Disturbed 40Ar-39Ar systematics in hydrothermal biotite and

Geochimica et Cosmochimica Acta, Vol. 61, No. 21, pp. 4655-4669, 1997
Copyright 0 1997 Elsevier Science Ltd
Pergamon
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PII SOO16-7037(97) 003504
Disturbed 40Ar-39Arsystematics in hydrothermal biotite and hornblende at the Scotia
gold mine, Western Australia: Evidence for argon loss associated with
post-mineralisation fluid movement
ADAM J. R. KENT’**
and T. CAMPBELL MCCUAIG’.+
‘Research School of Earth Sciences, The Australian National University, Canberra 0200, Australia
‘Deparhnent of Geological Sciences, University of Saskatchewan, Saskatoon, S7N 5E2, Canada
(Received May 6, 1996; accepted in revised form July 21, 1997)
Abstract-Hornblende
and biotite that formed during gold mineralisation at the Scotia mine, Western
Australia, have erratic 4oAr-3gAr release spectra and total gas ages that are -200-900
million year
younger than the ca. 2600-2620 Ma minimum age of gold mineralisation, as given by 4oAr-3gAr plateau
(muscovite) ages of crosscutting pegmatite dykes. Analysed homblendes are dominated by magnesio
hornblende but also contain small domains of ferro-actinolitic hornblende, actinolitic hornblende, and
actinolite. Biotite also appears to be substantially altered to chlorite along cleavage planes.
Relatively young apparent ages and high K/Ca ratios of argon released from homblendes at temperatures less than -1000°C are interpreted to be the result of degassing of contaminant biotite. However,
this cannot totally explain the young ages of homblendes. Gas fractions released at furnace temperatures
above lOOO”C,where the effect of biotite degassing is demonstrably negligible, still have apparent ages
that are -200-900
million years younger than the age of muscovite from post-gold pegmatite dykes.
The close proximity of disturbed hydrothermal hornblende samples to apparently undisturbed pegmatite
muscovite samples (less than a few metres in some cases) is difficult to reconcile with argon loss in
hydrothermal hornblende being the product of thermally-driven volume diffusion. Given a suitable
thermal history, argon loss could occur preferentially in homblendes if ( 1) the closure (for slow cooling)
and blocking (for reheating) temperatures of hydrothermal homblendes were lower than published
estimates, as has been observed in structurally complex metamorphic homblendes and/or (2) the closure
and blocking temperature of pegmatite muscovite were higher than commonly estimated. However,
neither of these interpretations can easily explain the large variation in hornblende ages.
It is instead suggested that argon loss occurred during mineral-fluid interaction during movement of
a retrograde fluid along the mineralised lode structures and that this occurred at ambient temperatures
below the blocking temperature of pegmatite muscovite. There is abundant geological evidence for the
passage of such a fluid at the Scotia mine, including the presence of numerous late brittle fractures
containing retrogressive low-temperature mineral assemblages. Late fluid movement is probably related
to Proterozoic erogenic activity along the nearby southeastern margin of the Yilgarn Craton.
The difference in argon systematics between hydrothermal minerals and pegmatite muscovite is largely
ascribed to the relatively low permeability of the more massive pegmatite dykes (with respect to ore
zones) preventing fluid egress to muscovite samples. The variations in ages of hydrothermal minerals
are probably related to the extent of fluid/mineral interaction as this is a function of parameters, such
as fluid/rock ratio, fluid P-T-X conditions, permeability, and mineral microstructures that may vary on
short time and length scales. Recognition of possible argon loss in hornblende via fluid interaction is
important for the interpretation of 4oAr-3gAr systematics in environments, such as many hydrothermal
ore deposits, where minerals may be exposed to fluids after crystallisation.
Copyright 0 1997 Elsevier
Science Ltd
1. INTRODUCTION
feldspar. This has especially been with respect to the extent
and nature of argon loss in these minerals during both laboratory (e.g., Hanson et al., 1975; Wartho et al., 1991; Lee,
1993) and geological (e.g., Turner et al., 1969; Harrison and
McDougall, 1980; Wartho, 1995; Lister and Baldwin, 1996)
processes and has provided a theoretical background on
which to base interpretations of 4oAr-3qAr data.
To date, understanding of geological argon loss has been
largely directed at that which occurs during slow cooling or
episodic reheating of minerals. This has been described using
theory based on volume diffusion whereby argon retentivity
of a specific mineral is dependent on factors such as the
diffusivity of argon and size of individual diffusion domains
within the mineral in question and the time-temperature his-
The 40Ar-39Ar technique constitutes one of the most widely
applied isotopic dating techniques in the geosciences and is
well suited to the investigation of many geological problems
(e.g,. McDougall and Harrison, 1988). Commensurate with
the wide application of the technique has been the detailed
investigation of the behaviour of argon systematics in appropriate minerals, such as hornblende, muscovite, biotite, and
*Present address: Department of Geological and Planetary Sciences, California Institute of Technology, Mail Stop 170-25, Pasadena, California 91125, USA (adam@expet.gps.caltech.edu).
+Present address: Etheridge Henley Williams, Suite 34A, 25 Walters Drive, Osborne Park, Western Australia 6017, Australia.
4655
A. J. R. Kent and T. C. McCuaig
4656
tory experienced by the sample (e.g.. Turner, 1969; Dodson.
1973; McDougall and Harrison, 1988; Lister and Baldwin,
1996). For slowly cooled samples (subject to the sample
experiencing a cooling rate that is linear with the reciprocal
of temperature) the dependence of argon diffusion on temperature is conveniently
expressed in terms of the closure
temperature of the mineral (Dodson, 1973); at temperatures
above the closure temperature diffusive argon loss is too
rapid to allow accumulation of argon within the mineral. It
is only at temperatures below the closure temperature that
argon accumulation can occur, and the geological clock can
commence to mark time. For minerals that have been held
at high temperatures for extended periods of time or experienced transient reheating events. following the usage of
Lister and Baldwin ( 1996), the term blocking temperature
is used to describe the temperature above which significant
argon loss occurs.
It is also known that argon loss can occur by mechanisms
other than volume diffusion. During laboratory heating in
vacua phase transitions and dehydroxylation
(in hydrous
minerals) are often more important than volume diffusion
in dictating argon degassing behaviour (e.g., Hanson et al..
1975; Gaber et al.. 1988: Lee et al.. 1991: Wartho et al.,
1991; Lee. 1993). In natural samples it is more difficult to
document argon loss by alternate mechanisms.
However.
there is emerging recognition
that, in addition to volume
diffusion, argon loss can be the result of chemical re-equilibration of potassic minerals (Wartho. 1995 )
This is especially germane to the interpretation
of geochronological
data from many hydrothermal
ore deposits
where the structures responsible for localising hydrothermal
activity and mineralisation may often act as locii for further
fluid flow after mineralisation.
In this contribution we report
an example of the disturbance of argon systematics in hydrothermal biotite and hornblende that formed during gold mineralisation at the Scotia gold mine, located adjacent to the
southeastern margin of the Yilgam Craton in Western Australia (Fig. 1). In this location argon loss in biotite and
hornblende was sufficient to produce variable apparent ages
-200-900
million years younger than the time of gold mineralisation. However, this did not result in argon loss in
nearby pegmatite muscovites. We suggest that, rather than
argon loss occurring via volume diffusion controlled by the
thermal history of this region. argon systematics of hydrothermal minerals can be best understood via argon loss during interaction between hydrothermal minerals and infiltrating post-mineralisation
retrograde fluids. In this situation the
thermal history of the sample may play a subsidiary role to
the hydrothermal history and parameters governing the extent of fluid-mineral interaction in controlling argon systematics.
2. THE SCOTIA MINE
The Scotia gold mine is located at the southern end ot
the Norseman-Wiluna
greenstone belt within the Norseman
Terrane of the Archaean Yilgam Craton (Fig. I : Swager et
al., 1990). The mine is situated about 70 km northwest of
the Fraser Front, which delineates the southern boundary of
the Yilgam Craton with the Proterozoic Fraser Province to
the south (Fig. I; Gee, 1979).
. .
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.......
.......
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.......
.......
.
.........
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.........
........
3
Granit
fl
Kalgoorlie Terrane greenstone sequences
oid
Norseman Terrane greenstone sequences
1
Undiffe rent iated greenstone sequences
a
Proterozoic Fraser Orogen
/Major
fault
Maior lode oold deoosit
Fig. 1. Regional map of the Kambalda-Norseman
region, showing
the location of the Scotia deposit with respect to the location of the
southeastern
margin of the Yilgam Craton and other major gold
deposits.
