# Uranium mobility and deposition over 1.3 Ga in the Westmoreland area (McArthur Basin, Australia)

Joséphine Gigon, Julien Mercadier, Irvine Annesley, Antonin Richard, Andrew

S. Wygralak, Roger G. Skirrow, Terrence Mernagh, Ion Probe Team Nancy

► To cite this version:

Joséphine Gigon, Julien Mercadier, Irvine Annesley, Antonin Richard, Andrew S. Wygralak, et al. Uranium mobility and deposition over 1.3 Ga in the Westmoreland area (McArthur Basin, Australia). Mineralium Deposita, 2021, 56 (7), pp.1321-1344. 10.1007/s00126-020-01031-2. hal-04772490

HAL Id: hal-04772490

https://hal.science/hal-04772490v1

Submitted on 13 Feb 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
Uranium mobility and deposition over 1.3 Ga in the
Westmoreland area (McArthur Basin, Australia)

Joséphine Gigon¹, Julien Mercadier¹, Irvine R. Annesley¹, Antonin Richard¹, Andrew S. Wygralak², Roger G.

Skirrow ³, Terrence P. Mernagh ⁴, Ion Probe Team Nancy ⁵

¹Université de Lorraine, CNRS, CREGU, GeoRessources, F-54500 Vandœuvre-lès-Nancy, France

²Northern Territory Geological Survey, PO Box 2901, Darwin, NT 0801, Australia

³Geoscience Australia, GPO Box 378, Canberra, ACT 2601, Australia

⁴Australian National University, 142 Mills Rd., Acton, ACT 2601, Australia

⁵CRPG, UMR 7358, CNRS, Université de Lorraine, Vandœuvre-lès-Nancy F-54501, France

Corresponding author: josephine.gigon@gmail.com; +33 3 72 74 55 59

JG: https://orcid.org/0000-0001-5971-9659

RS: https://orcid.org/0000-0002-9684-1554

TM: https://orcid.org/0000-0002-8555-0110

Abstract

The Westmoreland area is located south of the McArthur Basin (Australia) and hosts several uranium (U) deposits and prospects. Previous studies proposed that U mineralisation formed under conditions similar to the unconformity-related U deposits of the Alligator Rivers Uranium Field (ARUF), located north of the same sedimentary basin. Detailed mineralogical, geochemical and geochronological studies on mineralised intervals from the Redtree, Junnagunna and Huarabagoo deposits, and integration with previous studies, identified at least six generations of uranium oxides. These formed between ca. 1680 and 350 Ma, highlighting protracted mobility of U over 1.3 Ga. Each generation of uranium oxide has a specific chemical composition, indicating variable physico-chemical conditions for their formation throughout this period.

Although the 1680 Ma mineralising event formed at the same time and with similar mineralising fluids as the
unconformity-related U deposits from the ARUF, the physico-chemical conditions differed between the two
areas. Deposits of the Westmoreland area lack the typical Mg- and B-metasomatism of the unconformity-related U deposits in the ARUF: in the Westmoreland area, the chlorite is Mg-poor and no B minerals are observed, whereas the ARUF is marked by Mg-rich chlorite (between the clinochlore and Mg-amesite fields), Mg-foitite (tourmaline), and alumino-phosphate-sulphate minerals. The chondrite-normalized REE patterns of uranium oxides are enriched in LREE or flat, compared with the typical bell-shape of the unconformity-related U deposits. Moreover, chlorite thermometry indicates significantly higher temperature conditions (>300°C) than in the ARUF for the early U stages. Based on these results, it is thus unlikely that the mineralisation in the Westmoreland area can be linked to significant unconformity-related U deposits in this area. Collectively, the data point towards atypical ore-forming processes for basin-related U deposits in the Westmoreland area compared with known processes in the ARUF and in other sedimentary basins.

## Keywords

Uranium, Australia, Westmoreland, U-Pb isotopic ages, uraninite REE patterns, chlorite thermometry

## 1. Introduction

The Westmoreland area (Northern Territory and Queensland, Australia) hosts several U deposits and prospects located in the Paleo- to Mesoproterozoic McArthur Basin and related crystalline basement (Polito et al. 2005a; Ahmad et al. 2013; International Atomic Energy Agency 2018a, Figs. 1 and 2). The area hosts five deposits (Redtree, Junnagunna, Huarabogoo, Outcamp and Sue) that are classified as Westmoreland-Murphy-type U deposits (Lally and Bajwah 2006), and have indicated combined resources for all the deposits totalling ~ 16 000 t U contained within 18.7 million tonnes at an average grade of 0.089% U₃O₈, and an inferred mineral resource totalling ~ 7 000 t U contained within 9.0 million tonnes at an average grade of 0.083% U₃O₈ (JORC and NI 43-101 compliant resource, Laramide Resources Ltd 2019). These deposits and related occurrences are part of a set of deposits known in this large intracratonic basin, with several mineralised areas concentrating various types of metals (U, Pb, Zn, and Cu, Ahmad et al. 2013). U deposits are found in the Pine Creek Orogen (northern part of the basin) and in the Westmoreland area (southern part), Cu-bearing breccia pipes and minor vein-type mineralisation are present in the Redbank area in the southeastern part of the basin (e.g. Knutson et al. 1979), and SHMS Zn-Pb deposit are mostly found in the centre of the basin (e.g. Gigon et al. 2020).
The geological processes and physico-chemical conditions for the origin of the different U deposits found in the McArthur Basin have been studied for decades. The Alligator Rivers Uranium Field (ARUF) to the North hosts the Jabiluka, Ranger, Nabarlek and Koongarra U deposits among others belonging to the unconformity-related deposits type, for which the processes related to the U deposition have been studied extensively (e.g. Wilde and Wall 1987; McKay and Miezitis 2001; Derome et al. 2003; Polito et al. 2004, 2005b, 2011; Lally and Bajwah 2006; Fisher et al. 2013; Clauer et al. 2015; Skirrow et al. 2016; Gigon et al. 2019). U deposits from the Westmoreland area have comparatively attracted much less attention (Hills and Richards 1976; Fuchs and Schindlmayer 1981; Ahmad and Wygralak 1989 and references herein; Rheinberger et al. 1998; Polito et al. 2005a; Mernagh and Wygralak 2011), and the previous research (e.g. Polito et al., 2005a and Mernagh and Wygralak 2011) mainly examined if these U deposits and related occurrences could be analogues to the unconformity-related U deposits known at the border of the Northern part of the McArthur Basin, but also in the Athabasca and Thelon basins (Canada; Jefferson et al. 2007; Cuney and Kyser 2015; Grare et al. 2020). Unambiguously establishing such relationships with unconformity-related U deposits is indeed important because unconformity-related U deposits have the highest grades and tonnages among all the U deposits, and are therefore dominant U exploration target worldwide.

The results from the previous studies in the Westmoreland area indicate that the mineralogy, paragenesis and geochemistry of the Redtree, Junnagunna and Huarabagoo deposits share some characteristics with the U deposits of the ARUF (Polito et al. 2005a), thus suggesting that the timing and geological processes for the origin of both U provinces were similar. The fluid inclusions in the Eva Mine and Jackson Pit U deposits, yielded NaCl-rich and CaCl₂-rich brines and a low-salinity fluid that mixed in varying degrees (Mernagh and Wygralak 2011), as also documented for the unconformity-related U deposits from the ARUF (Ypma and Fuzikawa 1980; Durak et al. 1983; Derome et al. 2003, 2007; Polito et al. 2004, 2005b). These combined results seem to indicate that the first-order parameters and the timing for the formation of both uranium provinces in the McArthur Basin are identical. The main goal of the present work is thus to refine this comparison between the U deposits of the ARUF and the Westmoreland area by verifying the possible metallogenetic relationships between these two deposit subtypes.

New geochemical criteria applicable to uranium oxides, based on rare earth element content in uranium oxides, allow classification of the different types of U deposits (Mercadier et al. 2011b; Frimmel et al. 2014; Alexandre et al. 2016). Such criteria have been applied to the U deposits from the ARUF, confirming their affinity with the U deposits from the Athabasca Basin (Canada) and thus their classification as unconformity-related U deposits
(Mercadier et al. 2011b). Application of such geochemical markers is lacking for the U deposits of the Westmoreland area and this forms the first objective of this study. The second objective is to add to the existing data on ore mineralogy, geochemistry and age dating for different mineralised intercepts of three U deposits from Westmorland area: Redtree, Junnagunna and Huarabagoo. The final objective of this approach is to provide new geochronological constraints on formation of the U deposits and occurrences of the Westmoreland area, in particular the characterisation of potential multiple fluid events that favoured U mobility through time and space. These new results are compared to previously published models for the area, to the unconformity-related deposits of the ARUF, and to U deposits in other Proterozoic basins.

## 2. Geological settings

The Westmoreland area is adjacent to the Queensland-Northern Territory border on the south-eastern margin of the southern McArthur Basin (Fig. 1). The oldest rocks in the Westmoreland region are Paleoproterozoic metasedimentary, volcanic and felsic rocks from the Murphy Province that forms the basement to the McArthur Basin. The units in the Murphy Province are the Murphy Metamorphics and a suite of comagmatic felsic intrusive and volcanic rocks known as the Nicholson Granite and the Cliffdale Volcanics. The Murphy Metamorphics are essentially a succession of greenschist-facies metamorphosed shales and greywackes (Ahmad et al. 2013). Uranium-lead data on detrital zircon suggest a maximum depositional age of 1853 ±4 Ma (Hollis et al. 2010, Table SM1) and the sediments were metamorphosed prior to the extrusion of the Cliffdale Volcanics (Ahmad et al. 2013). The Nicholson Granite Complex (Fig. 2) is intrusive in the Murphy Metamorphics and can be divided into two broad suites based on chemical and mineralogical composition, one hornblende- and/or biotite bearing granite, quartz monzonite and granodiorite and one biotite- and/or muscovite bearing monzonite, granite and alkali granite. Two samples of the Nicholson Granite Complex have yielded SHRIMP U-Pb zircon ages of 1856 ±3 and 1845 ± 3 Ma (Page et al. 2000, Table SM1). The Cliffdale Volcanics are a succession of more than 4km of rhyolite, dacitic lavas, ignimbrites and intrusions that unconformably overlie the Murphy Metamorphics (Ahmad et al. 2013). SHRIMP U-Pb dating of zircons has yielded an age of 1851 ± 3 Ma (Page et al. 2000, Table SM1).

The Paleo-Mesoproterozoic McArthur Basin is an up to 10 km thick sequence of mostly unmetamorphosed sedimentary and volcanic rocks deposited between 1815 and 1492 Ma that unconformably overlies the Murphy Province (Ahmad et al. 2013). The basal unit of the McArthur Basin, the Westmoreland Conglomerate (Fig. 2), is composed mainly of fluvial sandstones and conglomerates and is up to 1,800 m thick. It has been subdivided into five stratigraphic units, each unit representing a major fining-upward cycle of sedimentation (Ahmad and Wygralak 1989). Each unit comprises proximal fluvial deposits typical of debris flows, alluvial fans, and braided river systems overlain by medium- to coarse-grained, well-sorted sandstones. Angular unconformities or
disconformities define breaks in sedimentation, with each new cycle of pebble or boulder conglomerate commonly defining the beginning of a new unit. Cobbles and coarse sand grains within the basal conglomerate are dominated by reworked quartz veins, chert, and clasts of felsic to mafic volcanic rocks that were likely derived from basement rocks such as the Murphy tectonic ridge or similar rocks that once existed to the North (Wygralak et al. 1988; Ahmad and Wygralak 1989; Polito et al. 2005a). The detrital lithic clasts, monazite, zircon, and apatite are likely to be among the source of uranium that was dissolved during peak diagenesis (Polito et al. 2005a and references herein). Most of the uranium mineralisation is hosted within the upper unit of the Westmoreland Conglomerate (Ptw4), which is a 80–90 m thick porous, coarse-grained sandstone, conglomeratic in part (International Atomic Energy Agency 2018a). U-Pb dating of detrital zircon from the Westmoreland Conglomerate has returned maximum depositional ages of 1865 ± 7 Ma for the lower section and 1843 ± 4 Ma for the upper section (Ahmad and Wygralak 1990, Table SM1). The Westmoreland Conglomerate is conformably overlain by the mafic Seigal Volcanics (basaltic lavas), followed by dolomite, sandstone, mafic and felsic volcanic rocks of the upper part of the Tawallah Group. The deposition age of the Seigal Volcanics is estimated at ca. 1780 Ma by stratigraphic correlation (Scott et al. 2000; Mernagh and Wygralak 2011).

Dolerite dykes and minor sills, which may have been feeders of the lava flows, intrude along North-Est trending fault and fracture zones that intersect the Westmoreland Conglomerate (International Atomic Energy Agency 2018b). Uranium mineralisation is associated with three major northeast-trending lineaments (NE Westmoreland, Redtree and El Nashfa dyke zones, Fig. 1), with the Redtree dyke zone extending over a strike of 20 km (Ahmad et al. 2013). The dykes has not been dated but they could have the same age as the Seigal Volcanics (Ahmad and Wygralak 1989).

