athabasca Basin. The spread of the Rb–Sr single-spot dates may reflect either a single, long-lived fluid-rock interaction event with a paroxysm at ca. 1640 Ma, or a sequence of pulsed hydrothermal episodes represented by the GMM and radial plot populations.

The main peak at ca. 1640 Ma is widely recognized across the basin as marking a major sedimentary reorganization event associated with a rapid westward deepening and the deposition of deeper-water successions (ca. 1664–1590 Ma; e.g., Pehrsson et al., 2023). This structural tilting correlates with the Racklan-Forward orogeny (ca. 1660–1590 Ma; e.g., Furlanetto et al., 2016; Pehrsson et al., 2023), a far-field intracontinental compressional event that reactivated deep crustal faults and deformed the northern Laurentian basement. Within the same time interval, zircon extracted from reworked tuff intraclasts of the Wolverine Point Formation (member b) yields a weighted mean ²⁰⁷Pb / ²⁰⁶Pb age of 1644 ± 13 Ma (2σ), interpreted as the crystallization age of the original tuff layers (Rainbird et al., 2007; Fig. 10). These tuffs are considered to be correlative with the Yavapai and Mazatzal Paleoproterozoic orogens, particularly within the Cochise block of the Mazatzal orogen, where juvenile arc terranes and syntectonic plutons range from ca. 1660 to 1600 Ma (Sims and Peterman, 1986; Karlstrom and Bowring, 1988; Eisele and Isachsen, 2001; Rainbird et al., 2007). In a broader geodynamic framework, this event may also be linked to the Racklan Orogeny of northwestern Laurentia, interpreted as the northwestern continuation of the Mazatzal orogenic system. Together, these orogenic events define a large-scale Paleoproterozoic accretionary system that affected the margins of the Canadian Shield and promoted significant crustal reworking, consistent with the hypothesis of a circum-Laurentian orogenic belt (e.g., Furlanetto et al., 2013). U–Pb data obtained from fluorapatite in the Athabasca Basin indicate a coherent phase of phosphate cementation, constrained by weighted mean ²⁰⁷Pb / ²⁰⁶Pb ages obtained on apatite with high radiogenic-to-common Pb ratios, and by Pb–Pb regression lines that are independent of common Pb (e.g., Davis et al., 2011). Fluorapatite from the Smart, Manitou Falls, and Wolverine Point formations of the Athabasca Basin yields ages of 1609 ± 30, 1638 ± 12, and 1625 ± 17 Ma, respectively, clustering around a mean age of approximately 1630 ± 9 Ma (e.g., Davis et al., 2011; Fig. 10). These ages are interpreted to record a basin-wide circulation of warm, oxidizing, P-rich brines produced by fluid-rock interactions between basinal brines and the altered crystalline basement during mesogenetic diagenesis. This large-scale fluid circulation is likely tectonically driven and related to the regional geodynamic evolution of the basin, as proposed by Pehrsson et al. (2023), and can be temporally constrained to ca. 1640 Ma based on the concordance between Rb–Sr and U–Pb apatite ages.