Mineralisation
at the Scotia mine is hosted within the
Woolyeenyer
Formation, a thick sequence of tholeiitic basalts. gabbroic dykes and sills. and high-MgO
tholeiitic
dykes. A gabbroic intrusion into the Woolyeenyer Formation
at Norseman has a U-Pb age of 2714 5 5 Ma (Hill et al.,
1992 ) . In the Scotia mine area this formation has been metamorphosed to middle amphibolite facies (McCuaig et al.,
1993 ) Hornblende-plagioclase
geothermometry
indicates
peak metamorphic temperatures between 530-700°C for the
rocks hosting the Scotia deposit (McCuaig et al., 1993).
Although the timing of metamorphism
is not well constrained, it is probable that this either occurred broadly contemporaneously
with intrusion of large volumes of granitoids
into the southern Norseman-Wiluna
Belt between 2660-
Disturbed MAr/‘9Ar release spectra in a gold mine
4657
the multiple generations of these structures as indicated by
crosscutting relationships, and the variable low-temperature
alteration assemblages associated with them, collectively indicate that brittle deformation occurred after, and at lower
pressures and temperatures, than gold mineralisation and
pegmatite intrusion. Retrogressive minerals similar to those
found in late fractures also occasionally occur as small flakes
and grains (generally +5 pm) on grain boundaries and cleavage surfaces of prograde hydrothermal minerals.
muscovite
- garnet
pegmatite
3. SAMPLE DESCRIPTION
7/6/70
Ramp
Fig. 2. Field relations between a mineralised quartz-diopside-homblende-biotite bearing shear zone and garnet-muscovite pegmatite
dyke underground at the Scotia gold mine (7/670 ramp).
2690 Ma (Binns et al., 1976; Hill et al., 1992) or during
metamorphism within the lower crust of the Yilgarn Craton
at ca. 2630-2650 Ma (Nemchim et al., 1994; Kent et al.,
1996).
Gold mineralisation at Scotia occurs within a NNW to N
striking, east-dipping ductile shear zone and extends discontinuously for 3 km along strike (Thomas et al., 1990). Ore is
mostly hosted within multiple quartz-diopside-calcite veins.
Within ore zones multiple veining events have produced
complex composite veins with alternating bands of quartz,
diopside, calcite, amphibole, biotite, and microcline (McCuaig et al., 1993 ) . Alteration selvages surrounding individual veins consist of an outer zone of biotite-homblendeplagioclase, which is replaced with increasing proximity to
veins by hornblende-plagioclase-quartz,
then diopsidequartz-calcite ( t microcline, zoisite, actinolite) assemblages. In composite veins, however, this zonation is often
complicated by overlapping alteration envelopes. Microprobe analyses reveal that hydrothermal amphibole compositions within diopside-quartz-calcite
veins are actinolite to
actinolitic hornblende, whereas hydrothermal amphiboles in
altered basalt adjacent to veins are actinolitic hornblende to
magnesio hornblende. These amphiboles are all coeval with
vein emplacement and associated hydrothermal activity.
Gold occurs as occasional small (<lo pm) grains, often
in association with pyrrhotite and chalcopyrite. Vein and
alteration assemblages are consistent with minimum temperatures for gold mineralisation in excess of 500°C (McCuaig
et al., 1993).
Quartz-muscovite-albite
+- biotite ? garnet pegmatites
crosscut ore bodies in many locations and in all observed
examples postdate mineralisation (Fig. 2; McCuaig et al.,
1993). Ore zones and pegmatites were locally cut by E-W
dolerite dykes dated elsewhere in the Norseman Terrane at
ca. 2400 Ma (Fletcher et al., 1987). Both ore zones and
pegmatites are disrupted by multiple generations of late brittle fractures (Thomas et al., 1990; McCuaig et al., 1993).
These structures are variably filled with assemblages of chlorite-quartz-sericite-albite
and albite-prehnite, fine-grained
cataclasite, and gypsum. The brittle nature of these faults,
Muscovite was separated from pegmatite samples taken
from two different localities: the 7/670 ramp (sample 91237) and 71440 north drive (sample ScPegS). In both these
locations pegmatites occur as shallowly dipping dykes which
crosscut ore zones. Field relations between pegmatite and
ore zone at the 7/670 ramp are detailed in Fig. 2. Muscovite
in pegmatites occurs mainly in books of large crystals, generally 0.5-2.5 cm across and up to 10 cm long. Garnet and
albite were also separated from sample 91-237 and analysed
for Sm-Nd, the results of which are detailed in Kent ( 1994)
and Kent et al. ( 1996). The grain size of analysed muscovite
samples was 400-200 brn for &Peg8 and 250- 150 pm for
91-237.
Hornblende and biotite were separated from three samples: 91-205, taken from within an ore zone on the 7/670
ramp; sample ScPvAm2, taken from the ore zone in the 7/
440 north drive; and sample 91-233 from the 4/550 ramp.
All samples consist of banded homblende-quartz-calcitepyrrhotite-diopside-biotite-plagioclase
altered metabasalt. In
sample 91-205 hornblende and biotite rich bands alternate
with quartz-plagioclase-calcite dominated bands. Both homblende and biotite form a foliation parallel to mineralogical
banding although regions of decussate biotite and sheaflike
whorls of hornblende are also occasionally apparent. Larger
crystals of hornblende also crosscut foliated biotite. Electron
microprobe investigations reveal that homblendes are predominantly magnesio hornblende (Table 1)) although there
is some chemical variations, generally best expressed in
A1203, MgO, and CaO contents. Typical compositions and
overall compositional ranges are given in Table 1. K20 contents are typically 0.20-0.25 wt%. In thin section, biotite
grains have pale brown to yellow-brown pleochroism and
an irregular ragged birefringence commonly associated with
partial retrogression to chlorite along cleavage planes. This
can be observed in backscattered electron images which indicate that interlayer chlorite intergrowths are ubiquitous
within biotite and are generally on the l-2 pm scale (Fig.
3). Consistently low K20 (5-7%; Table 1) from electron
microprobe measurements also implies that biotite from this
sample has suffered some degree of retrogression to chlorite.
This is a common feature in many Archaean gold deposits
formed at similar pressure-temperature conditions. The timing of chlorite retrogression is largely unknown.
Samples 91-233 and ScPvAm2 consist of banded homblende-quartz-calcite-pyrrhotite
bearing altered metabasalt.
Hornblende occurs as elongate prismatic crystals, largely in
hornblende dominated bands, but also intergrown with
quartz and calcite. As with sample 91-205, hornblende is
A. J. R. Kent and T. C. McCuaig
4658
Table
111
91-205
Ferri-actinolitic
hornblende
Cr20,
Fe0
MnO
MgO
CaO
Na,O
K20
Cl
Total
WCs
prop.
wt%
wt%
prop.
34.13
1.03
17.41
0.36
17.42
co.09
15.96
<0.06
5.220
0.119
3.139
0.043
2.228
6.900
0.017
1.691
0.009
1.630
0.035
2.716
1.812
0.285
52.28
0.15
3.08
0.08
16.43
0.28
14.51
11.05
0.12
7.416
0.017
0.511
0.009
1.952
0.034
3.069
1.679
0.017
0.22
0.00
97.93
0.043
0.08
0.01
98.06
0.034
0.021
15.138
0.008
Atom
prop.
48.08
0.17
9.99
0.08
13.61
0.29
12.71
11.77
1.02
14.738
<O.ll
7.01
na
93.32
from the Scotia mine.
[41
91-233
Magnesio
hornblende
131
91-205
Biotite
Atom
Atom
SiOz
Ti02
Al&s
data for minerals
PI
91-205
Magnesio
hornblende
wt%
1. Electron microprobe
3.638
1.368
15.755
>134
I51
91-233
Magnesio
hornblende
Atom
Atom
wt&
prop.
[61
91-237
Muscovite
wt%
prop.
Atom
wt%
7.044
0.061
1.469
0.009
1.694
0.026
2.697
1.190
45.27
0.29
12.74
0.37
15.34
0.33
10.62
11.48
6.550
0.035
2.174
0.044
1.862
0.044
2.287
1.784
0.78
0.21
0.02
98.60
0.216
0.035
I .38
0.24
0.03
98.09
0.392
0.044
0.42
11.21
0.109
1.904
15.216
9?26
14.099
0.020
47.18
<0.08
33.49
<0.09
4.84
0.12
<0.08
<0.07
prop.