The Westmoreland region was first prospected in the 1890s for Pb-Zn, then for Cu, and finally U was mined from the Peters Creek Volcanics which overlie the Westmoreland Conglomerate, 20 to 30 km west of Redtree (Laramide Resources Ltd 2016). The Redtree, Junnagunna and Huarabagoo U deposits are hosted largely within the shallow dipping Westmoreland Conglomerate near the unconformity with the Murphy Inlier and lie along the Redtree dyke zone (Fig. 1). The U-rich zones can be oriented parallel to the volcanic rocks or sills, discordant at the contact with mafic dykes, within fractures cross-cutting altered mafic volcanic rocks within fractures or along shear zones in altered felsic volcanic rocks. The studied deposits belong to “mafic dykes/sills in Proterozoic sandstone type” (International Atomic Energy Agency 2018b) or “Westmoreland-Murphy-type” (Lally and Bajwah 2006). The deposits of the “mafic dykes/sills in Proterozoic sandstone” category are situated in the Westmoreland area and within the Otish Basin (Canada). “Westmoreland-Murphy-type” deposits occur
within conglomerate, sandstone and mafic volcanic rocks, and are all associated with redox boundaries near the contacts between different lithologies. In addition, in the studied region, uranium occurrences were classified by Ahmad (1987) into five types (A–E), based on their hydrological and geological settings. Four of them are present in the basin, and reported on Fig. 1. Type A deposits lie at the contact between the Westmoreland Conglomerate and the Cliffdale or Seigal Volcanics (in sub-type A1 occurrences, the contact between the Cliffdale Volcanics and Westmoreland Conglomerate is a reverse fault; in sub-type A2 deposits, the contact between the Westmoreland Conglomerate and Seigal Volcanics is conformable). In type B deposits, mineralisation occurs as sub-horizontal and sub-vertical lenses in the Westmoreland Conglomerate. Type C deposits are hosted by intensely altered Cliffdale Volcanics; Type D occurrences are related to fractures in the lower part of the Seigal Volcanics.

The bulk of the Westmoreland resource base (~50%) is located in the Redtree deposit. It flanks the Redtree dyke zone immediately north of the northwest-trending Namalangi fault (Fig. 1). The Redtree dyke (20–30 m wide) broadly extends for 7 km from Redtree to Junnagunna and spatially controls the distribution of U, indicating that mineralizing fluids used this structure as a main drain uranium mineralising fluids probably used this structure (Laramide Resources Ltd 2019). The deposit is characterised by a mix of horizontal and vertical lenses of continuous mineralisation in distinctly coarse pebble conglomerates with grades ranging from 0.15% to >2% $U_3O_8$ (Fig. 2). Mineralisation is generally shallow, within 10m of surface and is associated with chlorite/hematite alteration. The mineralisation thickens and steepens near the dyke, where it is 30–40 m thick (International Atomic Energy Agency 2018a). The indicated resource is 12.86 Mt of ore at an average grade of 0.09% $U_3O_8$ and an inferred resource of 4.47 Mt at an average grade of 0.07% $U_3O_8$ (JORC and NI 43-101 compliant, Laramide Resources Ltd 2016).

The Junnagunna U deposit occurs at a fault intersection west of the Redtree dyke zone and south of the northwest trending Cliffdale fault (Fig. 1). Mineralisation is generally flat lying on both sides of the Redtree dyke and ranges from 0.5 m to 20 m in thickness beneath the Seigal-Westmoreland contact (Fig. 2). The deposit is covered by alluvium/colluvium clays, sand and weathered Seigal Volcanics (Laramide Resources Ltd 2019). The indicated resource is 4.36 Mt ore at an average grade of 0.08% $U_3O_8$ and an inferred resource of 2.15 Mt at an average grade of 0.8% $U_3O_8$ (Laramide Resources Ltd 2016).

The Huarabagoo deposit is about 3 km NE of Redtree along the Redtree dyke zone and straddles the contact of the Seigal Volcanics with the Westmoreland Conglomerate (Figs. 1 and 2). Mineralisation occurs in 20 m thick
lenses, which extend to 100–200 m length and to a depth of 80 m (Laramide Resources Ltd 2019). It outcrops at the southern end and is concealed to the north under 2 to 3 m of sandy alluvium and 5 to 8 m of weathered basalt of the Seigal Volcanics (Laramide Resources Ltd 2016). The deposit comprises a 3 km zone of vertical mineralisation associated with a complex dyke geometry with vertical and horizontal branches between the two principal dykes. The strongest mineralisation is associated with chlorite/hematite-altered coarse pebbly sandstones similar to Redtree. Most of the mineralisation is hosted within the Westmoreland Conglomerate adjacent to the dykes, and the remainder is hosted in the dykes (International Atomic Energy Agency 2018b). Indicated resources are1.46 Mt ore at an average grade of 0.08% U₃O₈ and inferred resources are 2.41 Mt at an average grade of 0.11% U₃O₈ (Laramide Resources Ltd 2016).

### 3. Sampling and methodology

A multi-method approach was conducted on different minerals (uranium oxides, chlorite and apatite), from the Junnagunna, Redtree and Huarabagoo deposits. A specific focus was the selection of samples within the highest-grade parts of mineralised zones of the deposits, all located within the Westmoreland Conglomerate (Fig. 2). This selection was also based on the paragenesis for the U deposits of the Westmoreland area established by Polito et al. (2005a).

Five drill holes were sampled in three deposits based on their high U grade, representative of the different mineralisation styles: JDD08-003 (Junnagunna), WDD9-128 (Huarabagoo), WDD10-151 (Huarabagoo), WDD07-2 (Redtree), and WDD08-106 (Redtree). Mineralised intersections are: at Junnagunna, 0.21% U₃O₈ in JDD08-003 from 71 to 76m; at Huarabagoo, 0.16% U₃O₈ in WDD10-151 between 22 and 31m; at Redtree, 0.25% U₃O₈ in WDD07-02 from 48 to 59m and 0.2% from 70 to 75m and 0.31% in WDD08-106 from 8 to 16m (Vigar and Jones 2009; Laramide Resources Ltd 2016). The following samples were studied: JDD08-003 at 74.4m, WDD07-2 at 51.7, 55.9 and 73.5m, WDD9-128 at 53.3 and 92.3m, WDD10-151 at 29.4m and WDD08-106 at 10.6m.

Selected thin sections were studied by optical and scanning electron microscope (SEM) to develop detailed paragenetic sequences. The chemical composition for the major elements of chlorite and uranium oxide was determined by Electron Probe Micro Analyses (EPMA). The isotopic U-Pb and Pb/Pb ages of uranium oxide and apatite were determined by Secondary Ion Mass Spectrometry (SIMS). The concentrations of minor and trace
elements in uranium oxides were determined by Laser Ablation-Inductively Coupled Plasma-Mass Spectrometer (LA-ICP-MS). The locations of the analyses are presented in Figs. SM1–SM8.

## 3.1. Petrography

Optical microscopy was carried out on thin sections using a numerical microscope VHX-200F with a high resolution objective VH-Z500W to cover the whole sections and a Zeiss Axio Scope A1 microscope for detailed petrography (GeoRessources Lab - Vandœuvre-lès-Nancy, France). SEM images and spectra were obtained using BackScattered Electron detector (BSE) Tescan Vega3LM with a conventional tungsten source at SCMEM (GeoRessources Lab - Vandœuvre-lès-Nancy, France). The analytical conditions were an acceleration voltage of 15 kV and beam current of around 0.1–1 nA.

## 3.2. Major elements composition of uraninite and chlorite

The quantitative measurements of major elements in uraninite and related chlorite were performed using a CAMECA SX-100 Electron Probe MicroAnalyser (EPMA) at SCMEM with a spot size of 1µm². For uraninite, the calibration was based on natural and synthetic oxides and/or alloys (V for V, andradite for Ca, hematite for Fe, albite for Si, apatite for P, zircon for Zr, UO₂ for U, MnTiO₃ for Mn and Ti, Al₂O₃ for Al, ThO₂ for Th, PbS for Pb and YPO₄ for Y). The analytical conditions were 20 nA current, accelerating voltage of 20 kV and counting time of 10 s. For chlorite, the calibration was also based on natural and synthetic oxides and/or alloys (V for V, andradite for Ca, orthose for K, hematite for Fe, albite for Si and Na, olivine for Mg, MnTiO₃ for Mn and Ti, Al₂O₃ for Al, and Cr₂O₃ for Cr). The analytical conditions were 12 nA current, accelerating voltage of 15 kV and counting time of 10 s. The limits of detection of each analysed element are presented in Table SM2 (uraninite) and Table SM12 (chlorite).

## 3.3. U-Pb isotopic ages of uranium oxides and apatite

SIMS is an established approach to date apatite (Li et al. 2012; Zhou et al. 2018) and uranium oxide (Fayek et al. 2000, 2002; Luo et al. 2015; Martz et al. 2019). The U-Pb and Pb/Pb analyses were carried out at the Centre de Recherches Pétrographiques et Géochimiques (CRPG, Nancy, France) with the Cameca IMS 1280 HR2 and the Hyperion RF O⁻ source for apatite and with the Cameca IMS 1270 E7 and the duoplasmatron O⁻ source for uranium oxide. The secondary positive ions were measured with a mass resolution (M/ΔM) of 6000 to separate peaks of interest from isobaric interferences, by ion counting in mono-collection mode on the axial electron multiplier.
For apatite, the primary O⁻ beam was set at an intensity of 30nA with a 10 µm spot size. Each measurement consisted of a 120 s pre-sputtering, secondary beam and mass centring and 16 cycles of measurement of masses: 178.5 (background), ⁴⁰Ca⁴⁴Ca³¹P¹⁶O₄, ²⁰⁴Pb, ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb, ²³⁸U, ²³²Th¹⁶O and ²³⁸U¹⁶O. Counting times were 4 s per peak for ²³²Th¹⁶O and ²³⁸U¹⁶O, 6 s per peak for ²³⁸U, 8 s per peak for ²⁰⁸Pb and 18 s per peak for ²⁰⁴Pb, ²⁰⁶Pb and ²⁰⁷Pb. An energy window of 30 eV was used with no energy offset. ⁴⁰Ca⁴⁴Ca³¹P¹⁶O₄ was used as a reference mass to calculate the U and Pb contents. The Madagascar apatite was used as standard to determine the relative ionization yield of U and Pb, with the Pb/U ratio expressed as a function of UO/U ratio. The external error on the ²⁰⁶Pb/²³⁸U ratios, measured on the Madagascar standard (dated at 473.5 Ma, Chew et al. 2019) for each session, was propagated on the final error of each sample measurements. With the measurement of the ²⁰⁴Pb isotope it is possible to calculate ²⁰⁶Pb/²³⁸U corrected ages using the measured ²⁰⁴Pb/²⁰⁶Pb ratios and the continental common lead compositions from Stacey and Kramers (1975) at the age of crystallisation. This allows us to calculate probability density plots and weighted average plots using ²⁰⁶Pb/²³⁸U corrected ages, and compare them to the common lead uncorrected Tera-Wasserburg intercept ages (Tera and Wasserburg 1972). All data were treated using ISOPLOT software (Ludwig 2007).

For uranium oxides, the primary O⁻ beam was set at an intensity of ~5 nA and focussed to a spot of ~15 µm in diameter, in Gaussian mode with a raster of 10µm. The energy window was opened at 30 eV and centred on the low energy side, 5 eV below the maximum value. An offset of -50eV was applied during analyses in order to avoid matrix effects due to mixes between uraninite and silicate in natural samples. Ions were measured by peak jumping in monocollection mode using the axial Faraday cup (FC) for ²³⁸U and ²³⁸UO and the axial electron multiplier (EM) for ²⁰⁴Pb, ²⁰⁶Pb, ²⁰⁷Pb, ²⁰⁸Pb and ²³²ThO. Each analysis consisted of 8 successive cycles. Each cycle began with measurement of the mass 203.5 and 203.6 for backgrounds of the FC and the EM. The counting time was 4 s for 203.5, 203.6, ²⁰⁸Pb and ²³⁸UO; 3 s for ²⁴⁸ThO; 6 s for ²⁰⁶Pb and ²³⁸U; 10 s for ²⁰⁴Pb and 20 s for ²⁰⁷Pb (waiting time of 1 s). The standard used was a uraninite from Zambia (Holliger 1988), which has a given concordant age at 540 ± 4 Ma (Cathelineau et al. 1990). Ages and error correlations were calculated using ISOPLOT (Ludwig 2007). Uncertainties in the ages are reported at the 2σ level. Common Pb corrections were based on the measured ²⁰⁴Pb content using the Pb isotope composition calculated from Stacey and Kramers (1975) at the inferred age of the uranium oxide crystallisation.

