©2012 Society of Economic Geologists, Inc.
Economic Geology, v. 107, pp. 401–425
Basin Evolution and Unconformity-Related Uranium Mineralization:
The Camie River U Prospect, Paleoproterozoic Otish Basin, Quebec
S. R. BEYER,1,† K. KYSER,1 E. E. HIATT,2 P. A. POLITO,3 P. ALEXANDRE,1 AND K. HOKSBERGEN4
1 Department
of Geological Sciences and Geological Engineering, Queen’s University, Kingston, Ontario, Canada K7L 3N6
2 Department
3 Anglo
of Geology, University of Wisconsin-Oshkosh, 800 Algoma Boulevard, Oshkosh, Wisconsin 54901
American Exploration Australia, Suite 1, 16 Brodie-Hall Drive, Bentley, WA 6102, Australia
4 Cameco
Corporation, 2121 11th Street West, Saskatoon, Saskatchewan, Canada S7M 1J3
Abstract
The Paleoproterozoic Otish Basin, Quebec, hosts several uranium prospects that until recently remained
underexplored and poorly understood. In this study, the Camie River U prospect, which shows similar characteristics to high-grade unconformity-related U deposits, is the focus of an integrated basin analysis in the western Otish Basin.
Conglomerate and sandstone of the Indicator Formation, which were deposited in at least six depositional
sequences, were affected by insignificant early diagenetic compaction and cementation. This allowed the formation of regional peak diagenetic aquifers, which became muscovite altered due to interaction with fluids
having δ18O and δ2H values similar to those of seawater-influenced basinal brines at 250°C.
U mineralization at Camie River occurred at 1721 ±20 Ma based on a 207Pb/206Pb date obtained by laser
ablation of uraninite, which coincides with a phase of the Otish Gabbro intrusion that has been dated at ca.
1730 Ma. The intrusive event promoted circulation of U-bearing basinal brines, triggering U mineralization at
several locations in the western Otish Basin. Interaction of basinal brines with the Otish Gabbro produced
coarse-grained chlorite, tourmaline, and phengitic muscovite, which decreased the fluid-conducting capabilities of diagenetic aquifers and resulted in fault zone- and fracture-dominated fluid flow.
Subsequent fluid alteration events produced limited U remobilization, sulfides, sudoite, and siderite
between ca. 1670 and 1410 Ma based on mineral paragenesis and 40Ar/39Ar dates of muscovite. Metamorphic
fluids having high δ18O values and temperatures around 300°C accompanied 1.2 to 1.0 Ga Grenville orogenesis and subgreenschist-grade metamorphism in the Otish Basin but were present at low water/rock ratios at
Camie River and therefore produced little alteration. Post-Grenville uplift of the Otish Basin likely produced
late, low-temperature alteration minerals that have been influenced by recent meteoric water, suggesting that
the fault zones and fractures the minerals occupy remain as preferential fluid-flow pathways to the present day.
Radiogenic Pb and the characteristic trace elements Mo + W + Nb have also preferentially dispersed from
the mineralization along fault zones, fractures, and depositional sequence boundaries, and can be used to
explore for Camie River-style U mineralization.
Introduction
THE PALEOPROTEROZOIC Otish Basin, Quebec, hosts several
uranium prospects that were discovered in the 1970s and
1980s but have since received little attention and remain underexplored. The Camie River U prospect is one such
prospect that has been the focus of revitalized U exploration
in the Otish Basin (Fig. 1). Of the several significant showings
discovered by Uranerz and joint venture partners in the western Otish Basin between 1974 and 1984 (Fig. 1), the Camie
River prospect was the only showing with characteristics similar to those of unconformity-related U deposits in the
Athabasca Basin, Saskatchewan (Gatzweiler, 1987). These
similarities include high-grade uranium mineralization located at the fault-offset unconformity between graphite-bearing metasedimentary rocks below and fluviatile conglomerate
and sandstone above, an alteration halo coincident with U
mineralization, and geochemical associations with Co, Ni, As,
and Cu, and other elements that can be associated with unconformity-related U deposits (Gatzweiler, 1987; Höhndorf
et al., 1987). The preliminary nature of these studies, however,
does not provide the necessary scope to evaluate basin-scale
† Corresponding
author: e-mail, [email protected]
0361-0128/12/4017/401-25
processes, such as establishment of fluid-flow pathways and
maturation of basinal fluids to produce oxidizing, metalleaching brines during diagenesis, which are required to form
large unconformity-related U deposits (cf. Kyser and Cuney,
2008a). Confirmation that such processes were operating in
the Otish Basin would indicate that unconformity-related U
mineralization in significant quantity may have occurred at
other locations in the basin and therefore increase the
prospectivity of the Otish Basin. Conversely, it is unclear
whether or not such processes acted similarly, or as efficiently,
in the Otish Basin as the Athabasca and Kombolgie Basins.
This study builds on the work of Gatzweiler (1987) and
Höhndorf et al. (1987) by placing the Camie River U prospect
in the context of basin evolution through an integrated basin
analysis. We present a sequence stratigraphic model for the
southwestern Otish Basin margin, which aids correlation of
potential fluid-flow pathways in the western Otish Basin.
Fluid evolution and the relative timing of fluid events in the
western Otish Basin are determined based on mineral crosscutting relationships and stable isotope geochemistry, and
summarized in a mineral paragenesis. The absolute timing of
significant fluid events, including U mineralization at Camie
River, is determined using U-Pb, Pb-Pb, and 40Ar/39Ar
401
Submitted: January 26, 2011
Accepted: October 16, 2011
402
BEYER ET AL.
74º
72º
70º
Canada
52º
o
Pr
USA
n
vi
ce
LG
BL
IL
M
Fig. 2
e
Zon
ic
i
in
l
L
M
is
ta
ss
A
lb
e
an
ac
il
le
Fr
o
nt
re
n
T
t
ec
l
vi
on
Camie
River
Pr
le
o
n
vi
ce
LEGEND
Grenville Province,
undivided
U showing
fault, unspecified
fault,
reverse/thrust
Mistassini Group
Temiscamie Fmn.
Upper Albanel Fmn.
Otish Group
Lower Albanel Fmn.
Otish Gabbro
Cheno Fmn.
Peribonca Fmn.
Papaskwasati Fmn.
Indicator Fmn.
km
Mistassini dikes
0
Superior Province,
greenstone belt
Superior Province,
undivided gneiss/intrusives
G
Gr
en
v
c
Mi
Su
Ba stas
pe
si
s
i
n
ri
ni
La
or
51º
Basin
Otish
N
50
FIG. 1. Generalized geologic map of the Otish and Mistassini Basins, Quebec, showing the location of the Camie River
unconformity-related U prospect in the Otish Basin. The mapped area shown in Figure 2 is also indicated. Other significant
U prospects: BL = Beaver Lake; IL = Indicator Lake; LG = Lorenz Gully; M = Matoush. Modified from Chown and Caty
(1973), Chown and Caty (1983), Fahrig and West (1986), and Gatzweiler (1987).
geochronology. Lastly, potential pathfinders for Camie Riverstyle mineralization are proposed based on the results of a 2%
HNO3 leach method (Holk et al., 2003), and guidelines for
exploration are presented.
Geologic Setting
The Paleoproterozoic Otish Basin is located approximately
300 km northeast of Chibougamau, Quebec. Both the Otish
Basin and the neighboring Mistassini Basin (Fig. 1) unconformably overlie an Archean basement complex consisting of
gneiss and migmatite, metavolcanic and metasedimentary
rocks (greenstone belts), and felsic plutonic rocks of the eastern Superior Province (Neilson, 1966; Chown, 1971; Chown
and Caty, 1973; Gatzweiler, 1987). Maximum ages ca. 2.55 Ga
(Rb-Sr and K-Ar on muscovite) for gneiss in the area are consistent with ages of the final assembly of the Superior craton
during the Kenoran orogeny (Höhndorf et al., 1987; Card,
1990). The Superior Province was intruded by the NW-trending Mistassini diabase dikes (Fig. 1), which are 2.2 to 2.0 Ga
in age (Fahrig and West, 1986; Fahrig et al., 1986) but may be
as old as 2.5 Ga according to unpublished work by Heaman
(2004).
The Archean basement complex was subaerially weathered
prior to formation of the Otish and Mistassini Basins, resulting
0361-0128/98/000/000-00 $6.00
in local high paleotopographic relief and extensive regolith
development (Chown and Caty, 1973, 1983). Basin formation
occurred sometime after ca. 2.2 Ga, as sedimentary rocks in
the Mistassini Basin unconformably overlie the Mistassini
dikes (Fahrig et al., 1986). Sedimentary rocks in the Otish and
Mistassini Basins belong to the Otish and Mistassini Groups,
respectively (Bergeron, 1957; Chown and Caty, 1973). The
Otish Group is composed of the Indicator Formation that is
conformably overlain by the Peribonca Formation (Fig. 1).
The Indicator Formation ranges from 300 to 1,000 m thick
and consists of conglomerate and sandstone (Chown and
Caty, 1973; Gatzweiler, 1987). The maximum observed thickness of the Peribonca Formation is 380 to 500 m, which consists of sandstone with variable dolomite cement, argillaceous
sandstone, and minor conglomerate and dolostone (Chown
and Caty, 1973; Gatzweiler, 1987). Höhndorf et al. (1987) and
Genest (1989) also described evaporitic horizons in the Peribonca Formation associated with magnesite and casts and
carbonate pseudomorphs after halite and gypsum.
The Otish and Mistassini Groups are intruded by dikes and
sills of the Otish Gabbro (Fahrig and Chown, 1973; Chown
and Archambault, 1987; Fig. 1). Wanless et al. (1965) reported a K-Ar age of 1465 Ma for the Otish Gabbro, but this
age likely represents a subsequent dike alteration event
402
403
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
Uranium mineralization in the Otish Basin
Approximately 30 uranium occurrences were discovered in
the Otish Basin in the 1970s and 1980s, which are associated
with basement veins and breccias, the sub-Otish Group unconformity, and intra-Otish Group faults and mafic intrusions
(Gatzweiler, 1987). Samples for this study were obtained
from the unconformity-related Camie River prospect, and
the fault-related Matoush deposit, both located in the western Otish Basin (Figs. 1, 2).
Archean basement rocks intersected at the Camie River
prospect (Fig. 3a) belong to the Hippocampe greenstone belt
(Genest, 1989) and consist of foliated quartz + chlorite ±
muscovite metapsammites interbedded with tightly folded
graphitic, cherty schists and zones of massive sulfides. These
metasedimentary rocks are structurally underlain by porphyritic metagabbros. Basement rocks are unconformably
overlain by conglomerates and sandstones of the Indicator
Formation. Uraninite and brannerite mineralization occurs
where the unconformity is in reverse-faulted contact with
basement massive sulfide and graphitic schist and extends between 20 and 50 m above and below the unconformity along
fractures (Gatzweiler, 1987). Uranium mineralization is associated with Mo, Nb, Co, Ni, As, Cu, Se, V, Ag, and Au, some of
which may be attributed to pyrochlore, columbite, and molybdenite, the latter of which follows uraninite (Gatzweiler, 1987;
Höhndorf et al., 1987). The best intersection was 7.83% U3O8
over 1 m (Gatzweiler, 1987). Chlorite, carbonate, pyrite, tourmaline, muscovite, and albite are reported from a zoned alteration halo, but paragenetic relationships are not indicated
(Ruzicka and LeCheminant, 1984; Gatzweiler, 1987; Höhndorf et al., 1987). Uraninite at Camie River was dated at 1723
±16 Ma using U-Pb methods (Höhndorf et al., 1987).
Uranium mineralization at the nearby Matoush deposit
(Fig. 3b) is focused by a major NNE-striking fault within the
0361-0128/98/000/000-00 $6.00
AM06 LEGEND
AM08
Otish Group,
N
AM21
5750000N
undivided
AM15
Archean
greenstone,
undivided
Eas
ac tmai
ces n Mi
s ro
n
ad
e
5755000N
5760000N
AM20
section line, fig. 3b
Matoush
deposit
Archean
gniess/intrusives,
undivided
exploration
drill hole
outcrop
(Fig. 3b)
EM conductor
EC08-003
OTS07-001
fault
0
5745000N
OTS07-002
Cameco
Otish South Camp
5740000N
Camie River
U prospect
5
OM49
OM09
OT08-1
section line, fig. 4
690000E
km
Camie River
(Fahrig and Chown, 1973). Gatzweiler (1987) and Höhndorf
et al. (1987) reported an undocumented Sm/Nd age of 1730
±10 Ma for the Otish Gabbro. Chown and Archambault
(1987) infer an age of ca. 1750 Ma based on a 1787 ±55 Ma
Rb/Sr age of the Temiscamie Formation (Fryer, 1972), which
they interpreted as a thermal event associated with intrusion
of the Otish Gabbro. More recently, the unpublished work of
Hamilton and Buchan (2007) interpreted the Otish Gabbro
to be ca. 2169 Ma. In this study we rely on the 1750 to 1730
Ma ages suggested in published reports by Gatzweiler (1987),
Höhndorf et al. (1987), and Chown and Archambault (1987),
and consider sedimentation in both basins therefore to have
occurred between about 2200 and 1730 Ma.
Gneiss, metavolcanic rocks, and anorthosite of the 1.2 to
1.0 Ga Grenville Front Tectonic Zone are in faulted juxtaposition with the Superior Province to the southeast (Chown
and Caty, 1973; Davidson, 1984; Rivers, 1997; Fig. 1). In the
Otish Group, Grenville Front-related faults show strike-slip
motion and trend 060 (dextral) and 010 (sinistral), the latter
dominating toward the southwest (Chown, 1984). Fault zones
were metamorphosed to amphibolite grade adjacent to the
Grenville Front and display ductile textures (Chown, 1979,
1984). In contrast, fault zones display subgreenschist-grade
metamorphism and brittle textures away from the Grenville
Front (Chown, 1971, 1979, 1984; Rivers and Chown, 1986).