48.95
0.56
8.68
0.16
14.10
0.23
12.59
12.37
15.161
[71
Amphibole
range
6.280
5.254
0.538
0.014
0.025
wt%
45.27-52.28
0.13-0.56
12.74-3.08
0.08-0.37
12.10-15.34
0.23-0.34
10.62-15.36
11.05-12.37
0.12-1.38
0.08-0.24
0.00-0.02
0.008-0.035
Note: Atomic proportions calculated using twenty-two (biotite, muscovite) and twenty-three (hornblende) oxygen atoms.
na-not analysed.
largely parallel to mineralogical
banding but occasional
sheaflike whorls are apparent. Magnesio hornblende is again
dominant, although actinolitic hornblende and ferro actinolitic hornblende domains are present on many large magnesio
hornblende crystals (Table 1).
In all samples amphibole compositions vary on small spatial scales. For example, in Fig. 3, within the space of a few
microns at the grain boundary, magnesio hornblende with
- 10 wt% A1203 changes to ferro-actinolitic
hornblende with
-3 wt% A&O3 (analyses 1 and 2 in Table 1) All hydrothermal hornblende analysed are predominantly
magnesio homblende, although small regions of actinolite or actinolitic
hornblende occur at grain boundaries, or as small regions
(- 10 pm) in the interior of larger grains (Fig. 3). The
observation that compositional
variations within individual
homblendes
mirror those observed between different alteration zones (e.g., vein-distal hornblende-dominated
assemblages and vein-proximal
actinolite and actinolitic homblende; see above and McCuaig et al., 1993) suggest that
fine-scale variations in amphibole compositions
are largely
related to superposition
of different alteration facies during
progressive hydrothermal
alteration. Thus, magnesio homblende which originally formed at the margin of a quartzcalcite vein was converted to actinolite and actinolitic homblende at grain boundaries during emplacement
of subsequent nearby auriferous quartz-calcite veins. This interpretation implies that variations in amphibole composition formed
during the original hydrothermal
mineralisation
and alteration process, and not during pegmatite intrusion or the postpegmatite
fracturing
and retrogression
event described
above. This is consistent with the observation that amphibole
is not part of the retrogressive
mineral assemblage
(e.g.,
chlorite, sericite, prehnite, gypsum, albite). However, it must
also be noted that given the complexity of amphibole compositional variations, it is impossible to rule out some degree
of chemical modification of amphibole (similar perhaps to
that described by Wartho, 1995 ) during subsequent retrogression.
Grainsizes of analysed amphibole separates were 400200 km for ScPvAm2 and 180- 100 km for 91-205 and
91-233.
Stable isotope data is also reported for samples SC16366, SC358-33, X358-40b,
SC-358-55, and SC358-99 from
hydrothermal alteration assemblages from drill cores through
the Scotia orebody. These samples are similar in appearance
to those described above. Hydrothermal
quartz, amphibole
and clinopyroxene
were separated from these samples, and
preliminary stable isotope analyses obtained for these mineral separates as part of a larger study ( McCuaig, 1996) are
reported here.
4. ANALYTICAL
METHODS
Mineral separates from all samples were prepared to >99% purity,
using standard heavy liquid, magnetic separation, and handpicking
techniques. Mineral compositions
were analysed at the California
Institute of Technology using a JEOL JXA-733 Superprobe equipped
with five crystal spectrometers and a Tracer Northern EDS detector
using an accelerating voltage of 15 kV and beam current of 15 nA.
Data was reduced with the CITZAF PRZ correction algorithms using
natural and synthetic standards (Armstrong,
1995). Errors in major
elements are on the order of 2% whereas errors for minor elements
are in the order of 10%. Amphibole compositions
were calculated
using the EMP-AMP software. Backscattered
electron and EDS Xray maps were acquired with a Camscan Series II Scanning Electron
Microscope operating at 15 kV.
Samples were analysed by the J0Ar-39Ar technique at two separate
facilities. the Research School of Earth Science at the Australian
National University, Canberra, Australia (samples 91-205, 91-233,
91-237) and Queen’s University, Kingston, Ontario, Canada (samples ScPeg8 and ScPvAm2).
Although a detailed interlaboratory
study between these two facilities has not been conducted, this does
not influence the final conclusions of this study which is more concerned with resetting of argon systematics rather than a detailed
comparison of the isotopic ages of these samples. Samples analysed
at the Australian National University were irradiated in a cadmium-
Disturbed
40Ar/39Ar release
spectra
in a gold mine
4659
Fig. 3. SEM images of hornblende and biotite from sample 91-205. (a) Backscattered
electron image of hornblende
crystal with biotite (shown at the left and bottom of the image) showing a chlorite-filled
fracture in top right comer
and lighter-coloured
ferri-actinolitic
hornblende rim around the bottom of the crystal. Analysis 1 from Table 1 is from
the core of this crystal and analysis 2 is from the rim. (b) Backscattered
electron image of biotite crystal showing
alteration to darker chlorite along cleavage planes. (c) and (d) EDS X-ray maps for aluminium and silicon for the
same field of view shown in (a). Note Al-poor, Si-rich rim, and relatively Al-rich, Si-poor chlorite inclusions. The
Si- and Al-poor regions in the bottom right and top left of the hornblende crystal are pits in the section.
lined aluminium canister for two 24 day cycles in the HIFAR reactor,
Lucas Heights, Australia along with fluence monitor minerals GA1550 biotite for biotite and muscovite and 77-600 hornblende for
hornblende. Argon was extracted from irradiated minerals in a double vacuum resistance furnace and analysed isotopically using a VG
MM12 mass spectrometer.
Analytical details are described in more
detail in Kent and McDougall (1995). Samples analysed at Queen’s
University were irradiated for 116 h in position 5c of the watermoderated, enriched uranium research reactor at McMaster University, Hamilton, Ontario, Canada (samples ScPeg8 8 and ScPvAm2)
using international hornblende standard HB3GR as a fluence monitor. Argon was extracted using a Lindberg tubular furnace and analyzed using a MS-10 mass spectrometer.
Argon isotopic ratios measured from both facilities were corrected
for neutron induced production of argon isotopes from K and Ca.
Ages from plateau-like segments in age spectra are calculated using
gas-fraction weighting of individual steps and associated errors, and,
in order to make ages from the two separate facilities comparable,
? 0.5% error in the irradiation parameter J is incorporated
into all
total gas and plateau-like
segment ages. All ages were calculated
A. J. R. Kent and T. C. McCuaig
4660
Table 2. ““Ar-39Ar data for minerals from the Scotia gold mine.