### 3.4. Minor and trace element concentrations in uranium oxides
Minor and trace elements in uranium oxides, including rare earth elements (REEs), were measured with a Laser Ablation-Inductively Coupled Plasma-Mass Spectrometer (LA-ICP-MS) at the GeoRessources laboratory (Nancy, France). The uranium oxides were ablated using a 193 nm Geolas Q Plus system with an aperture-imaged Complex 103 ArF excimer laser (Microlas, Göttingen, Germany) equipped with beam homogenisation optics. The laser spot sizes were 16 and 24 µm and the fluence was ~5 J/cm² with a repetition rate of 5 Hz. Ablated particulate material was analysed by an Agilent 7500c Quadrupole ICP-MS. The transport gas was a mixture of He and Ar, at typical flow rates of 0.5 and 0.8 1.min⁻¹ respectively. All data were acquired in raw counts using time-resolved mode. Data reduction was done using Iolite software (Paton et al. 2011). Absolute concentrations and limits of detection were obtained from the equation developed by Longerich et al. (1996). The calibration material (external standard) was a SRM NIST 610 glass standard and SRM NIST 612 glass was analysed to control the precision and accuracy of the measurements (concentrations from Jochum et al. 2011). The methodology was modified from Lach et al. (2013). The matrix effects when quantifying the REEs in a uraninite using SRM NIST 612 are limited (< 15%; Lach et al. 2013) and are considered to be in the same range for the other measured minor and trace elements in the present study, as previously proposed by Frimmel et al. (2014) and Alexandre et al. (2016). A series of 37 elements were analysed for each experiment: ¹¹B, ⁴⁷Ti, ⁵¹V, ⁵³Cr, ⁵⁵Mn, ⁵⁹Co, ⁶⁰Ni, ⁶³Cu, ⁶⁶Zn, ⁶⁹Ga, ⁷⁵As, ⁸⁸Sr, ⁸⁹Y, ⁹⁰Zr, ⁹³Nb, ⁹⁵Mo, ¹²¹Sb, ¹³⁷Ba, the 14 REEs, ¹⁸¹Ta, ¹⁸²W, ²⁰⁹Bi, ²³²Th and ²³⁸U. ²³⁸U was the internal standard measured by EPMA and the used concentration was adjusted depending of the analysed sample: 77.1 wt% U for WDD07-2 55.9A, 76.8 for WDD07-2 55.9B, 75.6 for WDD10-151 29.4, 75.2 for WDD08-106 10.6, 74.9 for WDD07-2 73.5B and 72.1 for JDD08-003 74.4B. The analytical precision calculated for all elements using standard deviation of recorded intensity is better than ± 20%. The limit of detection (LOD) varies for each element and each analysis as a function of signal intensity and was calculated for each ablation from signal intensities using the 3σ criterion (Longerich et al. 1996). The LODs were reported in Table SM3 when the calculated values for a given element fell below these values. Values used to calculated (ΣLREE/ΣHLREE)N are chondrite-normalised with chondrite values from Anders and Grevesse (1989) and the following elements were used for calculation: ΣREE=La+Ce+Pr+Nd+Sm+Eu+Tb+Dy+Ho+Er+Tm+Lu;

(ΣLREE/ΣHLREE)N=(LaN+CeN+PrN+NdN+SmN;)/(ErN+TmN+LuN)..

Correlations matrices (Tables SM4 – SM9) were generated using the ioGAS software. The chosen criteria for a positive correlation between two elements was a correlation coefficient higher than 0.9, and lower than -0.9 for an anti-correlation.
284

# 4. Results

285

## 4.1. Paragenesis

The paragenesis presented in this section is based only on the observations from the samples for this study. A combined paragenesis from this study and Polito et al. (2005a) is presented in Fig. 3 and discussed in the “Discussion” section, where the first uranium oxide generation (U1) and the third one (U3) are presented. In the mineralised intercepts, the early diagenesis is represented by hematite (H1) crystallisation and quartz overgrowths (Q1) on detrital quartz (Q0) grains (Fig. 4a). This was followed by an episode of dissolution of detrital quartz Q0, noted as Q⁻ in Fig. 4b, in relation to peak diagenesis. Chlorite in rosette-shape (C1, Fig. 4c), rutile (R1, Fig. 4d), clays minerals (dickite D1 and illite I1) and apatite (Ap; mostly 10–20 µm in width, Fig.5a) form during this stage. Chlorite C1 is altered into clay minerals (I1, Fig. 4e). A second generation of quartz (Q2, Fig. 4f) occurs at the end of this stage. A second generation of uranium oxide (U2) crystallises in contact with apatite (Ap) in voids produced by quartz dissolution (Fig. 5a). The U2 postdates apatite as U2 fills small fractures within apatite grains and surrounds them. Veins containing quartz (Q3), uraninite (U4), pyrite (Py3) and chalcopyrite (ccp) crosscut the sandstone (Fig. 5c) and they are thought to be coeval with U4 emplaced in voids (Fig. 5b) based on dating results. A fifth (U5) and sixth (U6) generations of uranium oxide occur (Fig. 5d), filling voids. The mineral assemblage is similar for three last uranium generations, but the ages obtained account for the different generations.

301

## 4.2. Chemical composition of the uranium oxides

The chemical composition of the different uranium oxides are reported in Table SM2 for the major elements and in Table SM3 for minor and trace elements. Fig. 6 shows the chemical maps obtained by EPMA, on which the scales were determined using element concentrations measured during *in-situ* spot EPMA analyses.

305

### *Major elements*

The U contents of the different generations of uranium oxides range from 63.2 to 90.15 wt% UO₂. The thorium contents for all the uranium oxides is low with concentrations below the detection limits (0.1 wt% Th). The different uranium oxides are characterised by variable contents of Pb, either between samples or within an individual sample (Figs. 6 and 7a). For each sample, the maximum Pb content, corresponding to the least altered zones, and the minimum Pb content which is measured in the most altered ones, are summarized in Table SM2.

306

307

308

309

310
The samples WDD10-151 29.4 (U4) and WDD08-106 10.6 (U5) have the most homogeneous Pb contents and can be considered as the least altered. The Ti contents are below the detection limit (0.07 wt% TiO₂) except for samples JDD08-003 74.4B (U2) with concentrations up to 1.39 wt% TiO₂, WDD07-2 55.9A and WDD07-2 55.9B (U4) up to 2.6 wt% TiO₂ and WDD07-2 73.5B (U6) up to 1.25 wt% TiO₂. The P contents are lower than 2 wt% P₂O₅ in sample WDD07-2 73.5B (U6), but can reach 6.93 wt% P₂O₅ in zones where U concentrations are low in samples WDD07-2 55.9A and WDD07-2 55.9B (U4). The calcium and silica contents are also highly variable, ranging from 0.98 to 4.16 wt% CaO with the highest values found in sample JDD08-003 74.4B (U2), WDD07-2 55.9A and WDD07-2 55.9B (U4), and from < LOD (0.1 wt% SiO₂) to 13.11 wt% SiO₂ in zones with low U concentrations. The silica contents are higher than the calcium contents for all the samples except WDD08-106 10.6 (U5). The alteration for samples WDD07-2 55.9A, WDD07-2 55.9B (U4) and WDD07-2 73.5B (U6) is marked by a loss in lead compensated by a gain in Si, Ca, Fe and P (Fig. 7a), which is not observed for the other samples.

### Minor and trace elements

Among all the minor and trace elements measured by LA-ICP-MS, 14 have contents below the limit of detection (LOD) for some of the analyses. Table SM3 presents the results, with the minimal and maximal values for each sample, and the maximum LOD value calculated for each element considering all samples and all analyses. The diagrams comparing the contents in As, Nb, W, Ba and V (Fig. 7b, c, d) demonstrate that the different uraninite generations are chemically distinct. Arsenic contents are low in sample WDD07-2 73.5B (U6, 807–4870 ppm), range between 2490 and 5130 ppm for WDD07-2 55.9A (U4), and the highest values are obtained in sample JDD08-003 74.4B (U2, 5750–17590 ppm), WDD10-151 29.4 (U4, 14300–25800 ppm) and WDD08-106 10.6 (U5, 10120–13130 ppm, Fig. 7b and Table SM3). Niobium contents range between 5.6 and 137 ppm for JDD08-003 74.4B (U2), between 78.2 and 147.8 ppm for WDD07-2 55.9A (U4), between 6.3 and 16.9 ppm for WDD10-151 29.4 (U4), between 1.56 and 2.7 ppm for WDD08-106 10.6 (U5), and between 0.24 and 108.8 ppm for WDD07-2 73.5B (U6). Tungsten contents range between 67.3 and 156.6 ppm for WDD10-151 29.4 (U4 in veins) and between 551 and 8350 ppm for the other samples. Barium contents range between 227 and 2280 ppm for JDD08-003 74.4B (U2), between 101 and 283 ppm for WDD10-151 29.4 (U4), between 32.6 and 856 ppm for WDD07-2 55.9A (U4), between 69.4 and 77.6 ppm for WDD08-106 10.6m (U5), and between 584 and 7780 ppm for WDD07-2 73.5B (U6). Vanadium contents range between 469 and 1637 ppm for WDD07-2 55.9A (U4), and between 1418 and 16510 ppm for JDD08-003 74.4B (U2), WDD10-151 29.4 (U4 in veins), WDD08-106 10.6 (U5), and WDD07-2 73.5B (U6).
341 Rare Earth Elements

342 The REE chondrite-normalised patterns of the uranium oxides from the different samples are variable, but are all
343 characterised by a relatively high concentrations in REEs (between 537 and 40169 ppm ΣREE, Fig. 7e) and an
344 enrichment in LREE and MREE compared to HREE. The REE patterns are homogeneous for a given sample,
345 except for samples JDD08-003 74.4B, WDD07-2 55.9A and WDD07-2 55.9B (Fig. 8). The sample JDD08-003
346 74.4B (U2) is marked by a variation of MREE and HREE (1709 to 42014 ppm Gd and 154 to 6778 ppm Lu for
347 example), whereas sample WDD07-2 55.9A and WDD07-2 55.9B (U4) are marked by a variation of LREE (19
348 to 2493 ppm La for example). Increased LREE contents correlate to higher degrees of alteration in samples
349 WDD07-2 55.9A and WDD07-2 55.9B, whereas no link between alteration and MREE and HREE variations
350 could be established in sample JDD08-003 74.4B. Some of the samples display a negative Eu anomaly
351 (WDD07-2 55.9A and WDD07-2 55.9B, WDD10-151 29.4, WDD07-2 73.5B).

352 Correlations between elements

353 Some minor and trace elements are correlated in all the samples, whereas some others are only correlated in
354 some of the uraninite generations. The correlation matrices are presented in Tables SM4–SM9. In all samples the
355 following elements have a positive correlation: Co and Ni; the REEs; and Ba and Ga. In sample JDD08-003
356 74.4B, B is correlated to V, Zn, Ga and Ba; Mn is anti-correlated to the REEs; Bi is correlated to Co and Ni
357 (Table SM4). In sample WDD07-2 55.9A, V is correlated to Ba and the LREE; Mn to Sb; Cu to Ga, Sr and Ba
358 (Table SM5). In sample WDD07-2 55.9B, the REEs, Zr and Th are correlated; B is correlated to Mn and Sb, and
359 anti-correlated to Sr (Table SM6). In sample WDD10-151 29.4, B is correlated to the REES but anti-correlated to
360 V and Mn (Table SM7). In sample WDD08-106 10.6, As is correlated to the HREE and W, and anti-correlated
361 to V, Mn and Co; Mo is anti-correlated to the LREE (Table SM8). In sample WDD07-2 73.5B, Co is correlated
362 to Cu and Ga; Ti to Nb; B to Co; Ba to Ga and Sr (Table SM9).

363

364 4.3. Isotopic ages of apatite and uranium oxides

365 Apatite

366 Apatite predates uranium oxides, as observed in sample JDD08-003 74.4B (Fig. 9a). A significant number of
367 analyses present anomalous U and/or Pb contents correlated to anomalous Pb/Pb or U/Pb ratios (Table SM10
368 and Fig. 9b). These anomalous values are probably linked contribution from uraninite U2 during the analyses of
some apatite grains, due to the relatively small size of the apatite grains. When discarding such data, the eight remaining analyses are marked by variable concentrations of common Pb and degrees of discordance, with $^{204}$Pb/$^{206}$Pb and $^{207}$Pb/$^{206}$Pb ratios ranging respectively from 0.0015 to 0.0350 and 0.1281 to 0.5846 (Table SM10). The analyses plot in a discordant position on a Tera-Wasserburg diagram with the calculated lower intercept giving a poorly-defined age of 1685 ± 65 Ma (MSWD=1.4; 8 data points).

### Uranium oxides

The data for the different uranium oxides dated in this study are given in Table SM11. The majority of the U-Pb isotopic ratios are discordant in the Wetherill Concordia diagram (Fig. 10), indicating that all the uranium oxides suffered from post-crystallisation alteration. The least altered zones (i.e. richest in Pb) were selected for the isotopic dating based on EPMA analyses and mapping (Fig. 6). All the uranium oxides have a contribution of common lead, with $^{204}$Pb/$^{206}$Pb ratios varying between 0.000001 and 0.001926. Sample WDD10-151 29.4 (U4) has the highest common lead contribution, with $^{204}$Pb/$^{206}$Pb ratios between 0.000763 and 0.0011327, followed by samples JDD08-003 74.4B (U2, $^{204}$Pb/$^{206}$Pb: 0.000002–0.001926), WDD07-2 55.9A and WDD07-2 55.9B (U4, $^{204}$Pb/$^{206}$Pb: 0.000010–0.000171), WDD08-106 10.6 (U5, $^{204}$Pb/$^{206}$Pb: 0.000030–0.000057) and WDD07-2 73.5B (U6, $^{204}$Pb/$^{206}$Pb: 0.000003–0.000122). All the analyses were corrected for common lead using the common lead contribution calculated at their crystallisation age: 1030 Ma for JDD08-003 74.4B (U2), 730 Ma for WDD10-151 29.4 and WDD07-2 55.9A and WDD07-2 55.9B (U4), 550 Ma for WDD08-106 10.6 (U5), and 350 Ma for WDD07-2 73.5B (U6).

The 54 analyses of the uranium oxides from sample JDD08-003 74.4B (U2) are moderately to strongly discordant and give an upper intercept at 1035 ± 57 Ma and a lower intercept at the origin, indicating a recent alteration event (Fig. 10a). The data are scattered with a MSWD of 76, with a majority of analyses with U/Pb ratios close to the origin.