OM47
OTS07-004
OTS07-003
Fig. 3a
695000E
700000E
705000E
FIG. 2. Generalized geologic map of the west-central Otish Basin (refer to
Fig. 1 for location), showing the location of the Camie River U prospect and
the Matoush deposit. Samples were collected from the outcrop and all exploration drill holes indicated. The mapped area shown in Figure 3a is also
indicated. EM = electromagnetic. Coordinates are for UTM projection (zone
18, NAD83).
Indicator Formation. Underlying basement rocks are barren
and consist of relatively unaltered tonalite and granodiorite.
Mineralization is dominated by uraninite that occurs as a series of subparallel, plunging shoots along the fault plane
(Gatzweiler, 1987; Fig. 3b) and a 10- to 50-m-wide zoned alteration halo that envelopes the fault and consists of tourmaline, chlorite, Cr-V−bearing muscovite, hematite, and
limonite (Gatzweiler, 1987; Höhndorf et al., 1987). Recent
exploration indicates that the deposit hosts an estimated indicated mineral resource of 436,000 metric tons (t) grading
0.78% U3O8, containing 7.46 million pounds (Mlbs) of U3O8
(Strateco Resources press release, 9 November 2010).
Methods
Archean basement rocks and Indicator Formation conglomerate and sandstone were collected from historic exploration drill core at Camie River and Matoush, recent drill
core at Camie River, and from one outcrop locality in the
western Otish Basin (Figs. 2, 3). Uranium mineralization
from historic core was not available for sampling but was
studied in recent Camie River drill core, where it exists solely
in the basement rocks. Seven stratigraphic sections were
measured from drill core at decimeter to meter scales. Approximately 180 polished thin sections were made and studied using standard transmitted- and reflected-light microscopy, and electron microscopy to identify minerals and
403
404
BEYER ET AL.
W
LEGEND
orebody trace
e River
Cami
Archean
greenstone,
undivided
exploration
drill hole
EM conductor
AM06 AM21
AM15
AM20
5742000N
Otish Group,
undivided
E
unconsolidated
glacial sediments
AM08
N
AM-15 zone
OM-09
OM-47
0
lithofacies
boundary
meters 500
OTS07-004
a
100
meters
5741000N
100 meters
U mineralization
707000E
MT-22/MT-34
zones
MT-34
zone
708000E
Indicator Formation −
conglomerate & sandstone
b
tonalite/granodiorite
Matoush
Fault
unconformity
FIG. 3. a. Generalized geologic map of the Camie River U prospect (refer to Fig. 2 for location). Samples were collected
from labeled exploration drill holes. U mineralization was collected from drill hole OTS07-004. Coordinates are for UTM
projection (zone 18, NAD83). b. Generalized cross section of the Matoush deposit (refer to Fig. 2 for location), showing the
location of drill holes sampled in this study in relationship to the Matoush fault. No U mineralization was available for sampling in these holes.
crosscutting relationships. Scanning electron microscopy was
performed using an Amray 1830 scanning electron microscope (SEM) in backscattered electron (BSE) mode, Queen’s
University, Canada. Qualitative chemical analyses were determined using an Oxford energy dispersive X-ray spectrometer
(EDS) with an Si(Li) detector crystal coupled to the SEM.
Quantitative chemical analyses of minerals were determined by electron microprobe analysis (EMPA) at Carleton
University, Canada, using an automated 4- spectrometer
Cameca Camebax MBX electron microprobe by the wavelength dispersive (WDS) X-ray method. Specimens were analyzed using a rastered electron beam 5 to 10 micrometers in
size, with a 15 kv accelerating voltage and 20 nA current. Elemental weight percentages were calculated using a Cameca
PAP matrix correction program. Analyses are accurate to 1 to
2% relative for major elements (>10 wt %), 3 to 5% relative
for minor elements (>0.5−<5.0 wt %). As detection limits
were approached (<0.1 wt %), relative errors approached
100%.
Phyllosilicates were separated from crushed samples by ultrasound disintegration and centrifugation for stable isotope
analysis and 40Ar/39Ar geochronology. Field identification of
phyllosilicates in drill core was aided by the use of an ASD
TerraSpec® short wave infrared spectrometer and AusSpec
TSG™ software. All phyllosilicate separates were analyzed by
X-ray diffraction (XRD) to determine mineralogical composition. XRD was performed using a Philips X-Pert instrument
at Queen’s University, Canada.
Oxygen isotope compositions of phyllosilicates were measured using a dual inlet Finnigan MAT 252 isotope ratio mass
spectrometer (IRMS) following oxygen extraction using BrF5
(Clayton and Mayeda, 1963). Hydrogen isotope compositions
0361-0128/98/000/000-00 $6.00
of phyllosilicates were determined using a ThermoFinnigan
TC/EA and a Deltaplus XP IRMS. Oxygen and hydrogen isotope ratios are reported in δ notation in units of per mil relative to Vienna Standard Mean Ocean Water (V-SMOW). δ18O
and δ2H analyses were reproducible to 0.2 and 3‰, respectively. Oxygen isotope fractionation factors used throughout
this paper are those proposed by Wenner and Taylor (1971)
for water-chlorite, and O’Neil and Taylor (1969) for watermuscovite. Hydrogen isotope fractionation factors used are
those proposed by Taylor (1974) for water-chlorite, and Vennemann and O’Neil (1996) for water-muscovite.
Muscovite temperatures of formation were estimated using
the molar fraction of pyrophyllite (Xprl) as determined by
EMPA (Cathelineau, 1988). Chlorite temperatures of formation were estimated using tetrahedral site occupancy as determined by EMPA (Cathelineau, 1988). However, the chlorite geothermometer of Zang and Fyfe (1995), a variant of the
chlorite geothermometer of Cathelineau (1988), was preferred, as the chlorites studied have Fe/(Fe + Mg) ratios
closer to those described in the former. Both muscovite and
chlorite geothermometers are accurate to within 25°C based
on replicate analyses.
Uranium-Pb and Pb-Pb isotope ratios were determined
using laser ablation, high-resolution, multicollector, inductively coupled plasma mass spectrometry (LA-HR-MCICPMS), using a ThermoFinnigan NEPTUNE instrument
equipped with a frequency-quintupled (213 nm) Nd-YAG
(LUV 213, New Wave-Merchantek) laser ablation system at
the Queen’s Facility for Isotope Research (QFIR), following
the procedures of Chipley et al. (2007). Uraninite ablation
was performed on polished thin sections using a 12- to 40-µm
spot size with 40 to 70% laser power at a frequency of 2 Hz.
404
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
Measurements on inhouse davidite and uraninite standards
and gas blanks bracketed individual sample measurements.
For each sample, 204Pb, 206Pb, 207Pb, 235U, and 238U were measured and corrections for common Pb were made scan-byscan for each spot.
40Ar/39Ar geochronology was performed at the Rare Gas
Geochronology Laboratory, University of Wisconsin-Madison, United States, using methods summarized by Smith et al.
(2003). Plateau ages were calculated using not less than 70%
of 39Ar released and three consecutive steps that overlap in
their 1σ error margin.
U-Pb and Pb-Pb isotope ratios and the concentrations of 54
elements were determined for 50 samples using 2% HNO3
leach and high-resolution inductively coupled plasma mass
spectrometry (HR-ICPMS), following the method of Holk et
al. (2003).
Results
Sedimentology and sequence stratigraphy of
the Indicator Formation
The Indicator Formation was studied along a 15-km transect trending approximately southwest (Fig. 4) and consists of
two lithofacies that are described and interpreted in Table 1.
Lithofacies 1 consists of relatively thickly bedded, massive to
trough crossbedded cobble to granule conglomerates. Clasts
are dominated by quartz, but can contain up to 25% feldspar
in some beds. The coarse grain size and sedimentary structures suggest lithofacies 1 was deposited in high-energy
braided stream channels. Lithofacies 2 contains more thinly
bedded, rippled to laminated, medium- to coarse-grained
quartz arenite to subarkosic arenite. These sandstones were
OTS07-003
OTS07-002
OTS07-001
EC08-003
unconsolidated
glacial sediments
2
6
datum
1
unconformity
5
405
deposited in shallow braided stream channels, or as unconfined sheetfloods in a braid delta setting, based on the reduced grain size relative to lithofacies 1, sedimentary structures, and intermittent mudstone content.
The Indicator Formation contains a hierarchy of bedding
discontinuities that aid sequence stratigraphic correlation
along the transect (Fig. 4). The contact between Archean
basement rocks and the Indicator Formation is the lowest ordered (most stratigraphically significant) discontinuity and is
a type 1 sequence boundary (Van Wagoner et al., 1988) that
represents a prolonged period of subaerial exposure and erosion of basement rocks (Chown and Caty, 1983). The next
highest ordered discontinuities are surfaces within the Indicator Formation across which grain size and proportion of
basement-sourced cobbles to pebbles substantially increase.
Several of these surfaces are correlated among the thicker
sections to the north and bound six depositional packages
that are between 75 and 2 m thick (Fig. 4). Each package
shows an upward increase in the proportion of lithofacies 2
sandstone. Toward the Camie River prospect to the south,
only the lowest two depositional packages are preserved due
to postdepositional reverse faulting (Fig. 4). The highest ordered bedding discontinuities consist of sharp, often erosive
surfaces and pebble- to granule-armored lags across which
there is little to no change in grain size or clast composition.
These surfaces are numerous, but laterally restricted and
likely represent autogenic processes such as channel avulsion, and therefore are of minimal sequence stratigraphic
relevance.
Lithofacies 1 and mudstone-free sandstones of lithofacies 2
contain relatively coarse-grained, poorly to moderately sorted
lithologic units that are potential diagenetic aquifers because
they do not become pervasively cemented during early diagenesis and are capable of conducting diagenetic basinal fluids
(Hiatt et al., 2003). The Indicator Formation along the transect does not contain thick diagenetic aquitards, such as mudstone or well-sorted, cemented sandstone that are barriers to
diagenetic fluid flow (Hiatt et al., 2003). However, mudstonebearing lithofacies 2b may partition the diagenetic aquifers
outside of the study area to the northwest (Fig. 4).
v
4
Lithofacies 2b
3
2
Lithofacies 2a
Lithofacies 1a & 1b
1
un
5 kilometers
100
meters
1 = depositional
sequence
co
nfo
rmity
qtz.+ chlorite metapsammite
porphyritic metagabbro
FIG. 4. Cross section of the southwest Otish Basin margin, showing the
distribution of lithofacies 1 conglomerates and lithofacies 2a and 2b sandstones within a sequence stratigraphic correlation. At least six depositional
sequences are recognized and are bounded by stratigraphically significant
surfaces across which the proportion of coarse-grained, basement-sourced
detritus sharply increases. Refer to Figure 2 for the location of drill holes.
0361-0128/98/000/000-00 $6.00
Mineral paragenesis
Archean basement rocks and the Indicator Formation were
altered throughout basin evolution resulting in numerous
mineral phases (Fig. 5). The distribution of key alteration
minerals that reflect fluid/rock interaction is presented in
Figure 6.
Mineralogy of unaltered host rocks: Unaltered metapsammitic basement rocks intersected at Camie River are green to
gray and typically consist of around 60 to 70% quartz, 30 to
40% chlorite (C0) and minor muscovite (M0), the latter two
phases being coarse grained and foliation controlled. Quartz
has a fine-grained, interlocking, microgranular texture. The
metapsammite contains several additional phases that follow
quartz + C0 ± M0 but appear to be the product of early
events that predate the Otish Basin based on crosscutting relationships with later phases. These are centimeter-scale-wide
veins of calcite (Ca0) and epidote, anhedral fine-grained rutile, and decimeter-scale zones of euhedral, coarse-grained
pyrite (S0) and magnetite.
405
406
BEYER ET AL.
TABLE 1. Description and Interpretation of Lithofacies
Lithofacies
Description
1. Conglomerate
2. Sandstone
Sedimentary
structures
Bedding
Interpretation
1a Cobble to pebble conglomerate, clast to
matrix supported, oligomictic; clasts are
rounded and dominated by quartz, with minor
K-feldspar and greenstone; matrix is very
coarse to medium grained, poorly to moderately
sorted quartz arenite to subarkosic arenite
1.0- to 2.0-m-thick
bedsets defined by crude
fining-upward successions to pebble- to
granule-bearing very
coarse grained sand
Massive, trough
crossbedding
Cobble to pebble
conglomerate was
deposited in braided
streams as channel
lags that were variably
reworked
1b Pebble to granule conglomerate, matrix
supported, oligomictic; clasts are dominated
by quartz, with 3–25% K-feldspar and minor
lithic clasts; matrix is very coarse pebble- to
granule-bearing sand
0.5- to 2.5-m-thick bedsets defined by common
fining-upward successions to variably granulebearing very coarse to
coarse-grained sand
Trough crossbedding; ripple cross
lamination in
sandstone
Pebble to granule
conglomerate and
subordinate sandstone
were deposited as bedforms in braided stream
channels
2a Medium- to coarse-grained, rarely fine to
medium grained, moderately to well-sorted
quartz arenite to subarkosic arenite
0.5- to 1.0-m-thick
bedsets
Planar lamination,
ripple cross lamination, trough
crossbedding
Sandstone was deposited
by shallow streams in a
braid delta depositional
environment, or as
unconfined sheetfloods
2b Medium- to coarse-grained, rarely fine to
medium-grained, moderately to well-sorted
quartz arenite to subarkosic arenite with common mm-scale mudstone laminae, mudstone
rip-ups, and cm-scale intervals of mudstone
0.5- to 1.0-m-thick
bedsets
Ripple cross lamination, trough
crossbedding;
desiccation cracks
in mudstone
Mud-bearing sandstone
was deposited by shallow,
intermittent streams in a
braid delta depositional
environment
HOST ROCKS
DIAGENESIS
PEAK
EARLY
POST-DIAGENETIC
ALTERATION
LATE
pyrite - S1
rutile - R1
apatite - P1
+REE-Th-Y phosphate
BASEMENT
BASIN
*tourmaline
*Fe-chlorite - C1
Ms + Pg
250˚C
Ph
300˚C
Grenville orogeny
muscovite - M1
intrusion of Otish Gabbro
quartz cement - Q2
albite cement
*calcite (in veins) - Ca1
kaolinite - K1
fracturing
deposition of Otish Group
Fe-oxide (rims) - F1
quartz (overgrowths) - Q1
fracturing
quartz - Q0
K-feldspar
zircon
monazite - P0
rutile - R0
300˚C
*sudoite - C2
*illite + montmorillonite
*kaolinite - K2
*Fe-oxide - F2
<150˚C
Mg-chlorite - C0
quartz - Q0
muscovite - M0
*calcite+epidote (in veins)
magnetite
pyrite+pyrrhotite - S0
+galena ±chalcopyrite
*dolomite
*muscovite - M1
rutile - R1
apatite - P1
*uraninite - U1a
*brannerite (±Nb) - U1b
*minnesotaite
*tourmaline
*U-phosphate - U2
*molybdenite - S2a
*pyrite - S2b
*galena - S2c
*siderite (in veins)
*kaolinite - K2
2.5
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time (Ga)
1.75
406
1.6
1.2
FIG. 5. Paragenesis of host-rock minerals and
minerals associated with early diagenesis, peak to
late diagenesis, and postdiagenetic alteration events.