Temp
(“C)
40Ar/3YAr
36Ar/39Ar
ScPeg8 Muscovite (400-200
(0.0161 g, J = 0.02325)
500
600
700
750
800
850
900
950
1000
1200
208.04
150.31
140.80
140.60
140.43
138.72
139.79
140.04
140.01
141.36
“Ar/39Ar
Vol 39Ar
(X10-* cm’)
Fraction
39Ar (%)
“OA@
0.009
0.079
0.006
0.001
0.000
0.001
0.000
0.000
0.000
0.128
0.032
0.268
4.550
2.363
1.781
2.285
3.995
3.173
1.516
0.225
0.2
1.3
22.6
11.7
8.7
11.3
19.8
15.7
7.5
1.1
60.8
97.6
99.9
99.8
99.8
99.8
99.9
99.8
99.8
97.1
126.38
146.66
140.76
140.36
140.15
139.45
139.65
139.75
139.65
137.23
(140.51)d
2473
2677
2620
2616
2614
2607
2609
2610
2609
2585
(2618
0.2503
0.4599
0.2438
0.1245
0.1493
0.0630
0.0596
0.0481
0.0552
0.0529
0.0412
0.0509
0.0530
0.0446
0.4182
0.1245
0.0354
0.0202
0.0305
0.0527
0.1058
0.1950
0.2782
0.3068
0.4167
0.7412
0.7656
0.6088
0.6014
0.3577
0.0233
0.0527
0.8
0.4
0.7
1.1
2.3
4.3
6.0
6.7
9.1
16.1
16.7
13.3
13.1
7.7
0.6
1.1
91.9
88.7
97.7
98.6
99.1
99.2
99.4
99.4
99.7
99.8
99.6
99.5
99.9
99.7
98.7
98.6
153.39
155.13
161.81
166.49
165.53
164.45
162.60
161.86
163.10
163.24
163.33
163.51
163.74
163.94
167.08
166.49
(163.34)*
2506
2521
2579
2618
2610
2601
2586
2579
2590
2591
2592
2593
2595
2597
2623
2618
(2592
+ 8
+ 8
t 5
? 8
2 3
k 2
? 3
_’ 2
k 2
? 2
? 3
2 3
2 3
+ 2
? 16
? 8
? 4)d
2.054
2.025
1.082
5.280
4.891
5.290
5.367
6.268
15.86
0.196
0.333
1.044
1.200
0.919
0.159
0.188
0.111
0.484
4.2
7.2
22.5
25.9
19.8
3.4
4.1
2.4
10.4
92.1
96.9
99.1
98.8
98.7
96.4
96.3
93.4
97.6
62.97
69.63
74.81
88.19
92.63
96.52
97.22
98.70
122.87
(89.43)d
1626
1736
1817
2011
2071
2122
2131
2150
2434
(2028
2 12
2 7
2 3
2 2
k 3
+ 11
k 7
1?1 7
k 2
k 4)d
0.15
0.16
0.29
0.06
0.06
0.06
0.06
0.05
0.02
2.856
5.530
12.68
7.878
3.575
2.491
4.843
13.96
11.92
13.45
22.15
17.84
17.55
13.43
0.0430
0.1199
0.1109
0.1376
0.3077
0.5503
0.7513
0.5895
0.4068
0.1189
0.1562
0.1855
0.0787
0.0036
1.2
3.4
3.1
3.9
8.6
15.6
21.2
16.5
11.4
3.3
4.3
5.2
2.2
0.1
73.4
88.6
93.9
96.9
98.3
98.5
99.1
99.0
99.0
98.6
98.6
98.4
98.0
51.5
64.04
51.01
63.91
60.85
56.86
39.97
61.90
89.60
91.73
88.85
120.17
95.85
98.08
77.48
(71.5)”
1459
1243
1457
1408
1343
1037
1425
1819
1846
1810
2173
1898
1813
1657
(1571
k 5
2 4
+ 2
t 2
t
1
5 1
IIZ 1
-e 1
2 2
2 2
-+ 2
2 2
t- 3
k 60
2 2)”
0.16
0.08
0.04
0.06
0.13
0.18
0.09
0.03
0.04
0.03
0.02
0.03
0.03
0.03
4oAr*/39ti
Age t la’
WCs
pm)
0.2763
0.0121
0.0002
0.0008
0.0009
0.0010
0.0004
0.0010
0.0011
0.0138
2
k
k
2
?
2
2
?
k
2
+
74
3
3
1
3
1
1
2
3
10
5)d
91-237 Muscovite (250- 150 pm)
(0.0014 g, J = 0.01963)
500
600
650
690
730
770
810
840
870
900
930
960
1000
1050
1100
1350
166.96
174.92
165.65
168.79
167.05
165.69
163.58
162.84
163.63
163.54
163.97
164.26
163.91
164.45
169.24
168.79
0.0460
0.0672
0.0131
0.0078
0.0051
0.0041
0.0033
0.0032
0.0017
0.0009
0.0021
0.0025
0.0005
0.0017
0.0075
0.0078
ScPvAm2 Hornblende (400-200
(0.0504 g, J = 0.02323)
600
700
800
900
950
975
1000
1050
1200
68.143
71.673
75.429
88.613
93.182
99.367
100.228
104.837
123.412
pm)
0.0183
0.0076
0.0024
0.0036
0.0040
0.0121
0.0126
0.0236
0.0099
9 l-205 Hornblende ( 180- 100 pm)
(0.0517 g. J = 0.01944)
600
700
750
800
850
890
930
970
1010
1040
1080
1150
1300
1500
87.069
57.351
67.406
62.414
57.663
40.491
62.260
89.572
91.870
89.235
119.95
96.162
89.751
149.03
0.0079
0.0237
0.0176
0.0088
0.0042
0.0027
0.0034
0.0072
0.0067
0.0083
0.0124
0.0106
0.0114
0.2490
Disturbed a.4r/39Ar release spectra in a gold mine
4661
Table 2 (Continued)
(X lo-’ cm3)
Vol j9Ar
Fraction
39Ar(%)
@A+
“Ar*P9A?
2.129
5.931
23.15
9.790
5.395
2.424
2.247
8.570
11.93
11.49
27.02
16.03
17.65
0.0485
0.1349
0.1040
0.1386
0.1781
0.4398
0.4909
0.4108
0.2815
0.1741
0.0894
0.1626
0.0112
1.8
5.1
3.9
5.2
6.7
16.5
18.5
15.5
10.5
6.5
3.3
6.1
0.4
64.1
90.8
96.0
98.8
99.2
99.7
99.5
99.3
99.5
98.8
98.8
98.6
86.4
73.14
74.65
94.81
88.23
86.06
89.34
95.32
116.14
129.07
128.71
140.77
118.10
117.42
(103.79)d
2.063
0.091
0.096
1.961
5.607
3.733
2.673
1.313
1.393
2.909
3.451
8.050
0.0321
0.1019
0.0529
0.0433
0.0469
0.0599
0.0957
0.1353
0.1111
0.1200
0.0639
0.0199
3.6
11.6
6.0
4.9
5.3
6.8
10.8
15.4
12.5
13.6
7.3
2.2
76.8
89.0
94.3
96.5
97.9
98.5
98.7
98.8
98.7
98.8
98.4
96.3
42.59
35.22
73.20
117.76
102.44
96.69
86.58
83.21
93.32
98.75
94.51
92.13
(8.98)”
Temp
(“0
Age ? la’
IUCa
91-233 Hornblende (180-100 pm)
(0.0380 g, J = 0.01938)
600
700
750
800
840
880
920
960
1000
1040
1100
1300
1500
113.95
81.842
97.089
88.655
86.375
89.464
95.672
116.26
128.54
129.15
139.73
118.34
134.09
0.1390
0.0271
0.0201
0.0065
0.0038
0.0016
0.0023
0.0054
0.0056
0.0085
0.0140
0.0127
0.0668
1593
1614
1881
1798
1770
1813
1888
2127
2261
2257
2374
2148
2140
(1989
2 14
2 4
+ 6
t 2
2 2
Ifr 2
t- 2
‘- 2
? 2
2 2
+ 2
2 2
2 17
+ 2)d
0.21
0.08
0.02
0.05
0.08
0.19
0.20
0.05
0.04
0.04
0.02
0.03
0.03
91-205 Biotite (150-90 pm)
(0.0019 g, J = 0.01952)
600
650
700
730
770
810
850
890
930
970
1010
1050
55.376
39.567
77.614
121.91
104.16
97.917
87.566
84.104
94.407
99.734
95.797
95.076
0.0440
0.0150
0.0153
0.0151
0.0089
0.0061
0.0046
0.0036
0.0043
0.0048
0.0061
0.0014
1092 2
944-c
1601 ?
2153 2
1982 +
1913 ?
1785 t
1741 +
1871 2
1938 +1886 +
1856?
(1751 2
6
6
4
3
4
3
2
2
2
3
5
9
4)d
a Calculated at 0°C 1 atm.
b 4oAr* = (“A&,,, - 40Ar~,-phe#‘Ar,otll.
’Quoted error does not include the error in the irradiation parameter J.
’ Calculated with weighting by mass 39Ar released per step.
using the decay constants of Steiger and Jager (1977), and final
errors are given as 2a including the error in J. KlCa plots are given
for amphibole only as the low contents of Ca in biotite and muscovite, coupled with the correction for 37Ar decay associated with the
long irradiation (and subsequent 4-6 month storage period) for
Archaean samples did not allow accurate determination of the KI
Ca ratio for these minerals.
Oxygen and hydrogen isotope analyses were carried out at the
University of Saskatchewan. Hydrogen isotope compositions were
determined using the uranium method of Godfrey ( 1962) as modified by Kyser and O’Neil ( 1984). Oxygen was extracted using the
BrFS technique described by Clayton and Mayeda ( 1963). Duplicate
analyses indicate a reproducibility of ? 0.2%0for oxygen values and
? 5%0 for hydrogen isotope values. 6l*O values of 9.6%0 for NBS28 quartz and SD values of -65%0 for NBS-30 biotite were obtained
using these techniques. S’*O and 6D values are reported relative to
V-SMOW.
5. RESULTS
5.1. “APPEAR Analyses
4oAr-39Ar data for all samples is given in Table 2 and
spectra and K/Ca plots (for homblendes only) are shown
in Fig. 4.