The uranium oxides from samples WDD07-2 55.9A and WDD07-2 55.9B (U4) give a global upper intercept age at 661 ± 39 Ma (49 analyses, lower intercept: 12 ± 11 Ma) but the data are highly scattered (MSWD = 170) and present highly variable discordancy behaviour, from almost concordant to strongly discordant (Fig. 10b). When using the less discordant data (25 analyses), the data are less but still scattered (MSWD = 84), with an upper intercept age at 738 ± 44 Ma and a lower intercept at 79 ± 45 Ma.
The 20 analyses of the uranium oxides from sample WDD10-151 29.4 (U4) are discordant but grouped (MSWD = 4.2) and give an upper intercept at 725 ± 23 Ma when anchoring the lower intercept at 0 ± 5 Ma (initial calculated lower intercept at -14 ± 6 Ma, Fig. 10c).

The 59 analyses of the uranium oxides from sample WDD08-106 10.6 (U5) are discordant but grouped (MSWD = 4.2) with an upper intercept at 554 ± 6 Ma when anchoring the lower intercept at 0 ± 5 Ma (initial calculated lower intercept at -16 ± 15 Ma, Fig. 10d).

The 85 analyses of the uranium oxides from sample WDD07-2 73.5B (U6) are highly scattered (MSWD = 139) and present a variable discordancy degree, giving an upper intercept of 350 ± 21 Ma when anchoring the lower intercept at 0 ± 5 Ma (Fig. 10e).

All the samples except two (WDD07 2 55.9A and WDD07-2 55.9B) are characterised by a lower intercept at almost 0 Ma, indicating that alteration of the uranium oxides is recent.

## 4.4. Chlorite composition and crystallisation temperatures

All the analysed chlorite grains have high iron contents (30.26–45.31 wt% FeO), low magnesium contents (2.01–11.99 wt% MgO), and plot along the Fe²⁺–Mg²⁺ chamosite–clinochlore substitution trend (Fig. 11a). Silica content ranges from 21.14 to 27.44 wt% and Al between 16.18 and 22.15 wt% Al₂O₃. Titanium, Mn, Ca, Na and K contents are consistently below detection limits except when chlorite is mixed with illite (Table SM12). Chlorite has a chamositic composition (Fig. 11b, Polito et al. 2005a; this study). To estimate temperature, the model developed by Bourdelle et al. (2013) was used because it was designed for diagenetic to low-grade metamorphic conditions and it avoids the problem of estimating the oxidation state of iron in chlorite. The results are presented in Table SM12, where the "Linear T°" column corresponds to the basic linear model for T°C < 350 °C, the "Quadratic T°" is based on a quadratic equation used above 350 °C, and the "Final T°" column gives the linear T° when it is below 350 °C and the quadratic one for temperatures above 350 °C. On 99 analyses, 13 range between 159 and 300 °C, most of the calculated temperatures are above 300 °C and 42 analyses could not be processed due to the sum of Si + Ti + Al + Fe²⁺ + Mn + Mg + Ca + Na + K - 4 (in apfu) higher than 6. Other thermometers were tested such as the ones developed by Cathelineau (1988) and by Jowett (1991) that also gave calculated temperatures higher than 300 °C.
422

# 5. Discussion

423

## 5.1. Timing of the U mineralisation in the Westmoreland area and its correlation to the geodynamic evolution of Australia

424

The paragenesis defined during this study is based on eight representative samples from the mineralised areas of the Redtree, Junnagunna and Huarabagoo U deposits and was combined with the results of Polito et al. (2005a), because the latter paragenetic sequence takes into account a larger diversity of samples from barren, altered and mineralised zones (Fig. 3). Some of their mineral assemblages were not observed in this study, mainly Fe-, Ti-oxides in the detrital minerals, dickite (D1) and pyrite (P1) linked to the ca. 1680 Ma diagenetic event. In the present study, the dickite was not identified and the clay mineral replacing chlorite C1 was illite. A first generation of micron-sized grains of uranium oxide U1 (referred to as “uraninite” in Polito et al. 2005a) is texturally associated with illite I2, hematite H2 and rutile R2. Polito et al. (2005a) described another generation of uranium oxide (noted U3 in the combined sequence) that crystallised in voids or in veins, possibly associated with pyrite (Py2) and galena. These two generations of uranium oxides were not observed in this study.

The U-Pb isotopic ages for the uranium oxides range from 1035 ± 57 Ma (sample JDD08-003 74.4B) to 350 ± 21 Ma (WDD07-2 73.5B), with 4 distinct stages at ca. 1035 Ma, 730 Ma, 550 Ma and 350 Ma (Fig. 10). These ages are younger than the oldest ages calculated for the Westmoreland area on very fine-grained intergranular uraninite U1 inclusions (6–8µm) in hematite at Redtree, with $^{207}$Pb/$^{206}$Pb ages of 1606 ± 80 Ma and 1655 ± 83 Ma from two analyses obtained by LA-HR-ICPMS (Polito et al. 2005a). Although the common lead contribution in U1 was not measured, this old age is younger than, but falls within error of the 1680 ± 21 Ma $^{40}$Ar/$^{39}$Ar age measured for diagenetic illite extracted from the Westmoreland Conglomerate (Polito et al. 2005a) and the 1685 ± 65 Ma age measured for apatite in association with uranium oxides at the Junnagunna deposit (sample JDD08-03 74.4B; Fig. 9). Therefore, it can be proposed that a diagenetic or hydrothermal fluid event was active in the area at ca. 1680 Ma, during which clay minerals, apatite and a first stage of uranium oxide crystallised. However, as the U1 ages are younger than the age of the illite and apatite, and considering the associated errors, U1 could also have formed shortly after this major fluid event, or at the end of it. From a geodynamic point of view, the ca. 1680 Ma period in the Westmoreland area can be correlated to the onset of magmatism in the Warumpi Province (Scrimgeour et al. 2005) and to a phase of rifting and intrusion of the 1678 Ma Carters Bore Rhyolite and 1670 Ma syn-extensional Sybella Batholith in the Mount Isa Province (Gibson et al. 2016). This stage was followed by at least five successive events of U mineralisation at ca. 1035 Ma, 870 Ma (Polito et al. 2005a), 730

435

436

437

438

439

440

441

442

443

444

445

446

447

448

449

450
Ma, 550 Ma and 350 Ma, the ages evaluated in this study correlating with the younger events of Polito et al. (2005a, Fig. 10f). These events could either have remobilised pre-existing mineralisation or a new stock of uranium could have been brought to the deposit site, and these two possibilities are discussed in section 5.2.

Hypotheses can relate these events to the geodynamic evolution of the McArthur Basin or Australia in a more global view. The timing of the ca. 1035 Ma event corresponds to the 1085–1040 Ma Giles Event, which included a failed east-trending intracontinental rift in the western Musgrave Province and the extensive 1075 Ma Warakurna Large Igneous Province (Wingate et al. 2004; Evins et al. 2010; Howard et al. 2011; Ahmad and Munson 2013). The ca. 870 Ma event corresponds to initiation of the Centralian Superbasin at about 840 Ma related to extension, rifting and crustal sagging that may have been an early expression of the breakup of the Neoproterozoic Rodinia Supercontinent (Lindsay 2002; Greene 2010). The 730 Ma phase corresponds to a period of tectonic quiescence, with little evidence for major tectonic events (Ahmad and Munson 2013). As all the other ages can be correlated with major tectonic events, it can be proposed that a same situation is valid for 730 Ma: a currently undefined geodynamic event favoured U mobility. The last two uraniferous events (ca. 550 and 350 Ma) take place when central Australia was affected by major intraplate tectonism, that dissected the Centralian Superbasin, exhumed rocks from deep crustal levels, and resulted in large amplitude east-trending gravity anomalies in central Australia (Ahmad and Munson 2013). The first of these events, the Petermann Orogeny, was focused in the Musgrave Province at 580–530 Ma and overlaps to the ca. 550 Ma event. Then, the Alice Springs Orogeny was a major intraplate orogenic event that variably affected large areas of central Australia from 450 to 300 Ma and resulted in the exhumation of the Arunta Region (Collins and Teyssier 1989; Haines et al. 2001) which could be related to the ca. 350 Ma mineralisation. The temporal correlation between the ages of the different uranium oxides in the Westmoreland area and successive geodynamic events having affected Australia demonstrates that these events favoured fluid circulations in the McArthur Basin and related basement, and uranium mobility during more than 1.3 Ga. However, none of the uraninite generations identified so far record the Isan Orogeny (1640–1490 Ma, with the peak around ca 1590 Ma, Murphy et al. 2011) that affected the nearby Mount Isa Basin.

## 5.2. Geochemistry of the successive U mineralisation: changing physico-chemical conditions in the Westmoreland area

The chemical composition of the ca. 1680 Ma uranium oxide U1 (“uraninite” in Polito et al. 2005a) is unknown, precluding comparison with later generations. The chemical compositions of the different generations of uranium oxides identified during this study are specific for each sample in terms of chondrite-normalised REE
patterns and minor to trace element concentrations (Figs. 7 and 8). Such results indicate at first order that the
physico-chemical conditions for the formation of these different uranium oxides were different (Mercadier et al.
2011b; Frimmel et al. 2014; Alexandre et al. 2015). Although the origin of this variability is not precisely
constrained (i.e., the chemical composition of the mineralising fluids, the source(s) of the fluids or of the metals,
or the pressure-temperature conditions; see Mercadier et al. 2011b), such variability suggests that the physico-
chemical conditions within the studied deposits have not been constant through time. The successive generations
of uranium oxides do not show any U-Pb isotopic signatures that indicate that they have been altered and/or
remobilised by later fluid events; all the lower intercepts of the Discordia lines are younger than 100 Ma (Fig.
10). When combined with the variability in the geochemical signatures of the different generations of uranium
oxides, these results do not allow conclusively define if U in the Westmoreland area was deposited at ca. 1680
Ma then locally remobilised during five fluid events under different conditions, or if each successive fluid event
brought a new batch of uranium and related metals to the deposits. The latter option is preferred, as
remobilisation of primary uranium in Canadian unconformity-related deposits leads to the formation of uranium
oxides in which some of the geochemical markers, like REE patterns, are preserved across the generations
(Mercadier et al. 2011a; Martz et al. 2019). The significant variations in the geochemical signatures of the
uranium oxides probably indicate new U inputs for each successive generation. In addition to unique chemical
signatures, each U-forming event in the Westmoreland area is characterised by a specific mineral assemblage
that reflects different formation conditions, or degrees of post-ore alteration (Polito et al. 2011). The U
mineralisation currently observed in the Westmoreland area is thus the result of a protracted history spanning at
least 1.3 Ga (1.68 to 0.35 Ga) of fluid circulations for which the physico-chemical conditions have varied, in
response to specific geodynamic events.

5.3. **Chemical composition of the clay minerals and related temperature of crystallisation**

Published data on C1 chlorite from the Westmoreland U deposits indicate that this unique generation of chlorite
is Fe-dominated but has variable Al contents and Mg/Fe ratios depending on their location (Westmoreland
Conglomerate and Seigal Volcanics), plotting between the chamosite, clinochlore and Mg-amesite compositions
(Fig. 11b; Polito et al. 2005a). The variation in chemical composition leads to variable range of calculated
temperatures of formation from 196 °C to 364 °C with the Cathelineau and Nieva (1985) and Cathelineau (1988)
methods (Polito et al. 2005a); and from 110 °C to 406 °C with the Bourdelle et al. (2013) method used for this
study. All the new analysed C1 chlorite, regardless of the location or deposit, lie in the chamosite field with
related temperature of formation mostly above 300 °C (Fig. 11, Table SM12). Polito et al. (2005a) proposed that the Fe-rich composition of the chlorite yielding the highest temperatures of crystallisation (364 °C) was linked to the proximity of the Seigal Volcanics that provided excess Fe during crystallisation of the chlorite. In addition, the Fe²⁺ in the Fe-rich chlorite adjacent to the mafic dykes and volcanic or mafic rocks themselves may have been the chemical reductant that reduced the hexavalent uranium to precipitate uraninite (Wall 2006; International Atomic Energy Agency 2018a). The distance between our samples and the Seigal Volcanics is unknown, but the same process could have resulted in the high Fe contents and the high temperature of crystallisation. Temperature estimates from fluid inclusions in the Redtree, Junnagunna and Huarabagoo deposits are lacking, but fluid inclusions were analysed from two related U deposits (Eva Mine and Jackson Pit) in the basement below the Westmoreland Conglomerate (Mernagh and Wygralak 2011). Based on their description and relation to other minerals, we assume that the quartz veins in which the fluid inclusions were studied are genetically comparable to the Q₂ drusy vein described by Polito et al. (2005a) which are co-genetic or post-date C1 chlorite. Temperatures of homogenisation for the different types of fluid inclusions are between 99 to 380 °C, with two populations: 99–270 °C and 190–380 °C (Mernagh and Wygralak 2011). Assuming that data obtained by Mernagh and Wygralak (2011) can be extrapolated to the Redtree, Junnagunna and Huarabagoo deposits, both chlorite and fluid inclusions indicate abnormally high thermal conditions (T>300 °C) during diagenesis at ca. 1680 Ma and reflect three possible origins. Firstly, the high temperatures could be due to a high regional geothermal gradient. Secondly, the fluids could be heated by the Seigal Volcanics whose age is not constrained but could be correlated to ca. 1780 Ma volcanic rocks (Mernagh and Wygralak 2011). Thirdly, fluids heated in the basement could have been injected into the basin. These high thermal conditions were probably short-lived because temperature estimates for illite I1 and illite I2, which post-date chlorite C1 and is co-genetic of U1 uranium stage at 1655 Ma, respectively, range from 150 °C to 250 °C (Polito et al. 2005a).