The relative timing of minerals was determined by
petrographic observations. Temperatures are calculated from EMPA data (see text for discussion). Ms =
muscovite; Pg = paragonite; Ph = phengite; * = preferential distribution in faults, fractures, or along the
basal unconformity.
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
OTS07-002
a
OTS07-001
EC08-003
unconsolidated
glacial sediments
OTS07-004
6
2
datum
1
5
Camie River
U prospect
4
3
M1 muscovite
paragonite
albite
2
un
5 kilometers
100
meters
1 = depositional sequence
con
1
formit
y
pervasive M1 in
basement rocks
OTS07-003
OTS07-004
b
albite
M1
muscovite
unconformity
fractures
major fault
basement rocks
W
U
E
FIG. 6. a. Cross section of the southwest Otish Basin margin, showing the
distribution of diagenetic alteration minerals (determined by petrography,
EMPA, SWIR spectrometry, and XRD). M1 muscovite marks extensive diagenetic aquifers in the Indicator Formation. b. Diagrammatic longitudinal
section showing the distribution of albite and M1 muscovite in relationship
to fractures and faults near the Camie River U prospect (indicated as “U”).
Refer to Figure 2 for the location of drill holes.
Graphite-bearing, black schistose metasedimentary basement rocks occur in meter-scale intervals at Camie River and
contain 60 to 90% C0 chlorite, with the remainder consisting
of quartz with a similar texture to that in metapsammitic rocks,
in addition to quartz in cherty lenses. Zones of massive sulfide
(S0) can replace up to 80% of this host rock and consist of
pyrite, pyrrhotite, galena, and minor chalcopyrite. Graphite
ranges in abundance from micrometers-thick, lustrous coatings on foliation planes to meter-scale intervals of graphitic
metapelite. Dolomite appears to heal a brecciated zone that
is a few meters thick within massive sulfides at the Camie
River U prospect. The dolomite paragenetically follows C0
chlorite and M0 muscovite but predates U mineralization.
Detrital phases in the Indicator Formation are dominated
by quartz (Q0), with up to 25% K-feldspar. Minor phases are
zircon, monazite (P0), and rutile, which can locally form up to
10% of the matrix in lithofacies 1 conglomerate, typically
within several meters from the unconformity.
Early diagenesis: Early diagenesis occurred under conditions of shallow burial in the Indicator Formation and produced mechanical compaction of detrital grains, minor
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amounts of Fe oxide (F1) rims, and syntaxial quartz cement
(Q1) as overgrowths on Q0 detrital quartz. Q1 cement is rare,
as most samples display evidence of partial to complete resorption of Q1 cement during later diagenesis.
Peak diagenesis: Deep burial of the Indicator Formation
resulted in the formation of extensive, well-developed secondary porosity that is occluded by several mineral phases.
Some of these phases also affected Archean basement
rocks. The earliest peak diagenetic minerals are quartz (Q2)
and albite cements (Fig. 7a) that are restricted to the Indicator Formation. Q2 cement is generally rare, restricted to
stratigraphically higher strata, and consists of dense mosaics of microcrystalline quartz. Albite cement is absent at
Matoush but is more prevalent at Camie River (Fig. 6a).
Near the Camie River prospect, albite pervasively occludes
secondary porosity and imparts a reddish-pink color to drill
core due to micrometer-scale Fe oxide inclusions within the
albite. Albite cement, and to a much lesser extent Q2, also
heals an early generation of fractures. Veins of calcite (Ca1)
in the Indicator Formation are millimeters wide and crosscut albite/Q2 quartz-healed fractures. Ca1 calcite also occurs in patches that are 10s to 100s of µm in size near the
veins (Fig. 7a).
Kaolinite (K1) and muscovite (M1) follow Ca1 calcite paragenetically. K1 fills secondary porosity and displays vermiform texture, is rare, and preferentially occurs in samples
near Matoush. M1 is the most pervasive and widely distributed peak diagenetic phase, and ranges from coarse to fine
grained (Fig. 7b). In the Indicator Formation, M1 imparts a
conspicuous pale-green color to drill core, partially to entirely replaces albite cement and K1 kaolinite where present,
and crosscuts Ca1 calcite (Fig. 7a). At the Camie River
prospect, M1 completely replaces albite in extensively fractured and faulted conglomerate within 30 m of the unconformity (Fig. 6b). M1-filled secondary porosity ranges from
20 and 35% and results in common “floating grain” texture
that is observed both in core and thin section (Fig. 7b). In
basement rocks, M1 muscovite commonly replaces quartz +
chlorite metapsammite directly beneath the unconformity
(Fig. 6a), converting the rock to pale-green schist, and gradually decreases in proportion with increasing depth over several meters.
M1 muscovite is followed by a suite of phosphates including monazite, apatite, and xenotime (P1) that occur as micrometers-sized overgrowths on detrital monazite and zircon
in lithofacies 1 conglomerate, and as anhedral masses tens of
micrometers in size within albite and M1 muscovite-filled
secondary porosity. In basement rocks, only rare anhedral apatite occurs. Euhedral pyrite (S1), with crystals that are hundreds of micrometers in size, is coeval with P1 phosphates in
the Indicator Formation and occurs within several meters of
the unconformity. Anhedral rutile (R1) also occurs during this
stage in both basement rocks and the Indicator Formation
and shows petrographic evidence to suggest it may have resulted from local remobilization and recrystallization of detrital and accessory R0 rutile.
Late diagenesis: Late diagenesis is characterized by uranium mineralization, followed by formation of coarse-grained
phyllosilicates and tourmaline in both basement rocks and the
Indicator Formation at Camie River. Uraninite (U1a) is the
407
408
BEYER ET AL.
a
b
c
d
e
f
FIG. 7. a. Detrital quartz (dq) grains in the Indicator Formation are cemented by albite, which is in turn crosscut by
patches of calcite (Ca1). Arrows indicate where M1 muscovite penetrates both albite and Ca1. Q1 quartz overgrowths are
absent due to complete resorption (sample OTS07-002-619.9m; cross-polarized transmitted light). b. Fine- to coarse-grained
M1 muscovite fills well-developed secondary porosity; dq = detrital quartz (sample OTS07-003-153.0m; cross-polarized
transmitted light). c. Fine-grained M1 muscovite is crosscut by coarse-grained, prismatic tourmaline (Tur) and chlorite (C1).
Arrows indicate where C1 is intergrown with coarse-grained, neoformed M1 muscovite (sample OTS07-002-621.4m; planepolarized transmitted light). d. Uraninite (U1a) replaces S0 pyrite and galena and paragenetically follows chlorite (C0) and
dolomite (sample OTS07-004-174.5m; plane-polarized reflected light). e. Uraninite (U1a) paragenetically follows chlorite
(C0) and dolomite, and replaces coarse-grained galena (S0). U1a is crosscut by fine-grained galena (S2c) (sample OTS07004-174.5m; plane-polarized reflected light). f. Brannerite (U1b) paragenetically follows chlorite (C0) and dolomite and partially replaces pyrite and galena (S0). Molybdenite (S2a) follows U1b (sample OTS07-004-174.3m; plane-polarized reflected
light).
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BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
most abundant U-bearing mineral, representing ~90% of
total U, and extensively replaces S0 pyrite and galena, a relationship that is conspicuous both in thin section (Fig. 7d, e)
and drill core. Paragenetically, U1a succeeds the assemblage
C0 + dolomite + M1 based on petrography. Brannerite (U1b)
is a relatively minor phase and was identified qualitatively
using EDS. It represents ~10% of total U mineralization, occurs at the periphery of the U1a mineralized zone, and is paragenetically indistinguishable from U1a. EDS spectra of U1b
revealed detectable to significant Nb contents. The highest
Nb contents were obtained on rare, coarse, rectangularshaped grains that show replacement textures (Fig. 7f), indicating that U1b may have replaced an earlier mineral such as
columbite, which was observed by Gatzweiler (1987) and
Höhndorf et al. (1987).
In the Indicator Formation, tourmaline is most abundant
within several meters of the unconformity throughout the
study area. Tourmaline can occupy up to 30% of M1 muscovite-filled secondary porosity as euhedral to subhedral
prisms that are tens of micrometers wide and hundreds of
micrometers long and rarely as clusters of radiating prisms.
Toward Camie River, tourmaline displays characteristically
strong olive-green pleochroism (Fig. 7c). Chlorite (C1) occurs
within a few meters of the unconformity in the Indicator Formation throughout the study area, where it can be intergrown
with coarse-grained, euhedral laths of M1 muscovite (Fig.
7c). C1 chlorite also occurs in fault zones that affect the Indicator Formation near the Camie River prospect.
In basement rocks, tourmaline is restricted to intensely
M1-altered metapsammite within a few meters of the unconformity. Minnesotaite is a minor phase that was observed
proximal to U mineralization in basement rocks, where it
crosscuts S0 sulfides and dolomite. C1 chlorite was not observed in basement rocks.
Postdiagenetic alteration: The U1a + U1b uranium minerals are paragenetically followed by sulfide phases in basement
rocks. Molybdenite (S2a) crosscuts U1 uraninite and brannerite and is present as fields of micrometer-scale, euhedral
laths and rarely in veinlets (Figs. 7f, 8a). Pyrite (S2b) is present in veinlets and fine anhedral masses. Galena (S2c) is
common as micrometer-scale, anhedral to subhedral inclusions within U1a uraninite and U1b brannerite (Fig. 7e), and
in veinlets. An unidentified U phosphate mineral (U2) is a
minor phase that follows minnesotaite in basement rocks (Fig.
8b). Siderite occurs in veinlets that crosscut both dolomite
and U1a + U1b + U2 uranium minerals (Fig. 8b, c) in the
a
b
c
d
FIG. 8. a. Uraninite (U1a) follows chlorite (C0), replaces S0 galena, and is in turn crosscut by molybdenite (S2a) (sample
OTS07-004-174.5m; plane-polarized reflected light). b. Uraninite (U1a) and minnesotaite (Mns) follow S0 pyrite. Mns is followed by U phosphate (U2), which is in turn crosscut by siderite (sample OTS07-004-174.3m; BSE-SEM image). c. Uraninite (U1a) follows C0 chlorite and dolomite and replaces S0 pyrite. All phases are crosscut by siderite veinlets (sample
OTS07-004-174.5m; BSE-SEM image). d. Detrital quartz (dq) grains in a fracture zone are rimmed by sudoite (C2) that
crosscuts M1 muscovite (sample OTS07-004-126.0m; BSE-SEM image).
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BEYER ET AL.
basement rocks. Chlorite (C2) is restricted to the Indicator
Formation and occurs in fractures that produced cataclastic
deformation of M1-muscovite-altered sandstone (Fig. 8d)
and that partially propagated along older fractures associated
with Ca1 calcite. C2 chlorite also fills voids produced from
the dissolution of P1 phosphates.
Illite after M1 muscovite, montmorillonite after C1 chlorite, and kaolinite (K2) are postdiagenetic alteration phases
that preferentially affect fault zones in the Indicator Formation and, to a much lesser extent, basement rocks near the
Camie River prospect. Illite is similar to fine-grained M1
muscovite in appearance and was distinguished by its characteristic interlayer site vacancy via EMPA. Montmorillonite
was identified in a C1 chlorite-altered fault zone by XRD,
specifically by the expansion of the 14Å peak to 17Å on glycolation. Late-occurring, rusty-colored Fe oxide (F2) is a
relatively minor phase that affects the Indicator Formation
throughout the study area.
Mineral chemistry
Prediagenetic minerals in the basement rocks: C0 chlorite
has the highest MgO contents of the three varieties of chlorite at Camie River, averaging 22.5 wt % (Table 2), and has a
composition similar to that of clinochlore (Deer et al., 1992;
Fig. 9). The estimated temperature of formation of C0 is
around 300°C (Fig. 10), which is a minimum temperature as
it approaches the upper limit of the geothermometer (Cathelineau, 1988). Dolomite has stoichiometric Ca contents of one
atom per formula unit. Approximately two of every 10 Mg
atoms are replaced by Fe and Mn in dolomite (Table 3).
Peak diagenetic minerals: Albite cement in the Indicator
Formation contains an average Na2O content of 11.6 wt %,
TABLE 2. Electron Microprobe Analyses of Chlorite, Tourmaline, and Albite at Camie River
Sample no.
1
n
Oxide (wt %)
SiO2
Al2O3
FeO
MnO
MgO
TiO2
Cr2O3
V2O3
BaO
CaO
Na2O
K2O
Cl
F
O⬅Cl
O⬅F
Total
8
Atomic proportions
number of oxygens
Tetrahedral sites
Si
AlIV
Sum
Octahedral sites
AlVI
Fe
Mn
Mg
Cr
Sum
26.97
22.48
13.31
0.36
22.48
0.02
0.02
0.03
n.a.
<D.L.
<D.L.
<D.L.
<D.L.
0.13
0
0.06
85.79
±
2
11
0.51
0.54
0.73
0.06
1.17
0.02
0.02
0.03
0.07
0.03
0.90
28
23.05
21.78
30.22
0.31
9.53
<D.L.
<D.L.
0.07
n.a.
0.04
0.04
0.03
<D.L.
<D.L.