X1.1. Hydrothermal minerals
91-205 Hornblende. The spectrum from this hornblende
sample is more erratic than the other two homblendes analysed, although a general progression to older ages with
progressive gas release is evident. The first 60% of argon
released has apparent ages between 1000 and 1500 Ma, this
is followed by steps with ages around 1800- 1900 Ma (with
the exception of step 11). The total gas age of this sample
1572 ? 12 Ma. As with other hornblende samples K/Ca
values are higher in early stages of gas release (maximum
of 0.2) and have uniformly lower values around 0.02-0.03
for the remainder of gas release.
ScPvAm2 Hornblende. This sample shows increasing apparent step ages with progressive gas release, ranging from
1700 to 2434 Ma. The total gas age of this sample is 2028
t 14 Ma. K/Ca values are high for the first 35% of gas
released (maximum 0.4) and then have uniform values
around 0.1 for the next 60% of argon released. The final
step has a K/Ca ratio of 0.02.
91-233 Hornblende. This sample is similar to ScPvAm2
hornblende with gradually increasing step ages with progres-
A. J. R. Kent and T. C. McCuaig
4662
2500I-
91-205 Hornblende
2300
ScPvAm2 Hornblende
2300-
2100
2100-
1900
isoo-
1,
L
1700
ASP
(W
1500
1500
1
1300
1300
1100
900
I
1700.
1100 I
Q~Ook------
T-
I1.0
0.2
0.4
0.6
0.6
1.0
0.2
Fraction3sAr released
0.4
0.6
0.8
1.0
Fraction3sAr reteased
wca
p-j
_
o.“o~o
bction
2500
2500-
91-233 Hornblende
91-205 Biotite
2300
2300.
rj , , , , , , ,
0.0
0.2
0.4
Fractkm
0.0
39~r releared
0.2
0.4
0.6
I
,
0.6
Fraction3gAr released
39Al reb&sed
0.6
0.6
Frtiton 39AJ retensed
Fig. 4. “Ar- ‘“AT spectra and K/Ca plot from hydrothermal
analysed from the Scotia gold mine.
sive gas release. Apparent ages show a general trend of
increasing age with progressive gas release from a minimum
of ca. 1600 Ma to a maximum age of 2373 Ma. The total
gas age of this sample is 1989 t 12 Ma. K/Ca values are
generally higher (maximum of 0.2) in the first 60% of argon
biotite and hornblende
and pegmatite
muscovite
samples
released and uniformly low (0.02-0.04)
for the final 40%
of argon released.
91-205 Biotite. Argon released in the early stages of gas
release ( 15%) has apparent ages in the order of 1000 Ma.
and the remaining steps form a saddle shaped spectrum. The
Disturbed
4oAr/39Ar release spectra
in a gold mine
4663
2620
2590
2590.
2560
256Oj
2530
2530.
2615f14Ma
1
25001
0.0
.
m
0.2
m
I
0.4
.
.
0.6
.
.
0.9
25ooJ
0.0
.
*
.
0.2
I
.
0.4
,
,
0.6
*
,
0.6
,
3
FractionSAI released
Fraction38~ released
Fig. 4. (Continued)
total gas age for this sample is 175 1 + 12 Ma, and the saddle
shaped portion of the spectrum (steps 4- 12) has an apparent
age of 1878 2 12 Ma. The saddle shape of the spectrum is
similar to that shown by spectra from biotite with interlayered chlorite alteration in Lo and Onstott ( 1989).
5.1.2. Pegmutites
In contrast to hydrothermal minerals, both pegmatite muscovite samples analysed exhibit 40Ar-3gAr spectra with well
developed plateau-like segments which contain >90% of
argon released. For sample 91-237, a plateau-like segment
between steps 3 and 16, comprising 93% of total 39Ar released, corresponds to an age of 2593 + 12 Ma. Muscovite
from sample ScPeg8 exhibits a plateau-like segment between
steps 3 and 19 comprising 97% of 3gAr released, and this
corresponds to an age of 2615 + 14 Ma. Within analytical
error (at 20) both ages are indistinguishable.
5.2. Stable Isotope Analyses
Stable isotope results for hydrothermal minerals from the
Scotia mine are shown in Table 3. Hydrothermal minerals
Table 3. Oxygen
Scotia mine.
and hydrogen
isotope data for minerals
6’*0
Temperature”
from the
Water
yield
Sample and mineral
6”O
6D
quartz
(“C)
(wt%)
&Peg8 muscovite
ScPvAm2 hornblende
91-205 hornblende
Sc163-66A amphibole
SC163-66B amphibole
SC358-99 amphibole
SC358-33 amphibole
SC358-33 diopside
SC358-4OB diopside
SC358-55 diopside
6.9
6.6
6.9
6.8
6.8
-58
-70
11.5
510
2.0
1.2
11.6
11.1
510
560
6.4
6.3
6.4
6.4
-65
-62
-65
1.3
1.9
11.0
500
All values reported in %o relative to V-SMOW. ’ Temperatures calculated using quartz-hornblende fractionation factor of Bottinga and Javoy
(1973) and quartz-clinopyroxene
fractionation factor of Clayton and
Keiffer (1991).
show a limited range in 6 “0 values (diopside = 6.3-6.4%0;
hornblende = 6.4-6.9%0; quartz = 11 .O- 11.6%0). Fractionations between quartz-hornblende and quartz-clinopyroxene
are consistent with formation of gold-related alteration assemblages at temperatures between 500°C and 560°C. These
calculated temperatures corroborate those estimated from
phase equilibria considerations, which also indicate formation of Scotia gold-related hydrothermal alteration assemblages at temperatures > 500°C (McCuaig et al., 1993).
Hydrogen isotopic compositions of hydrothermal amphiboles also show little variation, with 6D values ranging from
-70 to -61%0. Using a temperature range of 500-56O”C,
the quartz-water fractionation curve of Clayton et al. (1972)
for oxygen, and the hornblende-water fractionation curve of
Suzuoki and Epstein (1976) for hydrogen, yields 6”O and
SD values for the mineralising fluid of 8.8 to 9.%0, and -29
to -43%0, respectively. 6 ‘*OB,idand SDnuidestimates for other
deposits within the Yilgam Craton are similar to these values, although less data is available for hydrogen isotopes
(Golding et al., 1992). Calculated 6 180~uidvalues for deposits at Norseman and Kambalda, to the north of the Scotia
deposit, range between 5 and 10%0 and SDauid for the Victory-Defiance deposit at Kambalda is estimated at -30
t 12%0. Although the data available for the closest deposit
to the Scotia mine (the Princess Royal deposit at Norseman)
does suggest that this deposit has a different GDBuidvalue
than Scotia (- 10 + 4%0), it is unsure as to the degree
that minerals from Princess Royal are effected by the later
metamorphic effects of a large mafic intrusion in this mine
(Golding et al., 1992).
6. DISCUSSION
6.1. Tbe Age of Pegmatite
Mineraiisation
Intrusion
and Gold
The 2593 + 12 and 2615 ? 14 Ma ages of plateau-like
segments from pegmatite muscovites are within error of the
2620 + 36 Ma estimate of the age of sample 91-237 from
Sm-Nd (garnet-albite) analysis (Kent et al., 1996). Intrusion
of pegmatite dykes of comparable composition to those at
Scotia occurred elsewhere within the southern Yilgam Craton
at ca. 2630 Ma, and on a regional scale, post-mineralisation
A. J. R. Kent and T. C. McCuaig
4664
pegmatite intrusion at this time was contemporaneous
with an
episode of partial melting, granitoid intrusion. and metamorphism within the lower-middle crust (Kent et al., 1996).
Field relations such as those detailed in Fig. 2 unambiguously demonstrate
that pegmatite intrusion occurred after
gold mineralisation
at the Scotia mine, thus the ca. 26002620 Ma 40Ar- j9Ar from pegmatite samples provide a minimum estimate of the timing of mineralisation. Gold mineralisation at Scotia prior to ca. 2600-2620 Ma is also consistent
with the deposit forming during a widespread gold mineralisation event in the Yilgarn Craton at ca. 2630 Ma (e.g..
Groves, 1993a.b; Kent et al., 1996).
2500
I
6.2. Interpretation
of QAr-39Ar Spectra from
Hydrothermal Biotite and Hornblende
‘0Ar-39Ar spectra from hydrothermal
biotite and homblende are characterised
by a lack of plateau-like features
and erratic step ages in the earlier stages of gas release (Fig.