The temperature estimates and ages for the first U mineralising event are however loosely-constrained and the possibility that this U stage formed at ca. 1680 Ma in direct relation with the emplacement of the Seigal Volcanics has to be reconsidered, in particular due to spatial correlation between the U mineralisation and Seigal Volcanics (Fig. 2) and the possible older age of the volcanic rocks. Dyke emplacement driving formation of basin-hosted U mineralisation in Proterozoic basins such as McArthur Basin has been proposed for the Matoush and Camie River U deposits in the Otish Basin, Canada (Alexandre et al. 2015; Lesbros-Piat-Desvial et al. 2017), with temperatures of formation exceeding 300 °C. Such implications should be studied in more detail for the Westmoreland area.
No information regarding the type of alteration, nature of the fluids or temperatures can be currently provided for the five generations of uranium oxides postdating the ca. 1680 Ma U1 of Polito et al. (2005a), as these generations are not genetically associated with clay minerals, and fluid inclusions study in quartz Q3 associated with uraninite U4 are underway. A first-order assessment could be proposed based the REE contents measured in the uranium oxides (Mercadier et al. 2011b; Frimmel et al. 2014; Alexandre et al. 2015; Martz et al. 2019), but must be considered with caution. As ΣREE for U2 (sample JDD08-003 74.4B) formed at ca. 1030 Ma is above 1 wt% (Fig. 6e), a temperature estimate > 350 °C is proposed, whereas the following generations of uranium oxides were probably formed at a temperature < 350 °C.

## 5.4. Comparison between the uranium deposits in the Westmoreland area and unconformity-related U deposits

### 5.4.1. Comparison at the scale of the McArthur Basin: Westmoreland area vs ARUF

The Alligator Rivers Uranium Field (ARUF) of the McArthur Basin hosts numerous unconformity-related U deposits such as Ranger, Nabarlek, Jabiluka and Koongarra (Polito et al. 2004, 2005b, 2011; Jefferson et al. 2007; Cuney and Kyser 2015; Skirrow et al. 2016). As these deposits and those in the Westmoreland area share similar first-order settings and are all associated with the McArthur Basin, it is critical to establish if they could be linked to the same ore-forming events and/or ore-forming processes.

Deposits in the ARUF are located exclusively in the basement rocks, near the unconformity between an Archean to Paleoproterozoic basement complex and the McArthur Basin, are structurally-controlled, and have grades mostly above 0.1% U₃O₈ (Rawlings 1999; Jackson et al. 2000). Deposits in the Westmoreland area are structurally-controlled, spatially-related to volcanic dykes and sills in the basin and have lower grades (Hein 2002; Polito et al. 2005a; Skirrow et al. 2016; Laramide Resources Ltd 2019).

The ARUF and the Westmoreland area were affected by a major diagenetic/hydrothermal fluid circulation event at ca. 1680–1600 Ma (Fig. 12). This episode is recorded in multiple isotopic ages measured on illite, apatite and uranium oxides for the two regions (Maas 1989; Polito et al. 2004, 2005a, b; Clauer et al. 2015; Skirrow et al. 2016). The fluids of the diagenetic/hydrothermal event in the ARUF and Westmoreland area at 1680–1600 Ma were basin-derived brines and low-salinity fluids (Ypma and Fuzikawa 1980; Wilde et al. 1989; Derome et al. 2003, 2007; Polito et al. 2004, 2005b, 2006; Mernagh and Wygralak 2011). The isotopic and fluid inclusions data at Westmoreland show the involvement of NaCl-rich and CaCl₂-rich brines derived from evaporated seawater, comparable to the two brines related to formation of the unconformity-related deposits in the ARUF
570 (Derome et al. 2007). The mixing between low-salinity fluids and brines is a key process for ore deposition in
571 both areas (e.g. Derome et al., 2003, 2007; Mernagh and Wygralak, 2011). Although contemporaneous and
572 related to the same types of fluids, the fluids at ca. 1680 Ma in the ARUF and Westmoreland area are physico-
573 chemically different based on the differences in the mineralogy and chemistry of the minerals which crystallised
574 at that time in both areas. The most recent studies on unconformity-related U deposits from the ARUF, in
575 particular for the Ranger U deposit, show significant differences with the deposits in the Westmoreland area in
576 terms of the composition of chlorite (Skirrow et al. 2016; Gigon et al. 2019) and the presence and composition of
577 hydrothermal tourmaline (Gigon et al. 2019).

578 The unconformity-related U deposits in the ARUF at ca. 1680–1600 Ma formed through Mg- and B-
579 metasomatism (Fisher et al. 2013; Skirrow et al. 2016; Gigon et al. 2019), but such metasomatism is lacking in
580 the Westmoreland area. The absence of hydrothermal Mg-tourmaline with the uranium oxides in the
581 Westmoreland area is a first striking difference, as this mineral is a pathfinder for unconformity-related U
582 deposits in the ARUF (Skirrow et al. 2016; Gigon et al. 2019). The second difference is the chlorite composition:
583 the chlorite in the Westmoreland area are dominantly chamosite (Fe-rich), compared to the clinoclore to sudoite
584 composition (Mg-rich) of the diagenetic/hydrothermal chlorite in the unconformity-related deposits in the ARUF
585 (Fig. 11b, Polito et al. 2004, 2005b; Beaufort et al. 2005; Gigon et al. 2019). As a consequence, the
586 diagenetic/hydrothermal chlorite composition for the ARUF yields temperatures between 101 °C and 163 °C,
587 with mean temperature of 128 °C and median temperature of 124 °C (Gigon et al. 2019), considerably lower
588 than the > 300 °C calculated in this study.

589 The present study identifies for the first time multiple U generations between 1035 and 350 Ma in three deposits
590 of the Westmoreland area. The succession of fluid events also characterises the unconformity-related U deposits
591 from the ARUF, with multiple ages ranging from 1680 Ma to Present measured on both uranium oxides and illite
592 (Hills and Richards 1976; Ludwig et al. 1987; Maas 1989; Polito et al. 2004, 2005b; Ahmad and Munson 2013;
593 Clauer et al. 2015; Skirrow et al. 2016). These post-1680 Ma ages in the ARUF mainly fall between 1600 and
594 1300 Ma for the clay minerals and some uranium oxides, but are dominantly younger than 900 Ma for the
595 uranium oxides. Based on the available datasets, the post-1680 Ma ages measured for the ARUF do not totally
596 correlate with the five post-1680 Ma ages measured at the Westmoreland area, which could reflect that the fluid
597 events in the two areas are partially disconnected in time from 1680 Ma to Present. The nature of the fluid(s) and
598 physico-chemical conditions of these events in the ARUF are not well constrained.
It is currently not possible to link the physico-chemical conditions for the U1 generation in the Westmoreland area with the primary uraninite in the ARUF as no analyses were done on U1. Both primary mineralisation could have been formed in a similar system, but under higher temperature conditions in the Westmoreland area. In this area, the ultimate origin of the different generations and the physico-chemical conditions prevailing between 1035 and 350 Ma also remain to be constrained. The chondrite-normalised REE patterns of uranium oxides, which are diagnostic features of deposit types, differ significantly between the ARUF and the Westmoreland area. The different generations of uranium oxides in the Westmoreland area do not exhibit the “bell-shape” (i.e. high concentrations of Tb and Dy) typical of unconformity-related U deposits from the ARUF such as Nabarlek and Koongarra (Mercadier et al. 2011b). Even if some post-crystallisation modification of the REE patterns may have contributed to the LREE enrichment (as described in the Athabasca Basin by Mercadier et al. 2011a and Martz et al. 2019), such phenomenon is not sufficient to explain the major differences of the REE patterns between the two areas. The uranium oxides from the Westmoreland area have significantly higher contents in common lead (Table SM11) when compared to typical uranium oxides from unconformity-related uranium oxides (Skirrow et al. 2016; Martz et al. 2019), which is also an indicator of differences in their conditions of formation. It can be concluded that the formation conditions of the different U generations in the Westmoreland area did not correspond to those known for unconformity-related U deposits in the ARUF.

### 5.4.2. Comparison between U deposits in the Westmoreland area and U deposits in Canadian Proterozoic basins

Several Paleo- to Mesoproterozoic basins in Canada host U mineralisation: the Athabasca Basin (Saskatchewan), the Thelon Basin (Nunavut) and the Otish Basin (Québec).

The U mineralisation in the Athabasca Basin is of unconformity-related type, like the one in the ARUF (Jefferson et al. 2007; Cuney and Kyser 2015). The ca. 1680-1600 Ma diagenetic/hydrothermal episode defined in the Westmoreland area is documented in the Thelon Basin, where apatite is dated at 1667 ± 5.1 Ma (Davis et al. 2011). In the Athabasca Basin, this fluid event has been recorded by TIMS analyses of hydrothermal zircon tips (uranium-rich, thorium-depleted) in a leucogranite sampled 200 meters east of the McArthur River high-grade U deposit with ages from 1680 to 1550 Ma. This was interpreted as the approximate timing of the thermotectonic event(s) and associated fluids that produced the new U-rich zircon (Annesley et al. 2007). In addition with the 1650-1700 Ma diagenetic apatite age (Cumming et al. 1987), such results confirm a global fluid event at 1680–1600 Ma affecting the Athabasca, Thelon and McArthur basins which could have belong to the same super basin at that time (Furlanetto et al. 2016). No U mineralisation has however been dated at 1680
Ma for the Athabasca and Thelon basins, with maximum ages at ca. 1590 Ma (Alexandre et al. 2009) and ca. 1500 Ma (Sharpe et al. 2015) respectively. Recent data on U mineralisation in the Thelon Basin indicate that these deposits had a polyphased history over hundreds of Ma in which unconformity-related U mineralisation formed from ca. 1500 Ma, followed by low-T and meteoric-related redox fronts (Sharpe et al. 2015; Chi et al. 2017; Shabaga et al. 2017; Grare et al. 2020).

In the Otish Basin, host of the Matoush and Camie River U deposits, no significant hydrothermal activity is known at ca. 1680 Ma. This basin formed earlier than the Athabasca and Thelon basins with a minimum deposition age of 2172 to 2162 Ma (Hamilton and Buchan 2016; Milidragovic et al. 2016). The oldest reported age of fluid circulation and mineralisation is ca. 1725 Ma for the Camie River U deposit (Beyer et al. 2012; Lesbros-Piat-Desvial et al. 2017). This ca. 1725 Ma U mineralisation was initially interpreted to be unconformity-related type (Beyer et al. 2012), but recent research contradicts this interpretation because of a Paleoproterozoic, higher-temperature hydrothermal event, whose origin remains speculative (Lesbros-Piat-Desvial et al. 2017). The timing of formation of the Matoush deposit is loosely constrained (1695 ± 110 Ma; Alexandre et al. 2015), and its main characteristics (spatial association with a dyke, alteration mineralogy, association between uraninite and eskolaite) do not support classification as an unconformity-related type deposit, although the U-bearing fluids were oxidising basinal fluids (Alexandre et al. 2015), like for the Westmoreland area.

The chemical composition of uranium oxides analysed in this study records major differences for minor and trace elements (including REEs) compared to typical unconformity-related U deposits from the Athabasca (Mercadier et al. 2011b; Alexandre et al. 2016; Martz et al. 2019) or Thelon basins (Grare et al. 2020). This is particularly visible when comparing the chondrite-normalised REE patterns, as the different generations of uranium oxides from the Westmoreland area do not have the typical bell-shape pattern of the unconformity-related U deposits (Fig. 8). Differences are also visible for numerous minor and trace elements like As, W, Ba and V (Fig. 7c and d) compared with McArthur River and Cigar Lake U deposits of the Athabasca Basin (Alexandre et al. 2015; Martz et al. 2019). The majority of the minor and trace elements have significantly higher concentrations in the uranium oxides from the Westmoreland area, which could indicate a higher temperature of formation and/or different sources of metals. The complete geochemistry of the uranium oxides from the Otish Basin is lacking, but the REE patterns of the ca. 1725 Ma uranium oxides from the Camie River are rather similar to those measured for different generations of uranium oxide from the Westmoreland area (Fig. 8). Such results allow consideration that the physico-chemical conditions for the formation of the U
mineralisation in the Westmoreland area were probably more similar to those of Matoush or Camie River U deposits of the Otish Basin than those of the Athabasca and Thelon basins. Temperatures estimates from the Matoush (~300 ±50 °C, based on illite and chlorite crystal chemistry, Alexandre et al. 2015) and Camie River deposits (275 to 300 °C, Beyer et al. 2012 and 320 °C, Lesbros-Piat-Desvial et al. 2017) are comparable to the Westmoreland area, but significantly higher than values reported from the Athabasca and Thelon basins (e.g. Kotzer and Kyser 1995; Richard et al. 2016; Chi et al. 2017; Grare et al. 2020).

Other similarities with the U deposits from the Otish Basin are the lack of Mg-rich minerals (sudoite, Mg-foitite) or alumino-phosphate-sulfate minerals, which are markers of unconformity-related U deposits. The Matoush and Camie River deposits are spatially related to dykes (Beyer et al., 2012; Alexandre et al. 2015; Lesbros-Piat-Desvial et al. 2017), which played a role as the reductant of uranium, because it was a source of Cr and Fe at Matoush (Alexandre et al. 2015). An equivalent process was proposed for the Westmoreland area with the implication of iron from the dykes as reductant (Polito et al. 2005a). The mafic activity in the Otish Basin and in the southern McArthur Basin (Westmoreland area) could thus have favoured the formation of atypical U deposits, although linked to the same types of fluids as unconformity-related U deposits.