0
0
85.12
5.43
2.57
8.00
0.07
0.07
2.77
2.24
0.06
6.75
0.00
11.83
0.10
0.15
0.01
0.26
3
±
12
0.34
0.31
0.35
0.04
0.38
0.02
0.01
0.02
0.01
0.65
28
0.04
0.04
2.95
5.68
0.06
3.19
0.00
11.88
0.06
0.07
0.01
0.12
42.70
30.43
2.10
0.02
11.55
<D.L.
<D.L.
n.a.
n.a.
0.95
0.03
0.36
0.04
n.a.
0.01
0
88.19
1.71
1.77
0.58
0.02
1.07
0.02
0.01
0.15
0.02
0.13
0.01
0.01
1.72
0.04
0.04
16.00
±
36.74
32.34
10.65
<D.L.
4.20
0.36
0.03
0.03
n.a.
0.10
1.86
0.03
<D.L.
0.01
0
0
86.38
7.50
0.50
8.00
0.32
0.32
4.95
5.79
0.31
0.00
3.02
0.00
9.13
0.12
0.09
0.00
0.26
5.14
1.20
0
0.84
0
7.19
0.18
0.03
0
0.49
0
0.49
0.40
0.97
0.70
0.27
0.21
0.02
0.02
0.11
0.26
0.02
0.01
0.96
0
0.01
3.99
16.00
±
38.97
31.44
3.41
0.08
9.32
0.22
0.86
0.02
n.a.
<D.L.
2.83
0.04
<D.L.
0.11
0
0.05
87.24
0.04
5.04
1.30
2.00
1.65
0.15
0.66
0.12
1.45
0.04
0.09
0.02
0.08
0.03
1.61
0.05
0.07
0.01
4.80
0.37
0
1.80
0.09
7.06
0
0.71
0
0.71
0
0.05
3.95
±
70.00
20.03
0.03
<D.L.
<D.L.
<D.L.
<D.L.
<D.L.
<D.L.
0.05
11.60
0.09
<D.L.
<D.L.
0
0
101.85
1.02
0.28
0.05
0.03
0.28
0.04
1.33
8
0.13
5.04
0.10
0.09
6
8
20
4.95
0.03
5
20
20
0.08
0.26
0.08
15.92
4
12
28
5.18
2.82
8.00
Interlayer site
Ca
Na
K
Sum
Anions
Cl
F
OH
±
3.00
1.01
4.01
0.01
0.01
0.00
0.96
0.00
0.96
0.00
0.02
0.00
0.29
0.18
0.12
0.15
0.02
0.03
Notes: n = number of analyses; 1 = C0 (Mg chlorite, OTS07-004); 2 = C1 (Fe chlorite, OTS07-002); 3 = C2 (sudoite, OTS07-004); 4 = tourmaline (Camie
River; OTS07-002); 5 = tourmaline (Matoush; AM06, 21); 6 = albite cement (OM09, 47, OTS07-002); atomic proportions calculated on an anhydrous basis;
n.a. = not analyzed; <D.L. = below detection limit; OH calculated by subtraction
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BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
TABLE 3. Electron Microprobe Analyses of Carbonate Minerals
at Camie River
Fe
C1 chlorite
(proximal to ore)
(n=12)
C1 chlorite
(n=11)
C1 chlorite
(fault zone)
(n=13)
typical
chamosite
C0 chlorite
(n=8)
C2 chlorite
(n=12)
Athabasca
pre-ore
Athabasca
post-ore
Al
typical sudoite
Mg
typical clinochlore
FIG. 9. Ternary diagram showing the chemical variation of chlorites at
Camie River, in reference to the typical compositions of chamosite and
clinochlore reported by Deer et al. (1992), and sudoite reported by Lin and
Bailey (1985). Additionally, the fields for Athabasca pre- and post-ore chlorites (Cloutier et al., 2009; Alexandre et al., 2005) are shown.
with average K2O and CaO contents below 1.0 wt % (Table
2). Paragenetically later Ca1 calcite displays nearly stoichiometric Ca contents, with minor replacement by Fe (Table 3).
M1 muscovite in conglomerate and sandstone at Camie
River and Matoush shows considerable variability in the degree of octahedral-coordinated Al (AlVI) substitution, interlayer-cation substitution, and Si contents (Table 4; Fig. 11). At
Camie River, M1 in the Indictor Formation has low SiO2 contents that average 48.5 wt % and displays an average AlVI substitution factor of 0.90, meaning one of every 10 AlVI atoms is
replaced by Fe + Mg + Mn + Cr + Ti. As much as 1.0 wt %
Na2O substitutes for K2O in some M1 samples from Camie
River, indicative of paragonite. The same samples also have
the lowest SiO2 contents (46.0 wt %) and the highest AlVI substitution factor of 0.94. In contrast, M1 from Matoush has
50
Legend
illite, Indicator Fmn. fault zone
C1 chlorite, Indicator Fmn.
40
number
M1 muscovite Indicator Fmn.,
phengitic
M1 muscovite, Indicator Fmn.
30
M1 muscovite, Indicator Fmn.,
paragonitic
M1 muscovite, basement
rocks
20
C0 chlorite, basement
rocks
10
0
125
150
175
200
225
250
temperature (ºC)
275
300
325
FIG. 10. Histogram of estimated formation temperatures of white mica
and chlorite in the western Otish Basin. Estimated temperatures were calculated using the geothermometers of Cathelineau (1988) and Zang and Fyfe
(1995); see text for discussion.
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Sample no.
1
n
Oxide (wt %)
FeO
MnO
MgO
BaO
CaO
Total
2
Atomic proportions
number of
oxygens*
Fe
Mn
Mg
Ba
Ca
Sum
±
2
±
4
3
±
2
3.43
2.26
18.37
0.04
29.75
53.85
1.85
3.03
2.41
0.01
0.05
2.54
0.41
0.30
0.14
<D.L.
54.27
55.17
0.29
0.48
0.04
−
4.83
5.35
53.26
0.35
0.37
0.04
5.28
59.28
0.93
0.04
0.06
0.05
0.31
0.57
2
0.09
0.06
0.86
−
1.00
2.02
0.05
0.08
0.12
−
0.01
−
1
0.01
0.00
0.00
−
0.99
1.00
0.00
0.01
0.00
−
0.01
−
1
0.87
0.01
0.01
−
0.11
1.00
0.01
0.00
0.00
−
0.01
−
Notes: n = number of analyses; 1 = dolomite (OTS07-004); 2 = calcite
(Ca1; OTS07-002, -004); 3 = siderite (OTS07-004); *atomic proportions calculated assuming ideal CO2 contents; <D.L. = below detection limit; − =
value not able to be calculated
higher SiO2 contents, averaging 50.5 wt %, that is accompanied by a low average AlVI substitution factor of 0.80. This relationship is consistent with phengitic substitution ([R+2]VI +
SiIV D AlVI + AlIV; e.g., Guidotti et al., 1989). Basement-rock−
hosted M1 from Camie River shares a similar composition to
sandstone-hosted M1 at Camie River, with SiO2 contents averaging 47.4 wt %, and a slightly lower average AlVI substitution factor of 0.87 (Table 4; Fig. 11). The estimated temperature of formation for the majority of basement-rock− and
sandstone-hosted M1 at Camie River is between 250° and
275°C (Fig. 10). Temperatures between 275° and 300°C are
estimated for some basement-rock− and sandstone-hosted
M1 at Camie River, in addition to a high proportion of phengitic M1 sampled in the Indicator Formation at Matoush
(Fig. 10).
Late diagenetic minerals: Well-preserved C1 chlorite from
nonfractured or faulted conglomerate of the Indicator Formation contains lower MgO contents and higher FeO contents relative to C0 chlorite, the former averaging 9.5 and
30.2 wt %, respectively (Table 2). The composition of wellpreserved C1 is close to that of chamosite (Deer et al., 1992;
Fig. 9) and is estimated to have formed at a temperature
around 300°C (Fig. 10). In faulted conglomerate of the Indicator Formation, C1 displays widely varying FeO and Al2O3
contents (Fig. 9) and up to 1.0 wt % combined K2O + CaO,
suggesting small-scale intergrowth with late alteration illite
and montmorillonite. C1 in highly porous, desilicified sandstone less than 10 m from U mineralization contains the highest proportion of FeO relative to all Camie River chlorite
(Fig. 9) and has similar K2O and CaO contents to fault-hosted
C1.
Tourmaline from the Indicator Formation exposed in outcrop has an Mg-rich composition similar to that of dravite
(Table 2; Fig. 12). Much of the tourmaline from the Indicator
Formation at Matoush has compositions similar to dravite;
411
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BEYER ET AL.
TABLE 4. Electron Microprobe Analyses of Muscovite and Illite at Camie River
Sample no.
1
n
Oxide (wt %)
SiO2
Al2O3
FeO
MnO
MgO
TiO2
Cr2O3
V2O3
BaO
CaO
Na2O
K2O
Cl
F
O Cl
OF
Total
9
47.44
31.92
1.41
0.03
1.55
0.42
<D.L.
0.09
0.19
<D.L.
0.21
10.49
<D.L.
0.07
0
0.03
93.83
±
2
±
3
22
0.41
0.74
0.36
0.02
0.13
0.18
−
0.04
0.12
−
0.04
0.49
−
0.02
−
0.01
0.60
48.46
33.33
1.33
0.02
1.44
0.25
0.04
0.09
0.41
<D.L.
0.13
10.37
<D.L.
0.07
0
0.03
95.74
±
13
1.22
1.84
0.58
0.02
0.68
0.09
0.04
0.09
0.14
−
0.06
0.53
−
0.06
−
0.02
1.56
46.05
36.03
1.76
0.02
0.22
0.10
0.02
0.02
0.06
0.02
0.79
9.92
<D.L.
<D.L.
0
0
95.02
4
±
25
0.45
0.88
0.50
0.01
0.06
0.04
0.02
0.02
0.04
0.02
0.17
0.31
−
−
−
−
0.86
50.46
29.77
2.45
0.02
2.34
0.39
0.36
0.10
0.07
<D.L.
0.11
11.00
<D.L.
0.09
0
0.04
97.10
5
±
7
1.37
1.54
0.65
0.02
0.49
0.16
0.18
0.10
0.10
−
0.04
0.43
−
0.02
−
0.01
1.88
52.39
31.96
2.08
<D.L.
2.68
0.39
<D.L.
0.05
0.41
0.03
0.10
8.17
<D.L.
0.17
0
0.07
98.37
0.71
1.12
0.16
−
0.24
0.10
−
0.04
0.10
0.01
0.04
0.81
−
0.04
−
0.01
0.80
Atomic proportions
Number of oxygens
Tetrahedral sites
Si
AlIV
Sum
6.42
1.58
8.00
0.04
0.04
−
6.40
1.60
8.00
0.12
0.12
−
6.14
1.86
8.00
0.06
0.06
−
6.63
1.37
8.00
0.13
0.13
−
6.65
1.35
8.00
0.07
0.07
−
Octahedral sites
AlVI
Fe
Mn
Mg
Cr
Sum
3.51
0.16
−
0.31
−
3.98
0.06
0.04
−
0.03
−
−
3.60
0.15
−
0.28
−
4.03
0.13
0.07
−
0.13
−
−
3.81
0.20
−
0.04
−
4.05
0.05
0.06
−
0.01
−
−
3.24
0.27
−
0.46
0.04
4.01
0.10
0.07
−
0.10
0.02
−
3.42
0.22
0.00
0.51
0.00
4.15
0.08
0.02
0.00
0.05
0.00
−
Interlayer site
Ca
Na
K
Sum
−
0.06
1.81
1.87
−
0.01
0.09
−
−
0.03
1.75
1.78
−
0.02
0.07
−
−
0.20
1.69
1.89
−
0.04
0.05
−
−
0.03
1.84
1.87
−
0.01
0.08
−
0.03
1.32
1.35
0.01
0.14
−
Anions
Cl
F
OH
−
0.03
3.97
−
0.01
0.01
−
0.03
3.97
−
0.02
0.02
−
−
4.00
−
−
−
−
0.04
3.96
−
0.01
0.01
−
0.07
3.93
−
0.01
0.01
22
22
22
22
22
Notes: n = number of analyses; 1 = M1 muscovite, basement rocks (Camie River; OTS07-004); 2 = M1 muscovite, Indicator Formation (Camie River;
OM47, 49, OTS07-004) 3 = M1 muscovite, paragonitic, Indicator Formation (Camie River; OTS07-002); 4 = M1 muscovite, phengitic (Matoush; AM06, 15,
21); 5 = illite, fault-damaged Indicator Formation (Camie River; OTS07-004); atomic proportions calculated on an anhydrous basis; n.a. = not analyzed; <D.L.
= below detection limit; OH calculated by subtraction; − = value not able to be calculated
however, a trend toward higher Cr contents is observed. At
Camie River, tourmaline from the Indicator Formation is enriched in Fe and has a composition approaching that of schorl.
Postdiagenetic alteration minerals: Fracture-hosted C2
chlorite is rich in Al2O3, averaging 30.4 wt % (Table 2), with a
composition indistinguishable from that of sudoite (Lin and
Bailey, 1985; Fig. 9). Similar to fracture-hosted C1 chlorite,
C2 sudoite contains over 1.0 wt % combined K2O + CaO, indicating that C2 may contain small-scale intergrowth with late
alteration phyllosilicates. A temperature of formation for C2
based could not be reliably estimated due to the influence of
the possible intergrowths on calculations of tetrahedral site
occupancy. Siderite that occurs subsequent to U mineralization in the basement rocks displays replacement of one out of
0361-0128/98/000/000-00 $6.00
every 10 Fe atoms by Ca and minor replacement by Mg and
Mn (Table 3).
Illite after M1 muscovite in an Indicator Formation-hosted
fault zone at Camie River displays a notable interlayer site deficiency, with less than seven of every 10 interlayer K atoms
present (Table 4; Fig. 11). Low K2O content is accompanied
by a low average AlVI substitution factor of 0.82, which is typical for illite (Deer et al., 1992). The calculated temperatures
of illite formation vary widely between 130° and 220°C but
are lower relative to M1 (Fig. 10).