Taken together, these observations support the interpretation of a major basin-wide hydrothermal event at ca. 1640 Ma. Illite-sudoite alteration dated at ca. 1640 Ma can be interpreted in two alternative, but not mutually exclusive, ways. First, this alteration may predate the main stage of uranium oxide precipitation, commonly dated between ca. 1590 and 1200 Ma based on U–Pb analyses of uraninite and K–Ar/Ar–Ar geochronology of clay minerals (Cumming and Krstic, 1992; Philippe et al., 1993; McGill et al., 1993; Fayek and Riciputi, 2002; Alexandre et al., 2009; Cloutier et al., 2011; Powell et al., 2022). In this scenario, the ca. 1640 Ma illite-sudoite alteration would represent a basin-scale hydrothermal event that prepared the basement-basin interface for subsequent uranium mineralization by enhancing permeability, redox reactivity, and uranium mobility. Alternatively, the 1640 Ma alteration may record an early phase of uranium mineralization itself, subsequently overprinted and partially reset by younger tectono-hydrothermal events. Such a model is consistent with studies demonstrating that the U–Pb systematics of uraninite and associated alteration minerals in unconformity-related uranium deposits are commonly affected by multiple fluid-flow episodes, leading to partial Pb loss and a wide dispersion of apparent ages (Kotzer and Kyser, 1990; Fayek and Kyser, 1997; Fayek et al., 2002; Fayek and Riciputi, 2002; Kyser et al., 2015). Consequently, an older hydrothermal and/or mineralizing system active during the ca. 1680–1640 Ma interval cannot be excluded, but may remain cryptic due to subsequent uranium remobilization and isotopic resetting during younger basin-wide fluid circulation events. This interpretation is further supported by analogues from northern Australia, where in situ U–Pb analyses of uraninite indicate initial mineralization or hydrothermal alteration as early as ca. 1680–1640 Ma, followed by multiple resetting episodes extending into the Mesoproterozoic and Paleozoic (Clauer et al., 2015; Skirrow et al., 2016).

In addition to the main population associated with illite-sudoite alteration at ca. 1640 Ma, several consistent statistical sub-populations are identified by the GMM. However, no chemical or mineralogical modifications are observed in the zones where these single-spot data were measured. A first sub-population calculated around ca. 1680 Ma can be interpreted in two ways. It may represent a mixed population between preserved and altered data (Fig. 6F), reflecting the onset of alteration and a partial loss of radiogenic ⁸⁷Sr caused by this early alteration stage at ca. 1640 Ma. Such mixing would tend to skew the single-spot Rb–Sr dates of metamorphic muscovite, which crystallized prior to early hydrothermal muscovite, toward older dates. Alternatively, this ca. 1680 Ma population could correspond to an argillization episode dated to this period by Alexandre et al. (2009) using Ar–Ar analyses on basement-hosted illite, which yields pseudo-plateau ages at McArthur River of 1669 ± 4, 1666 ± 6, and 1671 ± 7 Ma, interpreted as a pre-ore alteration event (Fig. 10). However, the actual geological significance of each published age and its attribution to a specific event are often difficult to evaluate in this context, due to successive episodes of intense fluid circulation and associated fluid-rock interaction processes that promote substantial Ar loss and consequently yield younger ages (Chi et al., 2018). Nevertheless, during this time interval, the basin is characterized by a half-graben architecture, with sediment transport predominantly directed toward the northwest and enhanced sediment accumulation along the southeastern basin margin. This configuration reflects an intracontinental extensional regime established after the THO orogeny, which may have facilitated the circulation of early fluids within the basement and may also have reset the Rb–Sr system.