## 6. Conclusions

The new data provided on U mineralised intercepts from three deposits of the Westmoreland area (Queensland and Northern Territory, Australia) indicate that this region of the McArthur Basin experienced several episodes of fluid circulation between ca. 1680 Ma to Present, which document a protracted mobility of uranium over more than 1.3 Ga. Based on published and new data, six different events of U mobility were active at ca. 1680 Ma, 1035 Ma, 850 Ma, 730 Ma, 550 Ma and 350 Ma. The ca. 1680 Ma mineralisation formed due to the circulation and mixing of NaCl-rich, CaCl₂-rich basin-derived brines and low-salinity fluids. The timing and type of fluids involved were comparable to those of the unconformity-related U deposits in the ARUF to the North of the McArthur Basin. The physico-chemical conditions for the formation of these contemporaneous uranium mineralisations are different between the two regions, with higher chlorite temperature (> 300 °C) for the Westmoreland area and absence of Mg- and B-metasomatism typical of unconformity-related U deposits. Although the precise chemical composition of the uranium oxides formed at ca. 1680 Ma has not been measured, all the other U generations found in the Westmoreland area are chemically different from the uranium oxides found in unconformity-related U deposits. The Westmoreland area samples contain higher contents of numerous
minor and trace elements, like REEs, which supports a higher temperature of crystallisation. If the timing and the
types of fluids between the two parts of the basin were similar at ca. 1680 Ma, fluids in the Westmoreland area
were at higher temperatures, driven by a high regional geothermal gradient controlled by volcanism (Seigal
Volcanics), or injection of fluids heated in the basement into the basin. This elevated thermal regime may have
triggered fluid circulation in the Westmoreland area, with the dykes providing the structural and chemical traps
that focused the uranium-bearing brines and favoured U deposition at ca. 1680 Ma. Such conditions appear
similar to those known for the formation of U deposits in the Proterozoic Otish Basin (Québec, Canada) at ca.
1725 Ma. The new data proposed in this study indicate that the geochemical signature of the successive uranium
oxides from the Westmorland area are poorly compatible with uranium oxides formed in unconformity-related U
deposits, as confirmed by the difference in mineral association. It is therefore considered as unlikely to find
unconformity-related U deposits in this area. The significant changes in the chemical compositions of the
different generations of uranium oxides are in favour of changing conditions between 1680 and 350 Ma.
Although the precise physico-chemical are loosely defined, the signatures observed in the Westmorland area for
the first U stages are closer to those proposed for U deposits from the Otish Basin than those for the Athabasca
or Thelon basins.

702

703 Acknowledgements

The Northern Territory Geological Survey provided logistical support for the sampling mission. Laramide
Resources Ltd. provided authorisation, logistical support and guidance for sampling on the Westmoreland
exploration camp (especially Adrian Buer and Evan Hughes). Alexandre Flammang and John Moine
(GeoRessources Lab) as well as Thin Section Lab company staff are acknowledged for careful thin section
preparation. Chantal Peiffert, Andreï Lecomte and Olivier Rouer (GeoRessources Lab) are acknowledged for
support in LA-ICP-MS, SEM and EPMA analyses respectively. This research was funded by (i) a French
Ministry of Higher Education and Research PhD salary Grant to J.G., (ii) a Région Lorraine-FEDER grant to
J.M.: « Rôle des phases fluides dans la distribution spatiale des ressources métalliques dans les bassins
sédimentaires paléoproterozoiques australiens », (iii) OSU OTELO grants to J.M.: « Conditions de transport des
métaux dans un mégabassin protérozoïque » and A.R.: « Transferts de fluides et métaux dans le bassin de
McArthur (Australie) » and (iv) a CNRS-INSU-CESSUR grant to J.M. « Transferts des fluides et métaux dans
les méga-bassins paléoprotérozoïques ». The French National Agency through the national program « Investissements d'avenir » of the Labex Ressources 21 (reference ANR-10-LABX-21-RESSOURCES21) partly funded the last developments of the SIMS lab. The comments of David Quirt and an anonymous reviewer greatly helped improving the manuscript. Mostafa Fayek is warmly thanked for careful editorial handling.

## 721 References

722 Ahmad M (1987) Uranium-gold occurrences of the Murphy Inlier and surrounding region. Northern Territory Geological Survey. Technical Report 87/2

724 Ahmad M, Dunster JN, Munson TJ (2013) Chapter 15: McArthur Basin: in Ahmad M and Munson TJ (compilers). 'Geology and mineral resources of the Northern Territory'. Special Publication 5

726 Ahmad M, Munson TJ (compilers) (2013) Geology and mineral resources of the Northern Territory. Northern Territory Geological Survey/ISBN 978-0-7245-7257-1

728 Ahmad M, Wygralak A (1989) Calvert Hills, Northern Territory. 1: 250 000 metallogenic map series explanatory notes, SE 53-08. Northern Territory Geological Survey

730 Ahmad M, Wygralak AS (1990) Murphy Inlier and environs—regional geology and mineralisation. Geology of the Mineral Deposits of Australia and Papua New Guinea (editor Hughes FE) The Australasian Institute of Mining & Metallurgy, Melbourne 819–826

733 Alexandre P, Kyser K, Layton-Matthews D, et al (2016) Chemical compositions of natural uraninite. Canadian Mineralogist 53:1–30. https://doi.org/10.3749/canmin.1500017

735 Alexandre P, Kyser K, Layton-Matthews D, et al (2015) Formation of the enigmatic Matoush uranium deposit in the Paleoprotozoic Otish Basin, Quebec, Canada. Mineralium Deposita 50:825–845. https://doi.org/10.1007/s00126-014-0569-5

738 Alexandre P, Kyser K, Thomas D, et al (2009) Geochronology of unconformity-related uranium deposits in the Athabasca Basin, Saskatchewan, Canada and their integration in the evolution of the basin. Miner Deposita 44:41. https://doi.org/10.1007/s00126-007-0153-3

741 Anders E, Grevesse N (1989) Abundances of the elements: Meteoritic and solar. Geochimica et Cosmochimica acta 53:197–214. https://doi.org/10.1016/0016-7037(89)90286-X

743 Annesley IR, Madore C, McCready AJ, et al (2007) U-Pb ages of zircon, monazite, and zircon tips in Hudsonian granites at the McArthur River uranium deposit: a record of U mineralizing fluids. In: Proceedings of the Ninth Biennial SGA Meeting. Dublin, pp 1125–1128

746 Beaufort D, Patrier P, Laverret E, et al (2005) Clay alteration associated with Proterozoic unconformity-type uranium deposits in the East Alligator Rivers uranium field, Northern Territory, Australia. Economic Geology 100:515–536

749 Beyer SR, Kyser K, Hiatt EE, et al (2012) Basin Evolution and Unconformity-Related Uranium Mineralization: The Camie River U Prospect, Paleoproterozoic Otish Basin, Quebec. Economic Geology 107:401–425. https://doi.org/10.2113/econgeo.107.3.401
752 Billault V, Beaufort D, Patrier P, Petit S (2002) Crystal chemistry of Fe-sudoites from uranium deposits in the Athabasca Basin (Saskatchewan, Canada). Clays and Clay Minerals 50:70-81.
753
754 https://doi.org/10.1346/000986002761002847

755 Bourdelle F, Parra T, Chopin C, Beyssac O (2013) A new chlorite geothermometer for diagenetic to low-grade metamorphic conditions. Contributions to Mineralogy and Petrology 165:723-735.
756
757 https://doi.org/10.1007/s00410-012-0832-7

758 Cathelineau M (1988) Cation site occupancy in chlorites and illites as function of temperature. Clay minerals 23:471-85
759

760 Cathelineau M, Boiron MC, Holliger P, Poty B (1990) Metallogenesis of the French part of the Variscan orogen. Part II: Time-space relationships between U, Au and Sn-W ore deposition and geodynamic events—mineralogical and U-Pb data. Tectonophysics 177:59-79. https://doi.org/10.1016/0040-1951(90)90274-C
761
762
763

764 Cathelineau M, Nieva D (1985) A chlorite solid solution geothermometer the Los Azufres (Mexico) geothermal system. Contributions to Mineralogy and Petrology 91:235-244
765

766 Chew D, Drost K, Petrus JA (2019) Ultrafast, > 50 Hz LA - ICP - MS Spot Analysis Applied to U-Pb Dating of Zircon and other U- Bearing Minerals. Geostand Geoanal Res 43:39-60.
767
768 https://doi.org/10.1111/ggr.12257

769 Chi G, Haid T, Quirt D, et al (2017) Petrography, fluid inclusion analysis, and geochronology of the End uranium deposit, Kiggavik, Nunavut, Canada. Mineralium Deposita 52:211-232.
770
771 https://doi.org/10.1007/s00126-016-0657-9

772 Clauer N, Mercadier J, Patrier P, et al (2015) Relating unconformity-type uranium mineralization of the Alligator Rivers Uranium Field (Northern Territory, Australia) to the regional Proterozoic tectono-thermal activity: An illite K-Ar dating approach. Precambrian Research 269:107-121.
773
774
775 https://doi.org/10.1016/j.precamres.2015.08.007

776 Collins WJ, Teyssier C (1989) Crustal scale ductile fault systems in the Arunta Inlier, central Australia. Tectonophysics 158:49-66. https://doi.org/10.1016/0040-1951(89)90314-4
777

778 Cumming GL, Krstic D, Wilson JA (1987) Age of the Athabasca group, northern Alberta. In: Geological Association of Canada - Mineralogical Association of Canada Joint Annual Meeting, Program with Abstracts. p 35
779
780

781 Cuney M, Kyser K (2015) Geology and geochemistry of uranium and thorium deposits. Mineralogical Association of Canada. ISBN 0-921294-57-3
782

783 Davis WJ, Gall Q, Jefferson CW, Rainbird RH (2011) Fluorapatite in the Paleoproterozoic Thelon Basin: Structural-stratigraphic context, in situ ion microprobe U-Pb ages, and fluid-flow history. Bulletin 123:1056-1073
784
785

786 Derome D, Cathelineau M, Fabre C, et al (2007) Paleo-fluid composition determined from individual fluid inclusions by Raman and LIBS: Application to mid-proterozoic evaporitic Na-Ca brines (Alligator Rivers Uranium Field, Northern Territory Australia). Chem Geol 237:240-254.
787
788
789 https://doi.org/10.1016/j.chemgeo.2006.10.015

790 Derome D, Cuney M, Cathelineau M, et al (2003) A detailed fluid inclusion study in silicified breccias from the Kombolgie sandstones (Northern Territory, Australia): inferences for the genesis of middle-Proterozoic unconformity-type uranium deposits. Journal of Geochemical Exploration 80:259-275.
791
792
793 https://doi.org/10.1016/S0375-6742(03)00194-8

794 Durak B, Pagel M, Poty B (1983) Températures et salinités des fluides au cours des silicifications diagénétiques d'une formation gréseuse surmontant un gisement d'uranium du socle: l'exemple des grès Kombolgie (Australie). Paris
795
796
797 Evins PM, Smithies RH, Howard HM, et al (2010) Redefining the Giles Event within the setting of the 1120–
798 1020 Ma Ngaanyatjarra Rift, west Musgrave Province, Central Australia. Geological Survey of Western
799 Australia, Record 6:36

800 Fayek M, Harrison TM, Ewing RC, et al (2002) O and Pb isotopic analyses of uranium minerals by ion
801 microprobe and U–Pb ages from the Cigar Lake deposit. Chemical Geology 185:205–225.
802 https://doi.org/10.1016/S0009-2541(01)00401-6

803 Fayek M, Harrison TM, Grove M, Coath CD (2000) A Rapid In *Situ* Method for Determining the Ages of
804 Uranium Oxide Minerals: Evolution of the Cigar Lake Deposit, Athabasca Basin. International Geology
805 Review 42:163–171. https://doi.org/10.1080/00206810009465075

806 Fisher LA, Cleverley JS, Pownceby M, MacRae C (2013) 3D representation of geochemical data, the
807 corresponding alteration and associated REE mobility at the Ranger uranium deposit, Northern
808 Territory, Australia. Mineralium Deposita 48:947–966. https://doi.org/10.1007/s00126-013-0463-6

809 Frimmel HE, Schedel S, Brätz H (2014) Uraninite chemistry as forensic tool for provenance analysis. Applied
810 Geochemistry 48:104–121. https://doi.org/10.1016/j.apgeochem.2014.07.013

811 Fuchs HD, Schindlmayer WE (1981) The Westmoreland uranium deposit, Queensland, Australia. In: Uranium
812 exploration case histories. International Atomic Energy Agency

813 Furlanetto F, Thorkelson DJ, Rainbird RH, et al (2016) The Paleoproterozoic Wernecke Supergroup of Yukon,
814 Canada: Relationships to orogeny in northwestern Laurentia and basins in North America, East
815 Australia, and China. Gondwana Research 39:14–40. https://doi.org/10.1016/j.gr.2016.06.007

816 Gibson GM, Meixner AJ, Withnall IW, et al (2016) Basin architecture and evolution in the Mount Isa mineral
817 province, northern Australia: Constraints from deep seismic reflection profiling and implications for ore
818 genesis. Ore Geology Reviews 76:414–441. https://doi.org/10.1016/j.oregeorev.2015.07.013

819 Gigon J, Deloule E, Huston DL, et al (2020) Tracing metal sources for the giant McArthur River Zn-Pb ore
820 deposit using Pb-isotopes. Geology: https://doi.org/10.1130/G47001.1