Stable isotope geochemistry
M1 muscovite from the Indicator Formation at Camie
River has δ18O values that range from 7.9 to 13.7‰ and δ2H
412
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
413
ideal
muscovite
LEGEND
Camie River
illite (ss, fault zone; n=7)
1.00
M1 muscovite (ss; n=22)
M1 muscovite (ss, paragonitic; n=13)
M1 muscovite (basement; n=9)
Matoush
0.90
M1 muscovite (ss; n=25)
phengite
0.85
0.95
0.80
0.90
2.0
substitution for AlVI
1.00
K+
Na
(ap
fu)
1.8
0.85
alteration to
phengite
1.6
illi
te
to
tio
n
alt
e
ra
Si (apfu)
1.2
1.0
6.8
6 .6
6.4
6.2
6 .0
ideal
illite
Si (apfu)
(a
pf
u)
1.4
K+
Na
0.80
AlVI
AlVI (apfu)
+ Mg + Fe + Mn + Cr + Ti (apfu)
0.95
FIG. 11. Relationship between contents of interlayer cations K + Na, Si, and the degree of substitution for octahedral Al
in white mica at Camie River and Matoush, in reference to typical compositions of muscovite and illite reported by Deer et
al. (1992), and phengite reported by Konopásek (1998). M1 muscovite shows an alteration trend towards phengite and an alteration trend towards illite. Vertical tie lines for some data points have been removed for clarity. apfu = atoms per formula
unit; ss = sandstone or conglomerate of the Indicator Formation.
Mg
Mg
Cr+Fe
tourmaline
(Matoush; n=20)
typical
schorl
Fe + 2
Cr+3
tourmaline
(Camie River; n=12)
Cr+Fe
typical
dravite
tourmaline
(outcrop; n=6)
Mg +
2
Al+3
Al
60/40
Al
FIG. 12. Ternary diagram showing the chemical variation of tourmaline at
Camie River and Matoush, in reference to the typical compositions of dravite
and schorl reported by Deer et al. (1992).
0361-0128/98/000/000-00 $6.00
values that range from −83 to −56‰ (Table 5). Basementrock−hosted M1 at Camie River has a similar δ18O value of
10.1‰, and a slightly higher δ2H value of −48‰ (Table 5). A
temperature of 275°C (Fig. 11) was used to calculate δ18Ofluid
values, which range between 3.9 and 9.7‰ and δ2Hfluid values
between −49 and −14‰ (Fig. 13). Phengitic M1 from the Indicator Formation at Matoush has slightly lower δ18O values
(7.5−10.9‰), and higher δ2H values (−66 to −45‰) relative
to M1 at Camie River (Table 5). A temperature of 300°C (Fig.
11) was used to calculate δ18Ofluid values between 4.1 and
7.5‰, and δ2Hfluid values between −32 and −11‰ (Fig. 13).
C1 chlorite from the Indicator Formation at Matoush has a
δ18O value of 12.1‰, and a δ2H value of −89‰ (Table 5). C1
from a fault zone within the Indicator Formation at Camie
River has an identical δ18O value of 12.1‰ but a lower δ2H
value of −102‰ (Table 5). Using the minimum estimated
temperature of formation of 300°C, the calculated δ18Ofluid
value is 12.0‰. Fluids in equilibrium with C1 at Matoush
had a δ2H value of −49‰, whereas fault zone-hosted C1 at
Camie River has a lower δ2Hfluid value of −68‰ (Fig. 13).
The δ18O and δ2H values calculated for the fluids in equilibrium with M1 and C1 form a mixing trend between two
components (Fig. 13). One component has a low δ18Ofluid
value around 3‰ and a high δ2Hfluid value near −5‰, similar
to the isotopic composition of seawater. The second component has much higher δ18Ofluid and lower δ2Hfluid values, near
413
414
BEYER ET AL.
TABLE 5. Oxygen and Hydrogen Stable Isotope Compositions of Phyllosilicates in the Western Otish Basin
Sample no.
Lithology
OTS07-01-605.3m
OM09-54.48m
OM47-33.46m
OM47-139.09m
OM49-47.76m
OM49-191.96m
OM49-229.66m
OM49-265.71m
OTS07-03-76.5m
OTS07-03-153.0m
OTS07-04-102.1m
AM06-41.4m
AM08-6.89m
AM20-480.38m
AM21-215.58m
AM15-260.33m
OTS07-04-147.0m
Comment
bs
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
Mineral
Fraction size
δ18Omin
δDmin
H2O (%)
Temp (°C)
δ18Ofluid
δDfluid
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
M1
C1
C1
m
<2 µm
<2 µm
<2 µm
<2 µm
<2 µm
<2 µm
<2 µm
m
<5 µm
<2 µm
<2 µm
2-5 µm
2-5 µm
<2 µm
<2 µm
<2 µm
10.1
7.9
13.7
9.0
10.6
9.9
10.2
8.6
8.3
9.4
8.5
7.7
7.5
8.7
10.9
12.1
12.1
–48
–62
–75
–58
–83
–61
–79
–72
–62
–56
–58
–46
–45
–54
–66
–89
–102
5.5
6.0
7.3
4.8
5.1
6.6
4.8
6.5
6.3
5.7
5.2
5.2
4.9
4.3
5.1
11.4
11.4
275
275
275
275
275
275
275
275
275
275
275
300
300
300
300
300
300
6.0
3.9
9.7
5.0
6.6
5.9
6.2
4.6
4.2
5.4
4.4
4.3
4.1
5.3
7.5
12.0
12.1
–14
–28
–41
–24
-49
–27
–45
–38
–28
–22
–24
–12
–11
–20
–32
–49
–68
Fracture
Fracture
Seq. bdy.
Fracture
Fault zone
Notes: Temperatures of white mica and chlorite were calculated using mineral chemistry and the methods of Cathelineau (1988) and Zang and Fyfe (1995)
and are presented in Figure 10; Seq. bdy. = depositional sequence boundary; bs = basement rocks; ss = Indicator Formation; m = sampled by microdrill; oxygen and hydrogen isotope ratios are in units of per mil relative to V-SMOW
+13 and −55‰, respectively. These values are consistent with
those of metamorphic fluids that attain high δ18Ofluid values
due to interaction with 18O-rich rocks at low fluid/rock ratios,
and low δ2Hfluid values due to a meteoric water component.
U-Pb and Pb-Pb isotope ratios of Camie River uraninite
U-Pb isotope analysis of U1a uraninite from Camie River
by LA-HR-MC-ICPMS indicates 207Pb/206Pb dates that range
MW
L
M
A
-25
influence of
modern
meteoric
water
2
δ HH
2O
V-SMOW
COMPONENT 1
seawater-influenced
basinal brine
V-SMOW
0
-50
-75
Temp.(˚C) Camie River Matoush
300
C1
C1
M1
275
M1
M1 (basement)
275
-10
-5
18
COMPONENT 2
metamorphismderived fluids
0
δ O H2O
δ18O
10
15
V-SMOW
40Ar/39Ar
FIG.13. Calculated
and
values of fluids in equilibrium with phyllosilicates at Camie River. The samples plot along a mixing line between seawater-influenced basinal brines and metamorphism-derived fluids. Samples
from fault and fracture zones, and sequence boundaries have preferentially
exchanged H isotopes with δ2H-depleted, modern meteoric water (cf. Table
5). Fields defining the range of δ18O and δ2H values of basinal fluids in equilibrium with muscovite/illite and sudoite in the Athabasca Basin (A), and
metamorphic fluids (M) reported by Taylor (1974) are plotted in gray for reference (MWL = meteoric water line; V-SMOW = Vienna Standard Mean
Ocean Water).
0361-0128/98/000/000-00 $6.00
δ2H
5
from 2042 to 1721 Ma (Table 6). A majority of the analyses
have low 206Pb/204Pb ratios <1,000 (Table 6), indicative of high
proportions of common Pb that is likely contributed from S0
and S2 sulfides (Fig. 7d, e). Uraninite with the highest proportion of common Pb (206Pb/204Pb = < 500) gives 207Pb/206Pb
dates that approach the age of basin formation, whereas
uraninite with the lowest proportion of common Pb
(206Pb/204Pb = 2,668) gives a 207Pb/206Pb date of 1721 ±20 Ma
(Table 6; Fig. 14a), which is indistinguishable from the 1723
±16 Ma U-Pb date reported for Camie River uraninite by
Höhndorf et al. (1987).
The majority of U-Pb isotope ratios in Camie River uraninite are negatively discordant (Table 6), indicative of Pb addition or loss of U subsequent to uraninite formation, and
yield a discordant U-Pb upper intercept date of 1692 ±32 Ma
(Fig. 14b). This date is within error of the youngest
207Pb/206Pb date of 1721 ±20 Ma. The regression line is poorly
constrained at the lower intercept, resulting in a date of −201
±380 Ma that most likely reflects recent modification of U-Pb
isotope ratios. The relatively high error margins for U-Pb isotope ratios reflect uraninite heterogeneity (Kotzer and Kyser,
1993; Polito et al., 2004, 2005; Chipley et al., 2007), typified
by the mottled texture observed in BSE-SEM images of
Camie River uraninite (Fig. 14c). Uraninite heterogeneity is
also manifested as domains of uraninite with differing proportions of common Pb that are commonly encountered during the analysis of a single ablation spot (Fig. 14d).
dating of muscovite
Muscovite (M1) from the basement rocks and the Indicator
Formation was dated using the 40Ar/39Ar method (Table 7).
One sample from intensely M1 altered metapsammite within
several meters of the unconformity gives a disturbed age
spectrum with a total fusion date of 1663 ±10 Ma (Fig. 15).
M1 sampled from within 2 m of the U mineralization zone in
the basement rocks has a bell-shaped age spectrum and a
total fusion date of 1528 ±10 Ma (Fig. 15). The disturbed age
414
0361-0128/98/000/000-00 $6.00
415
828
983
1134
1449
945
977
541
733
818
1305
785
584
1166
2013
1164
2240
1562
1633
653
1636
1677
1800
1487
1157
1132
2668
626
598
705
1286
920
206Pb/204Pb
1.76
1.43
1.24
0.98
1.49
1.44
2.60
1.92
1.72
1.10
1.80
2.42
1.24
0.76
1.25
0.65
0.97
0.96
2.15
0.95
0.90
0.72
1.08
1.43
1.27
0.53
2.47
2.38
2.15
1.15
1.66
0.348
0.315
0.312
0.348
0.394
0.325
0.385
0.350
0.395
0.358
0.359
0.386
0.486
0.440
0.399
0.390
0.398
0.413
0.435
0.372
0.379
0.389
0.357
0.373
0.334
0.337
0.343
0.265
0.294
0.263
0.263
206Pb/238U
0.019
0.016
0.021
0.010
0.017
0.019
0.019
0.020
0.053
0.014
0.016
0.029
0.020
0.022
0.033
0.024
0.030
0.032
0.042
0.040
0.026
0.030
0.023
0.033
0.019
0.020
0.040
0.006
0.028
0.004
0.019
±1σ
4.952
4.458
4.440
4.953
5.632
4.628
5.598
4.997
5.651
5.099
5.143
5.554
6.845
6.169
5.588
5.464
5.614
5.821
6.204
5.229
5.340
5.485
4.968
5.226
4.680
4.699
4.842
3.801
4.327
3.871
3.869
207Pb/235U
0.390
0.268
0.372
0.173
0.296
0.377
0.512
0.392
0.797
0.291
0.399
0.638
0.413
0.378
0.605
0.387
0.579
0.586
0.940
0.730
0.361
0.689
0.477
0.727
0.376
0.282
0.759
0.174
0.657
0.170
0.443
±1σ
0.87
0.93
0.96
0.91
0.91
0.96
0.84
0.87
0.93
0.91
0.82
0.84
0.89
0.90
0.95
0.91
0.93
0.90
0.87
0.93
0.91
0.88
0.86
0.89
0.73
0.90
0.82
0.87
0.94
0.85
0.90
R
0.117
0.114
0.113
0.111
0.116
0.114
0.126
0.119
0.118
0.112
0.118
0.123
0.112
0.108
0.111
0.107
0.110
0.110
0.120
0.109
0.109
0.108
0.109
0.113
0.112
0.105
0.119
0.121
0.121
0.113
0.117
207Pb/206Pb
0.004
0.001
0.002
0.001
0.001
0.002
0.003
0.004
0.004
0.002
0.004
0.005
0.002
0.002
0.002
0.002
0.003
0.004
0.005
0.003
0.002
0.001
0.004
0.006
0.004
0.001
0.005
0.002
0.006
0.003
0.005
±1σ
1925
1765
1751
1924
2139
1816
2102
1934
2147
1974
1979
2105
2552
2353
2164
2124
2162
2228
2330
2039
2074
2120
1968
2042
1856
1873
1903
1517
1664
1508
1507
206Pb/238U
91
79
102
47
81
94
87
97
245
65
78
137
85
98
150
111
138
146
189
188
121
137
108
156
92
97
194
31
138
22
97
±1σ
1811
1723
1720
1811
1921
1754
1916
1819
1924
1836
1843
1909
2092
2000
1914
1895
1918
1950
2005
1857
1875
1898
1814
1857
1764
1767
1792
1593
1698
1608
1607
207Pb/235U
67
50
70
29
45
68
79
67
122
48
66
99
54
54
94
61
89
87
133
120
58
108
81
119
67
50
133
37
126
36
93
±1σ
1913
1864
1851
1816
1889
1872
2042
1941
1924
1826
1929
2003
1832
1758
1823
1744
1797
1800
1958
1783
1785
1759
1785
1848
1827
1721
1943
1965
1969
1852
1912
207Pb/206Pb
Apparent ages (Ma)
58
23
30
11
17
30
48
57
62
37
65
70
33
41
41
28
47
61
75
50
41
17
63
95
64
20
75
36
85
46
70
±1σ
0
6
6
–5
–12
4
–3
1
–11
–8
–2
–4
–38
–33
–18
–22
–14
–18
–17
–13
–16
–20
–7
–10
–1
–8
3
24
19
19
22
disc. (%)
Notes: Uraninite sampled from the interval 174.3 to 174.5 m in drill hole OTS07-004; com. Pb = percent 206Pb from common Pb; R = error correlation coefficient; disc. = percent discordance
Camie River 1
Camie River 2
Camie River 3
Camie River 4
Camie River 5
Camie River 6
Camie River 7
Camie River 8
Camie River 9
Camie River 10
Camie River 11
Camie River 12
Camie River 13
Camie River 14
Camie River 15
Camie River 16
Camie River 17
Camie River 18
Camie River 19
Camie River 20
Camie River 21
Camie River 22
Camie River 23
Camie River 24
Camie River 25
Camie River 26
Camie River 27
Camie River 28
Camie River 29
Camie River 30
Camie River 31
Sample no.
com.