The younger age populations at ca. 1600–1550 Ma are consistent with hydrothermal and mineralizing events that have been independently documented throughout the Athabasca Basin. The sub-populations at ca. 1600 Ma and ca. 1550 Ma document hydrothermal activity previously identified in the Athabasca Basin and its underlying basement, although it remains unclear whether they reflect a single event, a series of pulsatile episodes, or processes potentially related to, or distinct from, the ca. 1640 Ma event. This hydrothermal activity is highlighted in the basement by anatase precipitation, a low-pressure, low-temperature TiO2 polymorph, yielding a weighted mean ²⁰⁷Pb / ²⁰⁶Pb age of 1569 ± 31 Ma (MSWD = 0.30, n = 5, 2σ; Adlakha and Hattori, 2021, Fig. 10). This age has been interpreted as reflecting the onset of oxidizing hydrothermal activity in the basement, contemporaneous with U precipitation at the world-class McArthur River deposit, which has been dated by U–Pb analyses of U oxides with two discordant ages of 1540 ± 19 Ma (Alexandre et al., 2009) and 1540 ± 38 Ma (Alexandre et al., 2005). Additional ages at ca. 1520 Ma, including 1514 ± 18 Ma (Cumming and Krstic, 1992), 1519 ± 22 Ma (Fayek and Riciputi, 2002), and 1521 ± 18 Ma (McGill et al., 1993), have been interpreted as minimum crystallization ages for primary uraninite. An episode of argillization contemporaneous with these fluid-circulation events has also been identified at 1577 ± 38 Ma (K–Ar muscovite from gneiss; Philippe et al., 1993) at Cigar Lake, and at 1533 ± 13 Ma (Ar–Ar on illite; Alexandre et al., 2009) at McArthur River. A further rare-earth-phosphate hydrothermal episode is documented in the eastern Athabasca Basin within the Maw Zone rare-earth-element deposit. U–Pb dating of xenotime yields a ²⁰⁷Pb / ²⁰⁶Pb age of 1547 ± 14 Ma (Rabiei et al., 2017; Fig. 10), and both isotopic signatures and fluid-inclusion characteristics support the interpretation that the fluids responsible for xenotime precipitation were part of the same oxidizing-reducing hydrothermal system involved in the development of major uraninite mineralization across the Athabasca Basin. More broadly, these events are associated with a renewed phase of subsidence and marine flooding, attributed to intracontinental rifting or extension and linked to the emplacement of the Kuungmi Formation basalts in the adjacent Thelon Basin. This rifting episode coincides with a reorganization of the Laurentian crust following the Australia-Laurentia collision, and precedes the stabilization of Nuna, potentially promoting large-scale fluid circulation through both the basement and the basin, capable of resetting the Rb–Sr geochronometer in muscovite.

Finally, the minor age population at ca. 1467 Ma may correspond to an early mineralization event, as recorded by discordant U–Pb ages of 1467 ± 47 Ma at the Cigar Lake deposit (Fayek and Riciputi, 2002) and by weighted mean ²⁰⁷Pb / ²⁰⁶Pb ages of 1458 ± 12 Ma at the McArthur River deposit (Kister, 2003). An argillization event contemporaneous with these fluid-circulation episodes has also been identified at 1456 ± 32 Ma (K–Ar on the < 2 µm fraction from gneiss; Philippe et al., 1993) at Cigar Lake, and at 1433 ± 31 Ma (Ar–Ar on illite; Alexandre et al., 2009) at McArthur River. Overall, current data do not allow discrimination between a single protracted fluid-rock interaction episode at ca. 1640 Ma leading to the formation of an illite-sudoite alteration halo and multiple episodic fluid pulses to account for the Rb–Sr single-spot age distribution in altered metamorphic and early hydrothermal muscovite. Nevertheless, these results clearly demonstrate the applicability of the Rb–Sr chronometer for constraining hydrothermal fluid circulation at basin-basement interfaces. The distinct age populations identified in muscovite and alteration products correspond to tectono-hydrothermal events that are independently recognized in the regional geological record and have been previously dated using other geochronological methods. Comparable multi-stage fluid-thermal histories have been documented in other hydrothermal systems, such as gold deposit systems, where primary mineral phases commonly preserve evidence for repeated isotopic resetting, trace-element redistribution, and microstructural overprinting associated with episodic fluid and heat input (e.g., Bevan et al., 2021; Chen et al., 2024; Drake et al., 2023; Huang et al., 2023; Liu et al., 2018; Mao et al., 2013; Olierook et al., 2020; Redaa et al., 2021, 2022; Şengün et al., 2019; Wang et al., 2022; Zametzer et al., 2022; Tian et al., 2024; Boschetti et al., 2025; Huang et al., 2025). Such behaviour underscores the capacity of mineral-scale chronometers to resolve protracted hydrothermal evolution. In this context, the present data suggest that the Rb–Sr system in muscovite and alteration minerals is sensitive to successive hydrothermal events at basin-basement interfaces, highlighting its potential as a robust tool for reconstructing the timing and duration of fluid circulation processes, while emphasizing the need for further work to evaluate mineralogical and geochemical controls on system behaviour. Although these initial results are promising, further work is required to assess potential mineralogical or geochemical specificities that may influence the Rb–Sr system under such conditions.