821 Gigon J, Skirrow RG, Harlaux M, et al (2019) Insights into B-Mg-Metasomatism at the Ranger U Deposit (NT,
822 Australia) and Comparison with Canadian Unconformity-Related U Deposits. Minerals 9:432.
823 https://doi.org/10.3390/min9070432

824 Grare A, Benedicto A, Mercadier J, et al (2020) Structural controls and metallogenic model of polyphase
825 uranium mineralization in the Kiggavik area (Nunavut, Canada). Miner Deposita.
826 https://doi.org/10.1007/s00126-020-00957-x

827 Greene DC (2010) Neoproterozoic rifting in the southern Georgina Basin, central Australia: Implications for
828 reconstructing Australia in Rodinia. Tectonics 29:20 p. https://doi.org/10.1029/2009TC002543

829 Haines PW, Hand M, Sandiford M (2001) Palaeozoic synorogenic sedimentation in central and northern
830 Australia: A review of distribution and timing with implications for the evolution of intracontinental
831 orogens. Australian Journal of Earth Sciences 48:911–928. https://doi.org/10.1046/j.1440-8326.2001.00909.x

833 Hamilton MA, Buchan KL (2016) A 2169 Ma U–Pb baddeleyite age for the Otish Gabbro, Quebec: implications
834 for correlation of Proterozoic magmatic events and sedimentary sequences in the eastern Superior
835 Province. Canadian Journal of Earth Sciences 53:119–128

836 Hein KAA (2002) Geology of the Ranger Uranium Mine, Northern Territory, Australia: structural constraints on
837 the timing of uranium emplacement. Ore Geology Reviews 20:83–108. https://doi.org/10.1016/S0169-8338(02)00054-9

839 Hills JH, Richards JR (1976) Pitchblende and galena ages in the Alligator Rivers region, Northern Territory,
840 Australia. Mineralium Deposita 11:133–154. https://doi.org/10.1007/BF00204477
841 Holliger P (1988) *Ages U-Pb définis in situ sur oxydes d'uranium à l'analyseur ionique: méthodologie et conséquences géochimiques*. Comptes rendus de l'Académie des sciences Série 2, Mécanique, Physique, Chimie, Sciences de l'univers, Sciences de la Terre 307:367-373

844 Hollis JA, Beyer EE, Whelan JA, et al (2010) Summary of results. NTGS laser U-Pb and Hf geochronology project: Pine Creek Orogen, Murphy Inlier, McArthur Basin and Arunta Region, July 2007-June 2008. Northern Territory Geological Survey, Record 1:

847 Howard HM, Smithies RH, Evins PM, et al (2011) Explanatory notes for the west Musgrave Province. Geological Survey of Western Australia, Record 4:349

849 International Atomic Energy Agency (2018a) Geological classification of uranium deposits and description of selected examples. ISBN 978-92-0-101618-8

851 International Atomic Energy Agency (2018b) Unconformity-related uranium deposits. ISBN 978-92-0-108518-4

852 Jackson MJ, Scott DL, Rawlings DJ (2000) Stratigraphic framework for the Leichhardt and Calvert Superbasins: review and correlations of the pre-1700 Ma successions between Mt Isa and McArthur River. Aust J Earth Sci 47:381-403. https://doi.org/10.1046/j.1440-0952.2000.00789.x

855 Jefferson CW, Thomas DJ, Gandhi SS, et al (2007) Unconformity-associated uranium deposits of the Athabasca Basin, Saskatchewan and Alberta. Bulletin / Geological Survey of Canada 588:23

857 Jochum KP, Weis U, Stoll B, et al (2011) Determination of reference values for NIST SRM 610-617 glasses following ISO guidelines. Geostandards and Geoanalytical Research 35:397-429. https://doi.org/10.1111/j.1751-908X.2011.00120.x

860 Jowett EC (1991) Fitting iron and magnesium into the hydrothermal chlorite geothermometer. In: GAC/MAC/SEG Joint Annual Meeting, Toronto, May 27-29, 1991, Program with Abstracts 16

862 Knutson J, Ferguson J, Roberts WMB, et al (1979) Petrogenesis of the copper-bearing breccia pipes, Redbank, Northern Territory, Australia. Economic Geology 74:814-826. https://doi.org/10.2113/gsecongeo.74.4.814

865 Kotzer TG, Kyser TK (1995) Petrogenesis of the Proterozoic Athabasca Basin, northern Saskatchewan, Canada, and its relation to diagenesis, hydrothermal uranium mineralization and paleohydrogeology. Chemical Geology 120:45-89. https://doi.org/10.1016/0009-2541(94)00114-N

868 Lach P, Mercadier J, Dubessy J, et al (2013) *In Situ* Quantitative Measurement of Rare Earth Elements in Uranium Oxides by Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry. Geostand Geoanal Res 37:277-296. https://doi.org/10.1111/j.1751-908X.2012.00161.x

871 Lally JH, Bajwah Z (2006) Uranium deposits of the Northern Territory. Northern Territory Geological Survey, Report 20

873 Laramide Resources Ltd (2019) Westmoreland Uranium Project website. https://laramide.com/projects/westmoreland-project/. Accessed 26 Jun 2019

875 Laramide Resources Ltd (2016) Westmoreland Uranium Project - National instrument 43-101; Technical report - scoping study

877 Lesbros-Piat-Desvial M, Beaudoin G, Mercadier J, Creaser R (2017) Age and origin of uranium mineralization in the Camie River deposit (Otish Basin, Québec, Canada). Ore Geology Reviews 91:196-215. https://doi.org/10.1016/j.oregeorev.2017.10.006

880 Li Q-L, Li X-H, Wu F-Y, et al (2012) In-situ SIMS U-Pb dating of phanerozoic apatite with low U and high common Pb. Gondwana Research 21:745-756
882 Lindsay JF (2002) Supersequences, superbasins, supercontinents evidence from the Neoproterozoic-Early Palaeozoic basins of central Australia. Basin Research 14:207-223. https://doi.org/10.1046/j.1365-2117.2002.00170.x

885 Longerich HP, Jackson SE, Günther D (1996) Inter-laboratory note. Laser ablation inductively coupled plasma mass spectrometric transient signal data acquisition and analyte concentration calculation. Journal of analytical atomic spectrometry 11:899-904. https://doi.org/10.1039/AA9961100899

888 Ludwig K (2007) Isoplot/Ex version 3.41b, a geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center Spec. Publ. No 4:

890 Ludwig KR, Grauch RI, Nutt CJ, et al (1987) Age of uranium mineralization at the Jabiluka and Ranger deposits, Northern Territory, Australia; new U-Pb isotope evidence. Economic Geology 82:857-874. https://doi.org/10.2113/gsecongeo.82.4.857

893 Luo J-C, Hu R-Z, Fayek M, et al (2015) In-situ SIMS uraninite U-Pb dating and genesis of the Xianshi granite-hosted uranium deposit, South China. Ore Geology Reviews 65:968-978

895 Maas R (1989) Nd-Sr Isotope Constraints on the Age and Origin of Unconformity-type Uranium Deposits in the Alligator Rivers Uranium Field, Northern-Territory, Australia. Economic Geology 84:64-90. https://doi.org/10.2113/gsecongeo.84.1.64

898 Martz P (2017) Caractéristiques, chronologie et rôles des circulations fluides dans le bassin d'Athabasca et son socle: implications dans la formation et l'évolution du gisement d'uranium de Cigar Lake. PhD thesis; Université de Lorraine

901 Martz P, Mercadier J, Perret J, et al (2019) Post-crystallization alteration of natural uraninites: Implications for dating, tracing, and nuclear forensics. Geochimica et Cosmochimica Acta 249:138-159. https://doi.org/10.1016/j.gca.2019.01.025

904 McKay AD, Miezitis Y (2001) Australia's uranium resources, geology and development of deposits. AGSO-Geoscience Australia/ ISBN 0-642-46716-1

906 Mercadier J, Cuney M, Cathelineau M, Lacorde M (2011a) U redox fronts and kaolinisation in basement-hosted unconformity-related U ores of the Athabasca Basin (Canada): late U remobilisation by meteoric fluids. Mineralium Deposita 46:105-135. https://doi.org/10.1007/s00126-010-0314-7

909 Mercadier J, Cuney M, Lach P, et al (2011b) Origin of uranium deposits revealed by their rare earth element signature. Terra Nova 23:264-269. https://doi.org/10.1111/j.1365-3121.2011.01008.x

911 Mernagh TP, Wygralak AS (2011) A fluid inclusion study of uranium and copper mineral systems in the Murphy Inlier, Northern Australia. Russ Geol Geophys 52:1421-1435. https://doi.org/10.1016/j.rgg.2011.10.011

914 Milidragovic D, Beaudoin G, Hamilton MA, King JJ (2016) The Paleoproterozoic Otish Gabbro suite and coeval dyke swarms of the Superior Province: Probing the ca. 2.17 Ga mantle. Precambrian Research 278:126-144

917 Murphy FC, Hutton LJ, Walshe JL, et al (2011) Mineral system analysis of the Mt Isa-McArthur River region, Northern Australia. Australian Journal of Earth Sciences 58:849-873. https://doi.org/10.1080/08120099.2011.606333

920 Page RW, Jackson MJ, Krassay AA (2000) Constraining sequence stratigraphy in north Australian basins: SHRIMP U-Pb zircon geochronology between Mt Isa and McArthur River. Australian Journal of Earth Sciences 47:431-459. https://doi.org/10.1046/j.1440-0952.2000.00797.x

923 Paton C, Hellstrom J, Paul B, et al (2011) Iolite: Freeware for the visualisation and processing of mass spectrometric data. Journal of Analytical Atomic Spectrometry 26:2508-2518. https://doi.org/10.1039/C1JA10172B
926 Plissart G, Féménias O, Mărunțiu M, et al (2009) Mineralogy and geothermometry of gabbro-derived listvenites
927 in the Tisovita-Iuti ophiolite, Southwestern Romania. The Canadian Mineralogist 47:81–105.
928 https://doi.org/10.3749/canmin.47.1.81

929 Polito PA, Kyser TK, Alexandre P, et al (2011) Advances in understanding the Kombolgie Subgroup and
930 unconformity-related uranium deposits in the Alligator Rivers Uranium Field and how to explore for
931 them using lithogeochemical principles. Australian Journal of Earth Sciences 58:453–474.
932 https://doi.org/10.1080/08120099.2011.561873

933 Polito PA, Kyser TK, Jackson MJ (2006) The role of sandstone diagenesis and aquifer evolution in the formation
934 of uranium and zinc-lead deposits, southern McArthur basin, Northern Territory, Australia. Economic
935 Geology 101:1189–1209. https://doi.org/10.2113/gsecongeo.101.6.1189

936 Polito PA, Kyser TK, Marlatt J, et al (2004) Significance of alteration assemblages for the origin and evolution
937 of the Proterozoic Nabarlek unconformity-related uranium deposit, Northern Territory, Australia.
938 Economic Geology 99:113–139. https://doi.org/10.2113/gsecongeo.99.1.113

939 Polito PA, Kyser TK, Rheinberger G, Southgate PN (2005a) A Paragenetic and Isotopic Study of the Proterozoic
940 Westmoreland Uranium Deposits, Southern McArthur Basin, Northern Territory, Australia. Economic
941 Geology 100:1243–1260. https://doi.org/10.2113/100.6.1243

942 Polito PA, Kyser TK, Thomas D, et al (2005b) Re-evaluation of the petrogenesis of the Proterozoic Jabiluka
943 unconformity-related uranium deposit, Northern Territory, Australia. Mineralium Deposita 40:257–288.
944 https://doi.org/10.1007/s00126-005-0007-9

945 Rawlings DJ (1999) Stratigraphic resolution of a multiphase intracratonic basin system: the McArthur Basin,
946 northern Australia. Aust J Earth Sci 46:703–723. https://doi.org/10.1046/j.1440-0952.1999.00739.x

947 Rheinberger GM, Hallenstein C, Stegman CL (1998) Westmoreland uranium deposits. Geology of Australian
948 and Papua New Guinean mineral deposits: Melbourne, Australiasian Institute of Mining and Metallurgy
949 807–814

950 Richard A, Cathelineau M, Boiron M-C, et al (2016) Metal-rich fluid inclusions provide new insights into
951 unconformity-related U deposits (Athabasca Basin and Basement, Canada). Mineralium Deposita
952 51:249–270

953 Schindlmayr WE, Beerbaum B (1986) Structure related uranium mineralization in the Westmoreland district,
954 Northern Australia

955 Scott DL, Rawlings DJ, Page RW, et al (2000) Basement framework and geodynamic evolution of the
956 Palaeoproterozoic superbasins of north-central Australia: an integrated review of geochemical,
957 geochronological and geophysical data. Australian Journal of Earth Sciences 47:341–380.
958 https://doi.org/10.1046/j.1440-0952.2000.00793.x

959 Scrimgeour I, Kinny P, Close D, Edgoose C (2005) High-T granulites and polymetamorphism in the southern
960 Arunta Region, central Australia: Evidence for a 1.64Ga accretional event. Precambrian Research
961 142:1–27. https://doi.org/10.1016/j.precamres.2005.08.005

962 Shabaga BM, Fayek M, Quirt D, et al (2017) Mineralogy, geochronology, and genesis of the Andrew Lake
963 uranium deposit, Thelon Basin, Nunavut, Canada. Can J Earth Sci 54:850–868.
964 https://doi.org/10.1139/cjes-2017-0024