Pb (%)
TABLE 6. Isotopic Data and Apparent Ages for Camie River Uraninite
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
415
416
BEYER ET AL.
2,100
a
207
b
0.5
2,000
2200
0.4
Pb/238U
1,950
1,900
1,850
206
Pb / 206Pb date (Ma)
2,050
1800
0.3
1400
0.2
1000
1,800
0.1 600
1,750
1,700
500
1,000
1,500
2,000
2,500
0.0
3,000
0
intercepts at
-201 ±380 & 1692 ±32 Ma
MSWD = 0.089
2
4
206
Pb / 204Pb
6
8
207
Pb/235U
approx. depth of ablation pit (µm)
c
0
d
5
10
15
20
25
30
0
500
1,000
1,500
2,000
2,500
3,000
206
Pb/204Pb
FIG. 14. a. Plot of calculated 207Pb/206Pb dates vs. 206Pb/204Pb from in situ isotopic analysis by LA-HR-MC-ICPMS of
Camie River uraninite (Table 6). The analysis with the highest proportion of radiogenic Pb yields a date of 1721 ±20 Ma. PbPb and U-Pb dates were calculated using the equations of Ludwig (2000). b. U-Pb concordia diagram from in situ isotopic
analysis by LA-HR-MC-ICPMS of Camie River uraninite. c. Uraninite (U1a) from Camie River, showing mottled texture indicative of alteration during fluid events that occurred subsequent to original crystallization (BSE-SEM image). d. Plot of
206Pb/204Pb vs. depth of a laser ablation pit in a polished thin section of Camie River uraninite. Laser ablation sampled at least
three domains (separated by dashed lines) of uraninite that are marked with differing 206Pb/204Pb values (analysis “Camie
River 20”; Table 6).
spectra for both of these samples could result from 40Ar* loss
during a thermal or fluid event subsequent to M1 crystallization. Mixed mineral phases resulting from new mineral
growth can produce bell-shaped age spectra (McDougall and
Harrison, 1988) consistent with phengitic substitution recorded
in basement-hosted M1 at Camie River (Fig. 11). The total
fusion date for each basement-hosted M1 sample is likely a
homogenization of an older age of formation and a younger
thermal or fluid event that subsequently affected the M1
muscovite.
M1 from a fracture in the Indicator Formation has a relatively undisturbed age spectrum with a plateau age of 1409
±29 Ma (Fig. 15). This date is unlikely to represent the age of
M1 crystallization, based on the older total fusion dates for
coeval basement-hosted M1, and instead reflects the time of
Ar closure in muscovite following a thermal or fluid event that
completely reset the Ar isotope systematics in M1. This date
probably represents a minimum age for the sealing of fracture-related fluid-flow pathways in the Indicator Formation
0361-0128/98/000/000-00 $6.00
and coincides with a period of widespread anorogenic magmatism across Laurentia and Baltica (Van Schmus et al., 1975;
Medaris et al., 2003; Anderson and Morrison, 2005).
U-Pb isotope and trace element geochemistry
of 2% HNO3 leachates
The 2% HNO3 leach method has been employed in the
Athabasca, Thelon, and Kombolgie Basins (Holk et al., 2003;
Alexandre et al., 2009a; Cloutier et al., 2009, 2010; Beyer et
al., 2010) to detect excess radiogenic Pb and pathfinder elements that dispersed from a potential U deposit, from which
vectors to ore can be determined by following prospective
trends. 207Pb/206Pb ratios less than 0.6 are termed radiogenic
because they reflect an increasing proportion of radiogenic
Pb that was possibly sourced from a Proterozoic U deposit.
Excess radiogenic Pb that cannot be accounted for by the
decay of leachable U contained in a sample is referred to as
unsupported radiogenic Pb, whereas radiogenic Pb that can
be accounted for by the decay of leachable U contained in a
416
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
TABLE 7.
40Ar/39Ar
417
Analytical Data for M1 Muscovite at Camie River
40Ar*/39Ar
K
Apparent
age (Ma)
± 2σ
4.8
17.1
28.0
14.1
12.3
6.9
7.0
5.0
4.9
41.69
50.29
55.64
58.10
57.37
53.02
56.12
59.42
25.85
1136
1302
1398
1441
1428
1352
1407
1463
783
284
64
38
71
74
142
147
221
218
6.4
16.5
19.6
31.3
13.1
1.8
0.9
1.8
8.6
65.79
65.91
70.68
75.83
81.21
82.34
77.88
69.41
60.79
1568
1569
1644
1721
1798
1814
1750
1624
1486
63
25
21
13
14
105
220
108
24
OTS07-004-173.5m (basement rocks, 1 m from U mineralization), total fusion date = 1528 ±10 Ma
1
P
0.0000
0.0150
0.54
100.0
2.3
2
P
0.0001
0.0163
0.02
97.6
7.0
3
P
0.0000
0.0129
0.08
98.6
10.7
4
P
0.0001
0.0129
0.13
98.3
12.1
5
P
0.0001
0.0136
0.08
97.8
10.4
6
P
0.0001
0.0137
0.03
96.3
5.2
7
P
0.0001
0.0138
0.43
95.6
3.2
8
P
0.0002
0.0154
0.18
93.8
9.4
9
P
0.0003
0.0173
0.15
92.1
39.6
66.70
59.83
76.27
76.38
72.12
70.09
69.56
60.88
53.16
1582
1470
1727
1729
1666
1635
1627
1488
1354
145
49
32
23
32
61
90
37
9
Step
36Ar/40Ar
39Ar/40Ar
Ca/K
40Ar
OTS07-004-102.1m (fractured Indicator Fmn.), plateau age = 1409 ±29 Ma
1
0.0003
0.0215
0.22
2
0.0001
0.0193
3
P
0.0001
0.0177
0.08
4
P
0.0000
0.0171
0.03
5
P
0.0000
0.0174
0.05
6
P
0.0002
0.0177
0.05
7
P
0.0002
0.0167
0.05
8
P
0.0002
0.0160
0.05
9
0.0009
0.0286
0.16
air (%)
89.8
97.2
98.3
99.3
99.7
93.7
93.9
94.9
73.9
OTS07-001-605.3m (basement rocks, near unconformity), total fusion date = 1663 ±10 Ma
1
P
0.0000
0.0150
0.00
98.6
2
P
0.0000
0.0151
0.01
99.5
3
P
0.0000
0.0141
0.02
99.8
4
P
0.0000
0.0132
0.00
99.8
5
P
0.0000
0.0123
0.00
100.0
6
P
0.0000
0.0120
0.00
98.8
7
P
0.0001
0.0124
0.00
96.6
P
0.0001
0.0141
0.03
97.8
8
9
P
0.0001
0.0159
0.03
96.5
39Ar
(%)
Notes: P = step used in calculation of total fusion/weighted plateau apparent age
sample is referred to as supported radiogenic Pb (Holk et al.,
2003).
Radiogenic Pb was leached from samples in drill core that
penetrate Camie River U mineralization and from samples
~13 km west of the Camie River prospect along strike of the
electromagnetic conductor trend (Table 8; Fig. 2). In the Indicator Formation, radiogenic Pb is restricted to zones near
depositional sequence boundaries, especially the basal unconformity, and zones that have been fault damaged or fractured
(Table 8; Fig. 16). Basement rocks above U mineralization
2600
= 1409 ±29 Ma
= 1528 ±10 Ma
(basement) total fusion = 1663 ±10 Ma
(Indicator Fmn.) plateau
apparent age (Ma)
M1
2400 muscovite
(basement) total fusion
2200
2000
1800
1600
and near the unconformity contain radiogenic Pb but become
nonradiogenic away from these zones. Unsupported radiogenic Pb is present in about one-third of the samples (Table
8; Fig. 17). The excess radiogenic Pb in these samples was
likely added by fluids that interacted with Camie River U
mineralization and carried the radiogenic Pb away from the
deposit. Supported Pb in the remaining samples was likely
produced in situ by detrital U-bearing monazite and zircon,
which are common in the Indicator Formation.
Camie River U mineralization is associated with elevated
leachable Mo, W, and Nb contents, which are also detectable
in leachates from fractures or fault zones in the Indicator Formation above the mineralization (Table 8; Fig. 16). The elements Co, Ni, and As, which can be associated with complextype unconformity-related U mineralization, are elevated
only in samples that contain S1 pyrite in the Indicator Formation, and S0 sulfides in the basement rocks. Cu and Zn are
elevated in highly altered fault gouge 7 m above U mineralization (Table 8).
1400
Discussion
1200
Evolution of the western Otish Basin
Early to peak diagenesis: Two key events relating to U mineralization in the Otish Basin, specifically, the establishment of
diagenetic aquifers and the interaction of basinal brines with
basin fill in the diagenetic aquifers, resulted from basin-scale
1000
0
10
20
30
40
50
60
70
80
90
100
Cumulative 39Ar Released (%)
FIG. 15. The 40Ar/39Ar apparent age spectra for M1 muscovite sampled
from the Indicator Formation and basement rocks.
0361-0128/98/000/000-00 $6.00
417
0361-0128/98/000/000-00 $6.00
418
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
bs
bs
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
bs
bs
ss
ss
ss
ss
ss
bs
bs
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
ss
bs
bs
bs*
bs
bs
bs
OM09-17.7m
OM09-31.9m
OM09-44.6m
OM09-54.5m
OM09-74.7m
OM09-80.6m
OM09-93.8m
OM09-101.8m
OM09-106.9m
OM09-111.2m
OM09-114.1m
OM09-118.6m
OM09-125.7m
OM47-24.8m
OM47-33.5m
OM47-59.5m
OM47-79.8m
OM47-93.9m
OM47-113.7m
OM47-139.1m
OM47-145.8m
OM47-154.1m
OM47-163.6m†
OM47-167.3m
OM47-181.1m
OTS07-003-52.0m
OTS07-003-78.1m
OTS07-003-153.0m
OTS07-003-171.5m
OTS07-003-174.1m
OTS07-003-174.4m†
OTS07-003-188.7m
OT08-1-1
OT08-1-2
OT08-1-3
OTS07-004-61.4m
OTS07-004-62.8m
OTS07-004-87.5m†
OTS07-004-102.1m
OTS07-004-126.0m†
OTS07-004-135.9m
OTS07-004-147.0m†
OTS07-004-154.0m†
OTS07-004-163.0m†
OTS07-004-167.0m†
OTS07-004-173.5m†
OTS07-004-174.5m†
OTS07-004-178.0m
OTS07-004-182.9m
OTS07-004-215.5m
23.6
28.8
64.8
32.5
24.6
25.1
26.3
23.4
38.2
31.8
27.2
28.5
20.6
22.8
22.4
23.8
23.5
37.5
37.0
29.5
29.8
27.5
228.9
137.6
18.7
28.9
23.2
32.7
32.0
80.1
84.2
24.8
28.2
18.9
14.6
32.9
28.2
54.9
59.8
68.8
83.1
59.3
81.6
66.2
196.2
99.4
50.5
38.5
35.3
25.9
204Pb
206Pb/
16.3
16.8
19.6
17.0
16.8
16.5
16.8
16.5
17.2
17.2
16.9
17.0
16.1
16.1
16.3
16.4
16.3
17.5
17.5
17.0
17.0
16.8
35.3
25.3
15.5
18.1
15.3
18.3
16.8
22.4
22.0
16.3
22.4
14.8
12.5
18.8
15.4
20.2
19.5
23.2
21.4
19.0
21.4
21.3
30.7
24.6
21.0
17.6
17.5
16.4
204Pb
207Pb/
0.69
0.58
0.30
0.53
0.68
0.66
0.64
0.71
0.45
0.54
0.62
0.60
0.78
0.71
0.73
0.69
0.69
0.47
0.47
0.58
0.57
0.61
0.15
0.18
0.83
0.63
0.67
0.56
0.53
0.28
0.26
0.66
0.80
0.79
0.85
0.58
0.55
0.37
0.33
0.34
0.26
0.32
0.26
0.32
0.16
0.25
0.42
0.46
0.50
0.63
206Pb
207Pb/
238U/
1.74
3.84
9.04
0.59
2.65
5.25
2.77
1.05
5.69
0.69
2.47
1.78
0.07
1.61
2.74
1.16
0.63
5.41
1.23
0.52
2.92
0.45
2.81
6.64
0.02
0.99
1.42
0.33
1.30
15.54
2.46
0.06
0.20
0.11
0.20
32.97
43.29
1.83
4.05
2.52
15.79
1.35
0.99
0.27
0.55
1.05
1.52
2.56
0.03
0.17
206Pb
<DL
<DL
78313
<DL
<DL
<DL
<DL
<DL
<DL
<DL
63682
<DL
<DL
<DL
<DL
<DL
13023
45551
<DL
<DL
13184
21145
<DL
31526
328602
<DL
<DL
<DL
<DL
50510
<DL
<DL
<DL
69926
85433
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
602469
<DL
125890
84470
P
(ppb)
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
1217
<DL
1271
1773
583
3500
<DL
<DL
1088
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
<DL
<DL
<DL
<DL
<DL
<DL
846
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
508
<DL
906
<DL
<DL
751
V
Cr
(ppb) (ppb)
<DL
<DL
690
<DL
<DL
855
<DL
<DL
475
<DL
<DL
<DL
2104
<DL
675
<DL
<DL
<DL
<DL
<DL
<DL
<DL
1504
1777
1553
<DL
<DL
<DL
<DL
439
<DL
<DL
<DL
<DL
<DL
570
<DL
<DL
<DL
<DL
430
4650
<DL
884
<DL
<DL
<DL
556
<DL
<DL
Co
(ppb)
<DL
<DL
534
<DL
<DL
<DL
<DL
<DL
1020
<DL
<DL
720
4605
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
2139
6903
3123
<DL
<DL
<DL
<DL
401
<DL
1137
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
1920
<DL
1003
<DL
711
940
4146
<DL
583
Ni
(ppb)
Zn
(ppb)
2081 <DL
<DL
<DL
<DL
1396
<DL
<DL
<DL
<DL
1557
764
<DL
<DL
1510
575
3637 2020
1642
548
<DL
<DL
<DL
864
<DL
1054
2178 <DL
2805 <DL
2742
571
<DL
<DL
1947 <DL
4239
585
1721
462
<DL
<DL
2470 <DL
3306 2026
1824 3525
3038 1391
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
3801 <DL
879 <DL
<DL
424
<DL
<DL
<DL
<DL
<DL
<DL
1644 <DL
2577 <DL
707 <DL
744 <DL
<DL
<DL
535 <DL
3656 1402
860 <DL
3816 5380
131620 15554
669
502
<DL
971
877
466
<DL
<DL
582
616
Cu
(ppb)
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
801
<DL
<DL
1398
3986
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
1008
3377
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
533
4405
<DL
<DL
As
(ppb)
<DL
470
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
1963
5296
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
433
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
106
<DL
248
<DL
<DL
113
<DL
637
<DL
<DL
<DL
Zr
(ppb)
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
114
<DL
996
<DL
<DL
<DL
Nb
(ppb)
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
<DL
<DL
<DL
<DL
<DL
104
<DL
<DL
<DL
105
<DL
<DL
<DL
<DL
<DL
<DL
101
<DL
<DL
<DL
275
13147
151
<DL
<DL
Mo
(ppb)
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
<DL
<DL
<DL
<DL
898
<DL
<DL
<DL
<DL
<DL
<DL
<DL
867
<DL
427
984
507
614
688
7227
<DL
4997
541
<DL
854
Nd
(ppb)
Pb
(ppb)
n.a.