4). Isotope correlation plots for the hydrothermal
minerals
analysed are also erratic, and this, coupled with the radiogenie composition
of argon released from these samples.
precludes these from providing useful information.
Several factors may contribute to the unusual shape of the
spectra from biotite and hornblende samples analysed: ( 1 )
The general increase of step ages with progressive degassing
in hornblende samples may be due, at least in part, to the
presence of an argon concentration gradient. However, given
the erratic nature of the spectra and the evidence that argon
release during hornblende step heating may not reflect internal argon concentration variations (e.g., Hanson et al., 1975:
Gaber et al., 1988; Wartho et al., 1991; Lee et al., 1991; Lee.
1993 ) , no attempt has been made here to model hornblende
spectra in terms of diffusive argon loss. (2) Degassing of
intergrown
phases or exsolution
lamellae with differing
argon release behaviour, coupled with recoil of ‘9Ar during
sample irradiation, is a documented cause of erratic spectra
in biotite and hornblende (e.g., Harrison and Fitz Gerald.
1986; Lo and Onstott, 1989). This phenomenon
can also
significantly lower apparent closure temperatures (Harrison
and Fitz Gerald, 1986; Lo and Onstott, 1989; Baldwin et al.,
1990). Backscattered electron images of biotite from sample
9 l-205 reveal numerous intergrowths of chlorite along cleavage planes on the 1-2 pm scale (Fig. 3 ) . The saddle-shaped
spectrum exhibited by biotite from sample 91-205 is also
similar to spectra observed in other samples where biotite is
known to be finely intergrown with chlorite (e.g., Lo and
Onstott, 1989)) including samples from other Australian Archaean gold deposits (A. J. R. Kent unpubl. data. 1994 ),
Likewise backscattered
images of hornblende from sample
91-233 and 91-205 show that small (generally
1 - 10 pm)
domains of actinolite or actinolitic hornblende occur within
larger magnesio hornblende dominated grains (Fig. 3 ). (3 )
Variable amounts of contaminating K-bearing phases in bulk
mineral separates can also effect argon systematics,
with
observed spectra and apparent ages deriving from superimposed degassing of the primary phase and the contaminant
(e.g., Rex et al., 1993). For this study contamination
of
hornblende separates by biotite is most important, due to the
difference in K contents of these minerals (Table 1) and the
0
Field of amphibole
_.--- K/Ca values
0.00
0.05
0.10
A.
0
1000 -a
0.15
0.20
0.25
0.30
K/Ca
30
B.
n
25 t
/
91-205
Hornblende
91-233
Hornblende
0
4
400
64nJ
800
loo0
12W
Temperature W)
1400
1600
Fig. 5. (a) Variation of apparent ages with K/Ca for gas release
steps for hornblende samples from the Scotia gold mine. For each
sample steps with low K/Ca tend to have older apparent ages. (b)
Fraction of released “Ar with progressive heating for biotite and
hornblende from sample 91-205 and hornblende 91-233.
common intergrowth of these phases in the analysed samples. Despite every precaution it is probable that some degree
of contamination occurs in bulk mineral separates, and even
small degrees of contamination of a hornblende mineral separate by biotite may have a marked effect on apparent
ages during the lower temperature
stages of gas release
(< - 1000°C; Rex et al., 1993; Wartho, 1995). For the
hornblende samples analysed in this study, high K/Ca values
in the earlier, lower temperature, stages of gas release generally coincide with relatively young apparent ages (Fig. 5a,
Table 2) and are probably due to degassing of contaminant
biotite. However, as shown in Fig. 5b, at temperatures above
Disturbed 4oAr/39Arrelease spectra in a gold mine
- 1000°C outgassing of biotite is largely complete and biotite-derived argon should not contribute to the apparent ages
of argon released from hornblende above this temperature
(also similar to the findings of Rex et al., 1993). K/Ca ratios
calculated for argon released from hornblende samples at
temperatures above 1000°C are in the range of primary values for hornblendes measured by electron probe (Fig. Sa),
further indicating that minimal biotite-derived argon has contributed to the argon released at these temperatures. For this
study, contamination of hornblende by biotite may be most
important for sample ScPvAm2, as the grainsize of analysed
crystals is larger than that for the other hornblende samples
analysed (400-200 pm compared to 200- 180 pm for 91205 and 91-233), and the degassing temperature schedule
provides minimal resolution of the gas released above
1000°C (Table 2).
For sample 91-205 the > 1000°C fraction consists of five
steps comprising 27% of the total argon released and has a
mean age of 2246 t 12 Ma and K/Ca ratio of 0.03. For
sample 91-233 the mean age of this fraction (6 steps and
27% of released 39Ar) is 1745 + 11 Ma with a K/Ca of
0.03. For ScPvAm2 the >lOOO”C fraction consists of only
one step (10% of the total 39Ar released) with an age of
2434 & 14 Ma and K/Ca of 0.02.
6.3. Argon Loss in Hydrothermal
Minerals
It is important to note that, irrespective of the causes of the
erratic spectra, all hydrothermal minerals analysed appear to
have suffered a considerable, but varying, degree of argon
loss. Biotite from sample 91-205 is almost 900 million years
younger than the ca. 2600-2620 Ma minimum age of gold
mineralisation and hydrothermal alteration. For homblendes
tbe ages of argon fractions released at temperatures above
1000°C are significantly younger (- 900-200 million years)
than the minimum age of gold mineralisation. In addition to
argon loss, the ages for the >lOOo”C fractions from the
hornblende samples are also highly variable, with ages covering a range of almost 700 million years.
The close spatial association between of hydrothermal
lodes, containing biotite and hornblende that have experienced argon loss, and pegmatites dykes containing apparently undisturbed muscovite provides an insight into the nature of the processes responsible for argon loss in the hydrothermal minerals. The similarity between tbe ages (either
plateau or integrated ages) of muscovites from pegmatite
dykes, and the Sm-Nd garnet-albite age of the pegmatite
dykes from Scotia and from elsewhere in the southern Yilgam Craton (Kent et al., 1996) suggests that little, if any,
argon loss has occurred from muscovite within the pegmatites dykes. From this it can be inferred that these dykes
cooled quickly through their closure temperatures and were
not subsequently reheated to a temperature close to the
blocking temperature for a time period sufficient to promote
any significant degree of argon loss. Thus argon loss in
hydrothermal minerals occurred at ambient temperatures that
were below the closure and blocking temperature of argon
diffusion in pegmatite muscovites.
Most estimates of muscovite closure temperature are in
the range - 275-400°C and are largely based on empirical
4665
studies of the age of muscovite with respect to other indicators of cooling history, such as other mineral ages (e.g.,
Jager, 1979; Blankenburg et al.) 1989), or fluid inclusion
temperatures (Snee et al., 1988). Similarly calculations by
Lister and Baldwin ( 1996). using re-evaluated data of Robbins ( 1972) for the diffusion parameters of argon in muscovite, an infinite slab model for argon diffusion, and a range of
thermal histories and diffusion domain sizes (see discussion
below), suggest that the closure temperature of argon in
muscovite is between about 300-420°C (for pressures between 0 and 5 kbar-commensurate with pressure estimates
for the Scotia deposits; McCuaig et al., 1993). Lister and
Baldwin ( 1996) also estimate the blocking temperature of
muscovite to be in the order of 255-285°C from the basis
of numerical calculations, again using a range of thermal
histories.
Empirical and experimental studies suggest that for a
given thermal history the closure temperature of biotite is
slightly lower than that of muscovite (e.g., McDougall and
Harrison, 1988; Lister and Baldwin, 1996). Thus, at the
Scotia mine, biotite could have experienced thermally-driven
diffusive argon loss without nearby muscovite being disturbed if the Scotia mine region was either: ( 1) reheated to
a temperature above that of biotite but below muscovite
blocking temperature (for a range of thermal histories the
blocking temperature of biotite is estimated at - 230-255°C
by Lister and Baldwin, 1996), or (2) the thermal history
of the Scotia region involved extremely protracted cooling
(albeit sufficient to generate a ca. 900 Ma difference in
apparent ages of muscovite and biotite) between the muscovite and biotite closure temperatures (biotite closure temperature is generally thought to be in the range 300 5 50°C;
McDougall and Harrison, 1988).