965 Sharpe R, Fayek M, Quirt D, Jefferson CW (2015) Geochronology and Genesis of the Bong Uranium Deposit,
966 Thelon Basin, Nunavut, Canada. Economic Geology 110:1759–1777.
967 https://doi.org/10.2113/econgeo.110.7.1759

968 Skirrow RG, Mercadier J, Armstrong R, et al (2016) The Ranger uranium deposit, northern Australia: Timing
969 constraints, regional and ore-related alteration, and genetic implications for unconformity-related
970 mineralisation. Ore Geology Reviews 76:463–503. https://doi.org/10.1016/j.oregeorev.2015.09.001
971 Stacey J, Kramers J (1975) Approximation of terrestrial lead isotope evolution by a two-stage model. Earth and planetary science letters 26:207-221. https://doi.org/10.1016/0012-821X(75)90088-6

973 Sun S-S, McDonough WF (1989) Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. Geological Society, London, Special Publications 42:313-345

975 Tera F, Wasserburg GJ (1972) U-Th-Pb systematics in three Apollo 14 basalts and the problem of initial Pb in lunar rocks. Earth and Planetary Science Letters 14:281-304. https://doi.org/10.1016/0012-821X(72)90128-8

978 Vigar AJ, Jones DG (2009) Laramide Resources Limited. Westmoreland Uranium project. Redtree Resource update/ Report MA823

980 Wall VJ (2006) Unconformity-related uranium systems: Downunder and over the top. ASEG Extended Abstracts 2006:1-12. https://doi.org/10.1071/ASEG2006ab189

982 Wilde AR, Mernagh TP, Bloom MS, Hoffmann CF (1989) Fluid inclusion evidence on the origin of some Australian unconformity-related uranium deposits. Economic Geology 84:1627-1642. https://doi.org/10.2113/gsecongeo.84.6.1627

985 Wilde AR, Wall VJ (1987) Geology of the Nabarlek uranium deposit, Northern Territory, Australia. Economic Geology 82:1152-1168. https://doi.org/10.2113/gsecongeo.82.5.1152

987 Wingate MT, Pirajno F, Morris PA (2004) Warakurna large igneous province: A new Mesoproterozoic large igneous province in west-central Australia. Geology 32:105-108. https://doi.org/10.1130/G20171.1

989 Wygralak AS, Ahmad M, Hallenstein CP (1988) Sedimentology of the Westmoreland Conglomerate, southern McArthur basin, Northern Territory, Australia. Journal of the Geological Society of Australia 35:195-207

992 Ypma PJM, Fuzikawa K (1980) Fluid inclusion isotope studies of the Nabarlek and Jabiluka uranium deposits, Northern Territory, Australia. In: Proceedings of International Uranium Symposium on the Pine Creek Geosyncline. IAEA, Vienna, pp 375-395

995 Zane A, Weiss Z (1998) A procedure for classifying rock-forming chlorites based on microprobe data. Rendiconti Lincei 9:51-56. https://doi.org/10.1007/BF02904455

997 Zhou J-L, Li X-H, Tang G-Q, et al (2018) Ca. 890 Ma magmatism in the northwest Yangtze block, South China: SIMS U-Pb dating, in-situ Hf-O isotopes, and tectonic implications. Journal of Asian Earth Sciences 151:101-111

1000
1001

1002 **Fig. 1** Schematic geological map of the study area, Northern Territory and Queensland, Australia. **a** Location of the McArthur and South Nicholson basins and Murphy Inlier in Australia. **b** Zoom indicating the location of the Alligator Rivers Uranium Field (ARUF) and the Westmoreland area (black rectangle). **c** Simplified geological map around the studied deposits of the Westmoreland area. After Ahmad et al. (2013).

1006
Fig. 2 Conceptual cross-section showing the position of the Junnagunna, Redtree, Huarabagoo and Long Pocket deposits in the stratigraphic sequence of the McArthur Basin, Australia (compilation after Fuchs and Schindlmayer 1981; Schindlmayer and Beerbaum 1986; International Atomic Energy Agency 2018a, b) and the approximate locations of the studied samples.

Fig. 3 Paragenesis of the Redtree, Junnagunna and Huarabagoo deposits (Westmoreland area, Australia); compiled from Polito et al. (2005a) and this study. Ages are based on zircon U-Pb ages (Page et al. 2000) for the detrital minerals, U-Pb, $^{207}$Pb/$^{204}$Pb and $^{40}$Ar/$^{39}$Ar ages from Polito et al. (2005a) and U-Pb and Pb/Pb ages on apatite and uraninite from this study. Italic: data from the paragenesis presented by Polito et al. (2005a); regular: mineral sequence observed in this study; bold: mineral sequence observed in this study and by Polito et al. (2005a). The temperatures estimates are from chlorite and illite thermometry (Polito et al., 2005a; this study) and from the $\Sigma$REE (this study).

Fig. 4 Photomicrographs and backscattered electron image (d) illustrating the diagenetic sequence for mineralized intercepts of the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). a Q0 detrital quartz grains coated by hematite (H1), and quartz overgrowths Q1, sample WDD10-151 27.9B. b Quartz partial dissolution (Q⁻) after Q1, sample JDD08-003 74.4A. c Rosette of chlorite C1 after partial dissolution of Q0 and Q1, sample WDD07-2 51.7A. d Diagenetic rutile R1 in a matrix formed by uraninite U4, sample WDD07-2 73.5B. e Illite alteration on chlorite C1, sample WDD9-128 53.3B. f Quartz Q2 after quartz Q1, sample WDD07-2 55.9B.

Fig. 5 Backscattered electron image images of uraninite mineral assemblages and location of the analyses from the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). For the location of the analyses: purple circle = LA-ICP-MS; blue cross = SIMS, dark green triangle = EPMA. a Uraninite U2 as matrix around apatite grains (Ap), sample JDD08-003 74.4B. b Uraninite U4 with altered zones, sample WDD07-2 55.9A. c Uraninite U4 in vein associated with chalcopyrite and pyrite, sample WDD10-151 29.4. d Uraninite U6 associated with clay minerals, sample WDD07-2 73.5B.

Fig. 6 EPMA maps of distribution of U, Pb, Ca, Fe and Si in uraninite for samples WDD07-2 55.9A (top) and WDD07-2 73.5B (bottom, Westmoreland area, Australia). The U- and Pb-rich zones correlate with Ca- and Si-poor zones. The element concentration values (in wt%) are estimated from *in-situ* spot EPMA analyses.

Fig. 7 Chemical compositions of uraninite generations from the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). a Pb vs Si + Ca + Fe. b Binary diagram of As vs Nb (in ppm). c Binary diagram of As vs W
(in ppm). **d** Binary diagram of V vs Ba (in ppm). **e** Rare Earth Element abundance (ΣREE) vs. REE fractionation ((ΣLREE/ΣHREE)N) relative to published values from Mercadier et al. (2011b), Alexandre et al. (2016) and Martz et al. (2019). REE fractionation normalised to chondrite values of Anders and Grevesse (1989).

**Fig. 8** Chondrite-normalised REE patterns of the different uraninite generations from the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). Each generation exhibits a unique REE pattern indicating different chemical environments, as previously observed for the other minor and trace elements (Fig. 7). The C1 Chondrite normalization is after Sun and McDonough (1989). **a** U2 generation. Sample JDD08-003 74.4B. **b** U4 generation. The LREE enrichment indicated by the arrow corresponds to alteration indicated by Pb loss. The numbers indicated to the left of the diagram corresponds to the LA-ICP-MS spots presented in Fig. 5. Sample WDD07-2 55.9A. **c** U4 generation. Sample WDD10-151 29.4. **d** U5 generation. Sample WDD08-106 10.6. **e** U6 generation. Sample WDD07-2 73.5B. **f** Uraninite data from unconformity-related basement-hosted deposit in the ARUF (Mercadier et al. 2011b) and the Camie River deposit, Otish Basin, Canada (Lesbros-Piat-Desvial et al. 2017).

**Fig. 9** $^{238}$U/$^{206}$Pb and $^{207}$Pb/$^{206}$Pb ratios measured and related age on diagenetic apatite from the sample JDD08-003 74.4B (Junnagunna deposit, Westmoreland area, Australia). **a** Detrital quartz Q0 and diagenetic apatite in a uraninite U2 matrix (BSE image). **b** Tera-Wasserburg Concordia diagram.

**Fig. 10** U-Pb Concordia diagrams from in-situ isotopic analysis by SIMS of uranium oxides from the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). Plots were constructed using ISOPLOT (Ludwig 2007) from the isotope ratios presented in Table SM11. **a** Sample JDD08-003 74.4B. **b** The black ellipses are the data obtained on a single uraninite and the grey ellipses are additional data from the same depth in the same drillcore. Samples WDD07-2 55.9A and WDD07-2 55.9B. **c** Sample WDD10-151 29.4 anchored at 0 ±5 Ma. **d** Sample WDD08-106 10.6 anchored at 0 ±5 Ma. **e** Sample WDD07-2 73.5B anchored at 0 ±5 Ma. **f** All data from this study (open circles, anchored at 0 ±5 Ma) and from Polito et al. (2005a, pink).

**Fig. 11** Chlorite composition and nomenclature for the Redtree, Junnagunna and Huarabagoo U deposits (Westmoreland area, Australia). **a** Fe-Mg diagram indicating that all Fe atoms are in the Fe$^{2+}$ state in the ferromagnesian chlorite. **b** Classification diagram after Plissart et al. (2009). **c** Al$^{VI}$-Al$^{IV}$ diagram for analysed chlorites. Structural formulae are calculated on the basis of 14 oxygen. **d** Classification diagram for rock-forming chlorites according to Zane and Weiss (1998). Compositions of chlorite from the Westmoreland area (this study; Polito et al., 2005a) are plotted together with
chlorite from deposits of the ARUF (Polito et al. 2004, 2005b; Beaufort et al. 2005; Gigon et al. 2019), the Athabasca Basin (Kotzer and Kyser 1995; Billault et al. 2002; Martz 2017) and the Thelon Basin (Grare et al. 2020) for comparison. For data from the ARUF and the Canadian basins, the plotted data corresponds to the mean compositions for each sample or generation.

**Fig. 12** Synthesis of the alteration and mineralisation ages obtained in the ARUF and in the Westmoreland area of the McArthur Basin (Australia). 1: Hills and Richards (1976). 2: Clauer et al. (2015). 3: Ludwig et al. (1987). 4: Skirrow et al. (2016). 5: Polito et al. (2004b). 6: Polito et al. (2005b). 7: Maas (1989). 8: Polito et al. (2005a). 9: this study.
<figure><img src="image_2.png" /><figcaption>Figure 1</figcaption></figure>

138°00′
<figure><img src="image_3.png" /><figcaption>Figure 2</figcaption></figure>
<table><caption>Figure 3</caption><thead><tr><th>Characteristics</th><th>Detrital minerals</th><th>Early diagenesis</th><th>Diagenesis</th><th>First stage uraninite</th><th colspan="5">Other U-forming stages</th></tr></thead><tbody><tr><td>Age / timing</td><td>ca. 1800 - ca. 1750 Ma</td><td>ca. 1750 - ca. 1680 Ma</td><td>ca. 1680 Ma</td><td>1655 Ma</td><td>ca. 1030 Ma</td><td>ca. 870 Ma</td><td>ca. 730 Ma</td><td>ca. 550 Ma</td><td>ca. 350 Ma</td></tr><tr><td>Est. temperature</td><td></td><td></td><td>&gt; 300 °C</td><td>150-250 °C</td><td>&gt; 350 °C</td><td>&lt; 350 °C</td><td>&lt; 350 °C</td><td>&lt; 350 °C</td><td>&lt; 350 °C</td></tr><tr><td>Quartz Q0</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Feldspar</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-, Mg-silicates</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-, Ti-oxides</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Hematite H1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Quartz Q1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Quartz dissolution</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Chlorite C1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Rutile R1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Dickite D1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Illite I1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Quartz Q2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Apatite</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pyrite P1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U1</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Hematite H2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Illite I2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Rutile R2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>JDD08-003 74.4B</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U3</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pyrite P2</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Galena</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U4</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WDD07-2 55.9</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WDD10-151 29.4</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Pyrite P3</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Chalcopyrite</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Quartz Q3</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U5</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WDD08-106 10.6</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Uraninite U6</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>WDD07-2 73.5B</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>
<figure><img src="image_4.png" /></figure>
<caption>Figure 5</caption>

<figure><img src="image_5.png" /><figcaption>b</figcaption></figure>

<figure><img src="image_6.png" /><figcaption>c</figcaption></figure>

<figure><img src="image_7.png" /><figcaption>d</figcaption></figure>

<figure><img src="image_8.png" /></figure>
<caption>Figure 6</caption>

Sample WDD07-2 55.9A

<figure><img src="image_9.png" /></figure>

Sample WDD07-2 73.5B

<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /><figcaption>Figure 7</figcaption></figure>
<figure><img src="image_12.png" /></figure>

<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>
<figure><img src="image_18.png" /><figcaption>Figure 9</figcaption></figure>

<figure><img src="image_19.png" /></figure>
<figure><img src="image_20.png" /><figcaption>Figure 10</figcaption></figure>
<figure><img src="image_21.png" /><figcaption>Figure 11</figcaption></figure>
<figure><img src="image_22.png" /><figcaption>Figure 12</figcaption></figure>
Click here to access/download
**Supplementary Material**
ESM1.pdf
Click here to access/download
**Supplementary Material**
ESM2.xlsx