<DL
n.a.
<DL
n.a.
3025
n.a.
<DL
n.a.
<DL
n.a.
980
n.a.
<DL
n.a.
<DL
n.a.
2188
n.a.
1317
n.a.
<DL
n.a.
<DL
n.a.
<DL
n.a.
<DL
n.a.
<DL
n.a.
<DL
n.a.
<DL
n.a.
1927
n.a.
2021
n.a.
1463
n.a.
<DL
n.a.
1484
n.a.
3846
n.a.
4985
n.a. 196390
<DL <DL
<DL <DL
<DL
1113
<DL
510
<DL
198
152
108
<DL
380
<DL <DL
<DL <DL
146 <DL
<DL <DL
<DL <DL
141
958
<DL
236
<DL <DL
<DL
362
135
849
<DL
798
<DL
8395
<DL 66058
<DL 11287
1263 382708
158
408
<DL
643
<DL
106
W
(ppb)
U
(ppb)
608
532
619
711
925
20742
<DL
<DL
<DL
622
<DL
1979
681
708
<DL
<DL
<DL
7981
1257
<DL
771
593
1372
885
<DL
<DL
1115
<DL
<DL
<DL
<DL
<DL
665
<DL
1226
5597
1153
1168
1157
<DL
<DL
965
643
<DL
<DL
37926
913
75448
595
1293
<DL
<DL
<DL
<DL
<DL
<DL
<DL
<DL
2383
2685
<DL
626
<DL
<DL
169
<DL
<DL
<DL
126
<DL
<DL
1045
<DL
861
982
1906
<DL
1456
114
211
1029
6730
293
1666
664
1742
614
5737
127
214216
<DL
47840
1694 1025413
<DL
1417
<DL
<DL
<DL
<DL
Th
(ppb)
Notes: Lith. = lithology (ss = Indicator Formation; bs = basement rocks); n.a. = not analyzed; <DL = below analytical detection limit, which is 1 ppb or less for all elements; * = U-mineralized; † =
unsupported radiogenic Pb
Lith.
Sample no.
TABLE 8. U and Pb Isotope Ratios and Selected Elemental Concentrations of 2% HNO3 Leachates from Camie River
418
BEYER ET AL.
BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
mod.
rad.
weakly
radiogenic
a (m)
0
highly
radiogenic
c (m)
0 glacial till
b(m)
0
glacial till
419
glacial till
50
50
50
100
100
100
150
150
0.9
0.6 0.3
207
Pb/206Pb
U min.
Mo ~13,000 ppb -->
200
0.9
207
0.6 0.3
0.9
d (m)
0
ppb
0
Pb/206Pb
1500 0
ppb
1500
LEGEND
207
Nb
Mo
glacial till
0.6 0.3
Pb/206Pb
W
Lithofacies 2
fault/fracture zone
Lithofacies 1b
50
Lithofacies 1a
unconformity
tightly-folded,
massive-sulfidebearing graphitic
schist
100
foliated qtz. +
chlorite (C0)
metapsammite
150
0.9
0.6 0.3
Pb/206Pb
ppb
0
1500
207
Mo
W
FIG. 16. a. Stratigraphic section of drill hole OTS07-004, accompanied by plots of 207Pb/206Pb ratios, and contents of Mo,
Nb, and W as determined by HR-ICP-MS isotopic and elemental analysis of 2% HNO3 leachates. This drill hole intersects
U mineralization in the basement rocks. b. Stratigraphic section of drill hole OM47, accompanied by a plot of 207Pb/206Pb ratios. c. Stratigraphic section of drill hole OM09, accompanied by a plot of 207Pb/206Pb ratios. d. Stratigraphic section of drill
hole OTS07-003, accompanied by plots of 207Pb/206Pb ratios and contents of Mo and W. Refer to Figures 2 and 3 for the locations of drill holes.
diagenetic processes in the Indicator Formation based on
mineral paragenesis and chemistry, and stable isotope geochemistry. Early quartz cements, such as those similar to Q1
1750 Ma
300
200
750 Ma
1250 Ma
206Pb/204Pb
250
zone of excess (unsupported) Pb
Camie River (n=47)
typical error
cement, in siliciclastic basins can result in significant reduction of porosity and permeability (Leder and Park, 1986;
Bjørlykke and Egeberg, 1993; Oelkers et al., 1996) and the
development of confining diagenetic aquitards (Hiatt et al.,
2003; Polito et al., 2006). However, extensively resorbed and
poorly preserved Q1 cement in the Indicator Formation suggest that water/rock interaction occurred during subsequent
peak diagenesis, and therefore the fluid-conducting characteristics of the Indicator Formation were little affected by early
150
100
a
250 M
50
common Pb
0
0
10
20
30
238U/206Pb
0361-0128/98/000/000-00 $6.00
40
FIG. 17. Plot of 238U/206Pb vs. 206Pb/204Pb from HR-ICP-MS isotopic
analysis of 2% HNO3 leachates of samples from Camie River. The evolution
of isotopic ratios from 250 to 1750 Ma (Holk et al., 2003) is shown. Radiogenic Pb in one-third of the samples is unsupported by leachable U (indicative of Pb produced from a high 206Pb source such as a uranium deposit)
and plots within the zone of unsupported Pb (hatched area). The remaining
samples have radiogenic Pb that is supported by leachable U, indicative of radiogenic Pb produced in situ.
419
420
BEYER ET AL.
quartz cementation. Paragenetically later albite cement, however, pervasively occluded secondary porosity in at least the
oldest two depositional sequences in the southeastern study
area, converting these strata to diagenetic aquitards. The distribution of albite cement is sharply reduced to the north,
where much of the Indicator Formation continued to conduct
fluids that precipitated paragenetically later M1 muscovite
and minor K1 kaolinite. Paragonitic M1 may have acquired Na
through the replacement of albite cement by M1 based on the
close spatial association between the two minerals (Fig. 6a).
The widespread distribution of M1 muscovite indicates that
diagenetic aquifers had been established in nearly all of the
Indicator Formation in the study area except albite-cemented
strata to the south (Fig. 18a). Diagenetic aquifers were conducting fluids that are isotopically similar to oxidizing basinal
brines in other Proterozoic to Phanerozoic basins (Wilson and
Kyser, 1987; Ayalon and Longstaffe, 1988; Bray et al., 1988;
Kotzer and Kyser, 1995; Fayek and Kyser, 1997; Polito et al.,
2004, 2005, 2006; Kyser and Cuney, 2008a) based on the O
and H isotope composition of end-member fluids associated
with M1 along the observed mixing trend (Fig. 13). At temperatures around 250°C, the estimated temperature of M1
formation (Fig. 10), basinal brines were within the metal
leaching window (Southgate et al., 2006; Kyser, 2007) and
were capable of leaching and transporting U. These basinal
brines probably also interacted with common detrital phosphates and rutile in the Indicator Formation to produce peak
diagenetic P1 phosphates and rutile. The addition of P to
basinal fluids may have additionally facilitated the transport
of U by PO4- complexes (Kyser and Cuney, 2008b).
Basinal fluids penetrated up to about 10 m beneath the
basal unconformity based on the distribution of M1 muscovite in basement rocks. Chown and Caty (1983) reported a
similar thickness for pre-Otish/Mistassini Basin regolith developed on the basement rocks, indicating that basinal brines
readily penetrated the regolith but were unable to penetrate
relatively fresh basement rocks beneath. Basement fault
zones, such as the dolomite-healed fault zone associated with
Camie River U mineralization, were not preferentially conducting peak diagenetic basinal brines based on a similar distribution of M1 in this zone relative to nonfaulted basement
rocks. This indicates that these older faults were closed to
fluid-flow during peak diagenesis and may not have been reactivated at this point in basin evolution.
H isotope ratios indicate that seawater was a component of
basinal brines in the Indicator Formation. Seawater was likely
entrained in Otish Basin sediments during the marine transgression from the southeast that is represented by the dolostone- and evaporite-bearing Peribonca Formation (Höhndorf et al., 1987; Genest, 1989). Marine transgression and
subsequent burial dissolution of the evaporites probably promoted the formation of albite cement by providing a source
of Na (Land and Milliken, 1981; Land et al., 1987). The association of pore-filling albite cement and diagenetic aquitards
in the Indicator Formation illustrates the potentially detrimental effect that marine influence can have on the paleohydrogeology and potentially ore-forming processes in continental sedimentary basins.
Late diagenesis and U mineralization: The 207Pb/206Pb date
of 1721 ±20 Ma for Camie River uraninite is within error of
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2000-1750 Ma
~250°C
a
U
M1
(o) basinal
fluids
U
paragonitic M1
U
albite
U
M1
S
graphitic,
sulfidic schist
1750-1600 Ma
250-300°C
gneiss/
intrusives
qtz.+ chlorite
metapsammite
N
b
Otish Gabbro
U
phengitic M1
U
C1
Tur
U
Tur
Camie River U
1200 Ma-recent
<150°C
recent
meteoric water
c
illite +
Mnt + K2
FIG.18. Schematic cross sections depicting the evolution of the western
Otish Basin and the Camie River U prospect. a. Oxidizing basinal brines
(black arrows) interacted with detritus of the Indicator Formation (horizontal
gray lines indicate depositional sequences) during peak diagenesis, which
mobilized U, produced muscovite (M1) in diagenetic aquifers (white fill),
and reacted with earlier albite cement to produce paragonitic M1 to the south.
b. The Camie River U prospect formed at 1721 ±20 Ma (Fig. 14a), coinciding with the intrusion of the Otish Gabbro at ca. 1730 Ma. U mineralization
was followed by the assemblage chlorite (C1), tourmaline (Tur), and phengitic M1 muscovite. Otish Gabbro intrusion resulted in preferential fluid flow
(white fill, black arrows) along fractures, faults, and depositional sequence
boundaries. c. Grenville orogenesis (1.2–1.0 Ga) uplifted the southern basin
margin. Fracture- and fault-preferred flow of late fluids (white fill) altered
M1 + C1 at temperatures <150°C to produce illite, montmorillonite (Mnt),
and kaolinite (K2), which have been influenced by recent meteoric water.
the 1692 ±32 Ma U-Pb upper intercept date, a previously determined 1723 ±16 Ma apparent age of Camie River uraninite (Höhndorf et al., 1987), and the 1717 ±20 Ma apparent
age of basement-hosted U mineralization at Lorenz Gully
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BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
(Höhndorf et al., 1987). This age of U mineralization coincides with the 1750 to 1730 Ma age reported for the intrusion
of the Otish Gabbro (Chown and Archambault, 1987;
Gatzweiler, 1987; Höhndorf et al., 1987), which was a significant thermal and tectonic event in the Otish Basin. The intrusion would have increased the thermal gradient, promoted
convective flow of basinal brines through diagenetic aquifers
and fault and fracture systems (Fig. 18b), and triggered U
mineralization throughout the western Otish Basin. Brittle
fractures, some of which crosscut M1 and could have propagated during intrusion of the Otish Gabbro, may have allowed
U-bearing basinal brines to penetrate the pervasive and fluidrestricting albite cement that is common near Camie River,
and the unconformity. Upon interaction with massive-sulfide−bearing basement rocks, U+6 in the basinal fluid was reduced to U+4 to produce uraninite. Based on petrographic observations, S0 pyrite and galena within the massive-sulfide
zone significantly contributed to the reduction of U (Fig. 7d,
e). Furthermore, the lack of abundant hematite near the deposit, which has been observed proximal to other unconformity-related U deposits (Hoeve and Sibbald, 1978; Polito et
al., 2004, 2005; Alexandre et al., 2005), may implicate sulfur
as a participant in the reduction of U.
Tourmaline + C1 chlorite ± phengitic M1 ± minnesotaite
comprises a relatively coarse-grained mineral assemblage that
is paragenetically indistinguishable from U1a + U1b uranium
minerals. However, that these minerals were related to the U
mineralizing process is inconsistent with their distribution
that includes barren areas away from U mineralization. The
assemblage instead probably formed following U1 minerals,
after basinal brines had interacted with the Otish Gabbro intrusions that would have provided the Mg, Fe, Ti, and Cr present in the minerals (Tables 2, 4). The coarse grain size of the
assemblage suggests the minerals formed at higher temperatures that accompanied Otish Gabbro intrusions, which may
be reflected by the 300°C estimated temperature of C1 formation (Table 2; Fig. 10). It is possible, however, that this
temperature instead reflects a subsequent fluid event, based
on stable isotope ratios in C1 that show the influence of a
later, metamorphic fluid (Table 6; Fig. 13).