However, such simple models of thermally-driven diffusive argon loss based on published closure and blocking
temperature estimates cannot explain the observed argon loss
in hydrothermal hornblende from the Scotia deposit. Estimates of the closure and blocking temperature of homblendes are 500 2 50°C (e.g., Harrison, 1981) and 400430°C (Lister and Baldwin, 1996), respectively. Given that
both the estimated blocking and closure temperatures of
hornblende are in excess of that of muscovite, it is difficult
to explain the observed argon loss in hornblende from the
Scotia mine via a model of diffusive argon loss during slow
cooling and/or post-mineralisation reheating. Any thermal
disturbance or period of protracted cooling capable of resulting in argon loss from hornblende would also be expected
to completely reset tbe adjacent muscovite.
In light of this, two explanations for argon loss at the
Scotia mine are discussed below. Tbe first involves variations in the blocking and closure temperatures of the analysed minerals away from estimated values, and the second
involves argon loss during post-mineralisation
fluid flow
along ore-hosting structures.
6.3.1. Variations in closure and/or blocking temperatures
One explanation for tbe contrasting argon systematics of
muscovite and hornblende at the Scotia mine could be that
(assuming argon loss in hornblende samples was a volume-
4666
A. J. R. Kent and T. C. McCuaig
diffusion driven process) the actual closure and/or blocking
temperatures for muscovite and hornblende differ considerably from the estimates given above for these minerals. For
example, if the closure temperature of hydrothermal
homblendes were sufficiently less than 500 t 5o”C, and/or the
closure temperature
of pegmatite muscovites
were sufficiently greater than 350 +- 50°C then, given a suitable thermal history, diffusive argon loss in hornblende could potentially occur during slow cooling, without argon loss in muscovite.
Variations in closure temperature are certainly possible.
For example, in amphiboles it is well documented that exsolution can result in lower apparent closure temperatures in
metamorphic homblendes (Harrison and Fitz Gerald, 1986;
Baldwin et al., 1990). The experimental results of Baldwin
et al. ( 1990) suggest that the depression of closure temperature in highly exsolved metamorphic homblendes
is on the
order of lOO- 150°C (these authors calculate closure temperature between 360 and 435°C for their samples of metamorphic hornblende).
In the analysed hornblende samples from
Scotia, the presence of structural defects, small domains of
actinolitic hornblende observed within larger magnesio homblende grains, and the presence of small inclusions of impurities, such as chlorite or biotite, may similarly have lowered
the closure temperature of this mineral (Baldwin et al., 1990:
Lee, 1993).
Another alternative is that the closure and/or blocking
temperature of the pegmatite muscovite analysed was significantly higher than that of the estimated values. In contrast
to the infinite plane geometry used by Lister and Baldwin
(1996) for calculations
of the muscovite
closure and
blocking temperatures, Hames and Bowring ( 1994) recommend a cylindrical argon diffusion geometry for this mineral.
Further, these authors, on the basis of recognised argon concentration gradients in large muscovite crystals, suggest that
the effective diffusion dimension in such grains may be controlled by the physical grainsize of the mineral. If this is the
case for muscovite grains the size of those used in this study
(0.5-2.5 cm), then muscovite closure and blocking temperatures could be significantly higher than estimates for small
grains. This is illustrated in Fig. 6 where the calculated closure temperature of muscovite is plotted against the effective
diffusion dimension for a range of cooling rates (note that
the calculations for this plot use a cylindrical diffusion geometry and the diffusion parameters
for muscovite
recommended by Hames and Bowring, 1994 from re-appraisal of
the data of Robbins, 1972). However, although high closure
temperatures
may be possible, it is by no means certain
that the effective diffusion dimension is controlled by the
physical grain size in such large muscovite crystals. Argon
concentration
gradients have been observed in grains only
up to a few millimetres in diameter, not the 50-250 mm
size of the muscovite crystals analysed in this study. Further,
Lister and Baldwin ( 1996) argue against the interpretation
that the argon concentration
gradients observed represent
diffusion within a single diffusional domain. Also, the authors of this study are unaware of any empirical estimate for
muscovite which independently
assess the closure temperature as significantly in excess of 400°C.
In summation, it is considered unlikely that the observed
SW-
dT/dt
----zw 1
.(x)1
,,,,,,
,,,,;
.Ol
;
.l
= lO”C/Ma
dT/dt=S’C/Ma
_,_,“‘“;t’““,_
1
a (an)
Fig. 6. Variation of calculated closure temperature in muscovite
with increasing effective diffusion dimension calculated from the
formulation of Dodson ( 1973) using cylindrical geometry and diffusivity data from Robbins ( 1972) and Hames and Bowring ( 1994).
argon loss in hornblende in solely the result of thermally
driven argon loss via volume diffusion. Further, even if it
were feasible that a combination of lowering of hornblende
and elevation of muscovite closure/blocking
temperatures
could produce argon loss in hornblende and not pegmatite
muscovite, it is difficult to explain the differences
in the
apparent ages of hornblende samples by such a model. The
apparent ages of gas released from hornblende samples at
temperatures above 1000°C differ by almost 700 Ma. If the
closure or blocking temperature, either during slow cooling,
or during a reheating event, provided the only control on
argon loss then similar ages for the different hornblende
samples would be expected, unless there were also dramatic
variations in closure/blocking
temperatures and/or thermal
history between samples from different locations. It is difficult to imagine that variations in the thermal history or closure/blocking
temperature would be sufficient over the few
hundred metres range over which samples were collected to
generate ages that differ by up to 700 million years.
6.3.2. Argon
loss during fluid jiow
An alternative explanation for the disturbed argon systematics in hydrothermal minerals from the Scotia mine is that
argon loss occurred by non-volume diffusion processes. As
discussed in the introduction to this paper, there is considerable evidence that argon loss during laboratory step heating
in hornblende and biotite is the result of nonvolume-diffusion processes such as dehydroxylation
and in vacua phase
transitions. Likewise other studies have suggested that, in
some situations, argon loss over geological time may also
be the product of nonvolume-diffusion
processes, such as
mineral-fluid
interaction
and/or chemical re-equilibration
(e.g., Miller et al., 1991; Wartho, 1995).
The suggestion, therefore, is that argon loss in hydrothermal hornblende, and possibly also biotite, at the Scotia mine
Disturbed 4oAr/39Ar release spectra in a gold mine
occurred during movement of a retrograde fluid along the
ore-hosting structures. Argon loss was associated with mineral-fluid interaction (and possibly also concomitant chemical re-equilibration) during this retrograde event (or events).
Further, it appears that argon loss in hornblende occurred at
temperatures below that of the closure and blocking temperatures of pegmatite muscovite, which if these are similar to
published estimates, suggests that argon loss in hornblende
occurred at temperatures below about 250-300°C.
This scenario is consistent with the abundance of field
and mineralogical evidence that show that the deposit has
been the site of major retrograde (post-mineralisation) fluid
flow. In all parts of the mine numerous brittle fractures crosscut and often displace both the ore-hosting structures and
post-ore pegmatites (e.g., Fig. 2; Thomas et al., 1990; McCuaig et al., 1993). The brittle nature of deformation associated with fractures and retrograde assemblages (quartz-chlorite-albite-calcite-sericite; albite-prehnite; gypsum) infer that
fractures formed at pressure-temperature conditions substantially less than those at which gold mineralisation occurred
(McCuaig et al., 1993). This is consistent with the suggestion that argon loss in hornblende occurred at relatively low
temperatures. In some areas small flakes (91-5
pm) of
retrogressive minerals (predominantly
chlorite) are also
found along cleavage planes, grain boundaries, and in microfractures in hydrothermal minerals (e.g., Fig. 3; McCuaig et
al., 1993) suggesting that retrograde fluid flow also occurred
along grain boundaries within ore zones. Further, as documented already, hornblende samples from the Scotia mine
show marked chemical variation, and although this is largely
interpreted to have occurred during gold mineralisation, the
possibility that some of the chemical variations in homblende are the result of post-crystallisation chemical re-equilibration cannot be ruled out.
The actual mechanisms of argon loss during fluid activity
or chemical re-equilibration, remain less well understood
than argon loss via volume diffusion. Wartho ( 1995) suggested that argon loss in amphiboles from the Rameka gabbro in New Zealand, and from Bayan Obo in Inner Mongolia
were most likely related to partial chemical re-equilibration
during post-ctystallisation
thermal events. Miller et al.
( 1991) argued that reset K-Ar ages in homblendes from the
Connemara region of Western Ireland were related to excess
structural water and anomalous 6D values, interpreting this
to indicate that exchange between the hornblende and fluid
was the most probable cause of argon loss.
Unlike argon loss via volume diffusion, argon loss associated with mineral-fluid interaction can potentially explain the
observed differences in argon systematics of hydrothermal
minerals compared to pegmatite muscovite at the Scotia
mine. If argon loss depends on fluid interaction, then the
thermal history of a particular sample becomes subordinate
with respect to the hydrothermal history. The degree of mineral re-equilibration, and thus also presumably the degree of
argon loss, will then depend on parameters such as the local
permeability, fluid P-T-X conditions, fluid/rock ratios, and
the extent of fluid-mineral reaction. Compared to thermal
history variations, fluid parameters can vary on small spatial
and temporal scales. Thus the large differences in ages of
the > 1000°C argon fractions from hornblende samples from
4667
Scotia could be explained by differing degrees of fluid-mineral interaction, controlled by localised variations in fluid
parameters. The apparent lack of argon loss in muscovite
may be a function of the reduced permeability of the pegmatites in comparison to the relatively more fractured and displaced ore zones. Field observations support this. Compared
to ore zones, pegmatite are less fractured and show no indication of grain-boundary retrograde phases evident in the minerals in lode zones. Where brittle fractures do occur within
pegmatites, they are more widely spaced and are marked
by thin zones of cataclastic deformation, with retrograde
minerals restricted entirely to these spaced fractures. This
suggests that fluid flow within pegmatites during the fracturing was restricted entirely to these fractures.
If fluid-mineral interaction is the cause of the observed
argon loss in hydrothermal minerals then, unlike the study
of Miller et al. ( 199 1) , this appears to have not significantly
affected stable isotope systematics. Both the hydrogen and
oxygen isotope values of hydrothermal minerals from Scotia
are consistent with those estimated for nearby gold deposits
(albeit with a small dataset for 6D values) and temperatures
calculated from oxygen isotope fractionation of 500-560°C
(Table 3) are the same as the estimates of the temperature
of mineralisation from mineral equilibria (McCuaig et al.,
1993). Further careful work and an enlarged regional database is required to investigate whether hydrogen isotopes at
Scotia have been perturbed during argon loss (cf. Kerrich
and Cassidy, 1994). Fluid-mineral interaction may have occurred at temperatures less than those required for re-equilibration of oxygen isotopes, however, it is also plausible that
any retrograde fluid involved in mineral-fluid equilibration
and argon loss had already attained oxygen and hydrogen
isotope equilibrium with hydrothermal minerals during
movement along the ore structures from deeper crustal levels. Also, microscale investigations may be required to recognise disturbances of oxygen and hydrogen isotopes if mineral-fluid re-equilibration occurred within limited spatial domains.
Although it is difficult to constrain the precise timing of
argon loss with the available data, the range of apparent
ages in hornblende and biotite samples is consistent with
movement of retrograde fluids throughout the Scotia deposit
in response to Proterozoic erogenic activity along the adjacent margins of the Yilgam Craton. Scotia is situated less
than 100 km from the southern and eastern boundary of the
Yilgam Craton (Fig. 1) and is within a zone of deformation,
as determined from aeromagnetic data, which is attributed
to Proterozoic interaction between the Yilgam Craton and
the Albany-Fraser Province (Whitaker, 1990). Several major erogenic episodes, involving high grade metamorphism
and granitoid emplacement, occurred within the Albany-Fraser Province between ca., 1900- 1100 Ma (Gee, 1979). This
activity may have caused successive pulses of fluid activity
along Archaean structures adjacent to the cratonic margin,
including those which host gold mineralisation at Scotia, in
a way analogous to that documented for Archaean faults in
the Superior Province (Kerrich, 1994; Powell et al., 1995).
In the Yilgarn Craton these retrogressive events are also
recorded in other isotopic systems. For example, lead isotope
results from ore sulfides associated with gold mineralisation
A. J. R. Kent and T. C. McCuaig
4668
in the Norseman and Kalgoorlie Terranes record at least two
Proterozoic events: one at ca. 2000 Ma and a later event at
ca. 1100 Ma, that have altered lead isotopic compositions
of gold-related
galena (Perring and McNaughton,
1990).
Craton-wide
intrusion of mafic dykes at ca. 2400 Ma
(Fletcher et al., 1987), including some in the Scotia mine
vicinity, may also be related to some argon loss events.
6.3.3. Ramfications for interpretation of 10Ar-.‘yArresults
The possibility that argon loss may be attributed to postmineralisation
fluid movement through the Scotia deposit
has important ramifications for the interpretation
of results
from 40Ar- 39Ar analysis of minerals that may have interacted
with fluids after formation. Where isotopic resetting of this
style has occurred the criteria required to assess the validity
of isotopic ages may differ from those used to detect thermally-driven
argon loss. The thermal history of a sample
becomes largely secondary in favour of the hydrothermal
history of a sample, and factors such as permeability, fluid/
rock ratios, fluid P-T-X conditions, and crystal microstructures (to facilitate fluid-mineral exchange reactions ) determine the degree of isotopic disturbance (e.g., Miller et al.,
199 1) As these parameters can vary on short length and
timescales, it is envisaged that argon loss via this process can
be much more variable than that driven by thermal history
variations, which are generally regional in scale. This is
especially germane to the studies of hydrothermal ore deposits as, by definition, many ore deposits of this type are located
within structures that act as locii for the passage of hydrothermal fluids. Many mineral deposits, such as Archaean
gold deposits, are associated with major crustal structures
that have experienced long post-mineralisation
fluid movement histories (e.g., Kerrich and Cassidy, 1994; Powell et
al., 1995). In this situation there is considerable opportunity
for argon loss to have occurred within hydrothermal minerals. Interpretations
from previous studies on Archaean gold
deposits that have entertained the possibility of argon loss
only under purely thermal regimes (e.g., Hanes et al., 1992:
Zweng et al., 1993) will require reassessment
in light of the
results outlined herein. Although some studies have recommended the use of stable isotope systematics to recognise
activity of retrograde fluids (e.g., Miller et al., 1991; Kerrich
and Cassidy, 1994). the conclusions
of this study (albeit
limited by restricted data for comparison) are that argon loss
can occur without obvious disturbance of the bulk oxygen
(and possibly also hydrogen) isotope systematics of coexisting silicates.
7. CONCLUSIONS
Argon loss in hydrothermal hornblende (and possibly also
biotite) at the Scotia gold mine, Western Australia is interpreted to have occurred during fluid/mineral interaction associated with post-gold movement of a retrograde fluid along
ore-hosting structures. Explanations for the observed argon
systematics whereby argon loss occurs via volume diffusion
and is controlled by the thermal history of the Scotia region
(such as during protracted cooling or post-mineralisation
reheating) are considered improbable.
If argon loss is controlled by fluid/mineral
interactions,
then factors such as fluid/rock ratio, fluid P-T-X conditions,
permeability, and mineral microstructures
may be more important in governing
40Ar-39Ar systematics
than factors
known to influence volume diffusion. Thus this study recommends caution for assessing jOAr- 39Ar behaviour in environments, such as many hydrothermal ore deposits, where fluids
may have interacted with minerals after crystallisation. Interpretations based exclusively on the thermal, rather than hydrothermal, history of analysed samples may be insufficient.
Although the example outlined herein is an extreme case,
more subtle forms of this phenomenon
may be difficult to
recognise.
Acknowledgments-AJRK
acknowledges
the receipt of Australian
Postgraduate
Research and ANUTECH
Supplementary
Scholarships, and TCM acknowledges
receipt of NSERC and University of
Saskatchewan
Postgraduate
Scholarships.
The Australian Institute
of Nuclear Science and Engineering
(AINSE) provided financial
assistance
for sample irradiation
in Australia. J. Fedorowich
is
thanked for argon analysis of samples at Queen’s University. Western Mining Corporation
and Central Norseman Gold Corporation
provided access to the Scotia mine, logistical support during field
work. and financial assistance for analytical work. K. Johnson, S.
Peters, and R. Waugh are thanked for assistance and inspiring discussions during field work. P. Carpenter is thanked for help with electron
microprobe and SEM analyses. Additional financial and logistical
support by BHP-Utah, Normandy Poisedon, Pancontinental
Mining,
Placer Exploration,
and Western Mining Corporation
to AJRK is
also gratefully acknowledged.
Discussions with J. Eiler and reviews
by I. McDougall,
T. Spell. J. Lee, K. Cassidy, R. Kerrich, K.V.
Hodges, and T.K. Kyser significantly improved the manuscript.
Editoriul
handling:
T. K. Kyser
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