Higher temperatures that accompanied intrusion of the
Otish Gabbro probably promoted coarse recrystallization of
M1 muscovite (Fig. 7b), including intergrowth with C1 (Fig.
7c), which may have decreased the fluid-conducting capability within diagenetic aquifers, based on the preferential distribution of tourmaline and C1 in fault zones and along the
basal unconformity (Fig. 18b). Higher temperatures also
probably promoted phengitic substitution in M1, which has
been described from low-T, low to high-P metamorphic terrains (e.g. Velde, 1965; Sassi et al., 1994; Massonne and Kopp,
2005). A high-P origin for phengitic substitution is not favored, considering the unlikelihood of a strong pressure gradient between Matoush and Camie River to explain the high
proportion of phengitic M1 at Matoush (Fig. 11). Instead, a
low-T, low-P origin for phengitic substitution is favored,
where substitution is influenced by the bulk composition of
the host rock, and is preferred in K-spar-dominated lithologies (Sassi et al., 1994). Thus the degree of phengitic substitution observed in M1 may reflect differing K-spar content
within the Indicator Formation, with higher stratigraphic
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421
levels at Matoush (Fig. 3b) containing more K-spar than the
lower stratigraphic levels at Camie River.
Postdiagenetic alteration: Post-U 40Ar/39Ar dates of M1 in
basement rocks and in Indicator Formation fractures may reflect the timing of a hydrothermal event responsible for the
post-U assemblage C2 sudoite + S2 sulfides + siderite. Based
on the ca. 1400 Ma plateau age of fracture-hosted M1 (Fig.
15), the hydrothermal event may have been associated with
ca. 1.4 Ga anorogenic magmatism that has been implicated
with hydrothermal alteration in other Proterozoic basins (e.g.,
Medaris et al., 2003). The event also remobilized a small
amount of U to form late U2 phosphate in basement rocks.
Additionally, the close spatial association shared by S2c galena
and U1a/U1b uraninite/brannerite suggests that radiogenic
Pb sourced from radioactive decay of U in the latter minerals
likely influenced the growth of the former.
Metamorphic fluids in the Otish Basin, typified by a fluid
mixing trend to higher δ18O values and lower δ2H values (Fig.
13), likely originated from subgreenschist-grade metamorphism in the Otish Basin during the ca. 1.2 to 1.0 Ga
Grenville Orogeny (Chown, 1979; Rivers and Chown, 1986).
C1 chlorite was the most affected by metamorphic fluids (Fig.
13) due to its distribution in preferential fluid-flow pathways
(basal unconformity and fault zones), whereas M1 and phengitic M1 were variably influenced by metamorphic fluids depending on their proximity to these zones. The high δ18O values of metamorphic fluids reflect much lower water/rock
ratios during metamorphism relative to diagenesis, which
most likely preserved pre-Grenville U-Pb and Ar isotope ages
at Camie River, despite the location of the study area within
about 20 km of the Grenville Front Tectonic Zone (Fig. 1).
Grenville-aged U-Pb isotope resetting in uraninite at Indicator Lake (Fig. 1), which is situated near a major intra-Otish
Group fault, is suggested by a date of 1072 ±5 Ma (Höhndorf
et al., 1987) and suggests that fluid flow in the Otish Basin
was restricted to regional-scale fault zones at this point in
basin evolution. Grenville orogenesis also resulted in reverse
and strike-slip faulting in the area (Chown, 1979; Rivers and
Chown, 1986) and is likely responsible for the ~400-m offset
observed in the southern study area (Figs. 4, 6).
M1 muscovite and C1 chlorite in fault zones are partially altered to illite and montmorillonite (Figs. 9, 11), which reflect
a late, low-temperature fluid event, based on the estimated
formation temperatures of illite (Fig. 10) that was also coeval
with the formation of K2 kaolinite in these zones (Fig. 18c).
The late event may represent the uplift of the Otish Basin subsequent to Grenville orogenesis and the incursion of meteoric
waters into fault zones and fractures. Additionally, fault zonehosted C1 and some M1 from Camie River have been influenced by fluids with low δ2H values (Figs. 13, 18c) not unlike
modern meteoric waters in the Otish Basin. The preferential
H isotope exchange affecting fault-hosted phases suggests
that these zones presently remain as fluid-flow conduits.
Implications for unconformity-related U exploration
Known U mineralization at Camie River provided the opportunity to test the 2% HNO3 method in the older Otish
Basin that evolved differently than the Athabasca, Thelon,
and Kombolgie Basins in which the method was developed
(Holk et al., 2003). In the latter basins, diagenetic aquifers
421
422
BEYER ET AL.
conducted fluids up to ca. 350 m.y. following U mineralization
(Kyser et al., 2000; Polito et al., 2006; Alexandre et al., 2009a,
b; Hiatt et al., 2010), which helped preserve intrastratal fluidflow pathways facilitating the dispersion of radiogenic Pb and
pathfinder elements from deposits. Following Otish Gabbro
intrusion and U mineralization in the Otish Basin, only fluidflow pathways along unconformities and in fault zones and
fractures were preserved, based on the preferential distribution of unsupported radiogenic Pb and Mo + W + Nb, which
are characteristic of Camie River U mineralization in these
zones (Fig. 16). Samples of basement rocks and the Indicator
Formation away from these zones, including locations proximal to U mineralization, do not contain unsupported radiogenic Pb in most cases or detectable Mo + W + Nb, suggesting little to no penetration by post-U mineralization fluids
away from preferential fluid-flow zones (Fig. 16). Due to the
selective distribution of radiogenic Pb and pathfinder elements near Camie River, the successful employment of the
2% HNO3 method as a vector to Camie River-style mineralization would likely have to be accompanied by structural
mapping and sampling of fault zones, fractures, and depositional sequence boundaries.
The portion of the Camie River U prospect examined
shares the key characteristics of basement-hosted, or simpletype unconformity-related U deposits in the Athabasca and
Kombolgie Basins (Fayek and Kyser, 1997; Polito et al.,
2004, 2005; Alexandre et al., 2005; Cloutier et al., 2009), as
Camie River U mineralization resulted from the interaction
of oxidizing, U-bearing basinal brines with reducing,
metasedimentary basement lithologic units near the unconformity during a major tectonic event (Otish Gabbro intrusion) in the basin. The Camie River U prospect appears different from these deposits by its apparent complex
geochemical association (cf. Gatzweiler, 1987). However, in
contrast to sulfides, arsenides, and selenides that are coeval
with U mineralization in Athabasca complex-type U deposits
(e.g., Fayek and Kyser, 1997), Mo, Nb, Co, and Ni at Camie
River are attributed to S2 molybdenite, columbite, and S0S1 sulfides, respectively, that are paragenetically distinct
from U1 minerals, and therefore may not be indicative of U
mineralization. The Camie River U prospect is truly different from simple-type U deposits by its lack of a pronounced
alteration halo. U mineralization at Camie River is spatially
associated with a narrow zone of M1 muscovite alteration in
the basement rocks, which could be easily missed during exploration relative to the chlorite + illite + hematite alteration
zones that can accompany basement-hosted U deposits in
the Athabasca and Kombolgie Basins. Additionally, while
previous studies describe a zoned alteration halo consisting
of chlorite, carbonate, and albite in the Indicator Formation
near the Camie River U prospect (Ruzicka and LeCheminant, 1984; Gatzweiler, 1987), mineral crosscutting relationships demonstrate that each of these phases is not genetically
related to U1 minerals, and therefore does not provide predictive indicators of U mineralization.
The Camie River U prospect expands the simple-type unconformity U model by adding massive sulfide-bearing basement rocks to the list of potential reductants for U, in addition to graphite- and Fe-rich chlorite-bearing basement
lithologic units (Polito et al., 2004; Alexandre et al., 2005).
0361-0128/98/000/000-00 $6.00
The prospectivity of the Camie River area, and additional
areas in the western Otish Basin and neighboring Papaskwasati Embayment that are underlain by greenstone
belts (Fig. 1), greatly depends on the presence and continuity of the massive sulfide + graphite belts that play a substantial role as a U trap based on unambiguous petrographic
relationships in both hand sample and in thin section (Fig.
7d, e). The basement rocks at Camie River may be best
suited for a reducing role in a simple-type unconformity-related U model. Narrow basement alteration zones marked
by M1 muscovite and tourmaline, and their sharp juxtaposition with unaltered basement rocks beneath a possible thin
regolithic zone (Chown and Caty, 1983) suggest little basement rock interaction with either basement or basinal fluids
and a decreased probability of forming a complex-type U
deposit.
Age of the Otish Gabbro
The unpublished work of Hamilton and Buchan (2007),
which reported a U-Pb age of ca. 2169 Ma from a single baddeleyite crystal in the Otish Gabbro, adds uncertainty to the
age of the intrusion and whether or not multiple intrusions
make up the Otish Gabbro. Together with an unpublished UPb baddeleyite age of ca. 2500 Ma for the pre-Otish Basin
Mistassini dikes (Heaman, 2004), the unpublished Otish
Gabbro age suggests that the Otish Basin formed between
ca. 2500 and 2170 Ma, similar to the 2450 to 2219 Ma
Huronian Supergroup (Bennet at al., 1991) in Ontario. The
results of our study, however, emphasize the fluid evolution
of the Otish Basin, specifically the development of basinal
brines in diagenetic aquifers and their evolution during ca.
400 m.y. of diagenesis and interaction with the Otish Gabbro
(Fig. 5). The duration of diagenesis in the Otish Basin is not
unlike that in the Athabasca Basin, where diagenetic aquifers
hosted fluids for ca. 350 m.y. following basin formation based
on 40Ar/39Ar dates of diagenetic illite (Kyser et al., 2000;
Alexandre et al., 2009a, b). This extended interval of diagenesis could have occurred during an earlier interval (ca.
2500−2170 Ma) in contrast to what our data suggests. In
such a scenario, Otish Gabbro intrusion at ca. 2170 Ma could
have been associated with initial U mineralization, and the
ca. 1720 Ma age of uraninite at Camie River and at the distant Lorenz Gully showing (Höhndorf et al., 1987; this study)
could have resulted from a later basin-wide Pb-loss event
possibly associated with post-Penokean orogenic magmatism
that is recognized in the Huronian belt (i.e., Cutler batholith,
Cannon, 1970). However, the Huronian Supergroup hosts
quartz-pebble conglomerate−associated U deposits that were
likely paleoplacer in origin and preserved under oxygen-poor
atmospheric conditions (e.g., Ono and Fayek, 2011), which
likely persisted during deposition of the basal two-thirds of
the Huronian Supergroup (Roscoe, 1973; Young et al., 2001).
Additionally, U deposits with redox-controlled trapping
mechanisms are not known from the Huronian Supergroup
but are the only such deposits found in the Otish Basin.
Therefore, it is more likely that the formation of basinal
brines that were sufficiently oxidizing to mobilize U throughout the Otish Basin occurred in a younger basin bracketed by
the dates reported in the peer-reviewed literature and used
in this study.
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BASIN EVOLUTION AND UNCONFORMITY-RELATED U MINERALIZATION: CAMIE RIVER, QUEBEC
Conclusions
1. Diagenetic aquifers were thick and largely unconfined
in the Indicator Formation near Camie River and conducted
basinal brines with a composition and temperatures appropriate for leaching and transporting U from basinal sediments. Therefore the Otish Basin had the hydrogeologic
framework and fluid characteristics required to form unconformity-related U deposits. However, the association of porefilling albite cement and diagenetic aquitards in the Indicator Formation illustrates the potentially detrimental effect
that marine influence can have on the paleohydrogeology
and potentially ore-forming processes in continental sedimentary basins.
2. U mineralization at Camie River, and at least one other
location, was triggered by the ca. 1730 Otish Gabbro intrusion. Our fluid evolution model for the Otish Basin is compatible with an interpretation of a Huronian-age equivalent
Otish Basin as suggested by the unpublished work of Hamilton and Buchan (2007), but is not favored due to contrasting
U mineralization styles in the Otish and Huronian Basins and
the oxygen-poor atmospheric conditions that accompanied
deposition of much of the Huronian Supergroup.
3. The preservation of ca. 1720 to 1410 Ma U-Pb and PbPb isotope and Ar-Ar dates at Camie River suggests that
minimal fluid interaction with uraninite and muscovite occurred during Grenville orogenesis, and that diagenetic
aquifers were probably indurated and unable to conduct fluids by 1200 Ma. Basin-scale fluid flow was probably restricted to major, regional-scale faults by this point in basin
evolution.
4. Fault zones, fractures, and depositional sequence
boundaries are preferential conduits for the dispersion of unsupported radiogenic Pb and the trace elements Mo + W +
Nb from Camie River U mineralization and can be utilized to
explore for additional deposits. Away from these zones, including locations proximal to U mineralization, unsupported
radiogenic Pb in most cases or Mo + W + Nb are not detectable, suggesting little to no penetration by post-U mineralization fluids away from preferential fluid-flow zones
5. The portion of the Camie River U prospect examined is
most similar to a basement-hosted, or simple-type unconformity-related U deposit but is unique in that basement-hosted
sulfide minerals, rather than graphite or Fe chlorite, played a
major role in reducing U. Additionally, alteration is restricted
to a narrow zone of muscovite and lacks chlorite and
hematite.
Acknowledgments
The authors thank Don Chipley, Kerry Klassen, Bill MacFarlane, and April Vuletich for assisting with stable and radiogenic isotope analysis at the Queen’s Facility for Isotope Research. Al Grant, Queen’s University, and Peter Jones,
Carleton University, Ottawa, assisted with SEM and electron
microprobe analyses. Brad Singer and Brian Jicha, University
of Wisconsin, are thanked for Ar/Ar dates. Clayton Durbin of
Cameco Corporation compiled maps of the Camie River area.
The research was funded by Cameco Corporation, a Natural
Sciences and Engineering Research Council Collaborative
Research Development grant, and a Society of Economic Geologists Canada Foundation grant.
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423
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