Contrib Mineral Petrol (1997) 129: 120±142
Ó Springer-Verlag 1997
C. O'Reilly á GRT. Jenkin á M. Feely
DHM. Alderton á AE. Fallick
A ¯uid inclusion and stable isotope study of 200 Ma
of ¯uid evolution in the Galway Granite, Connemara, Ireland
Received: 3 April 1996 / Accepted: 5 May 1997
Abstract Fluid inclusions in granite quartz and three
generations of veins indicate that three ¯uids have affected the Caledonian Galway Granite. These ¯uids
were examined by petrography, microthermometry,
chlorite thermometry, ¯uid chemistry and stable isotope
studies. The earliest ¯uid was a H2O-CO2-NaCl ¯uid of
moderate salinity (4±10 wt% NaCl eq.) that deposited
late-magmatic molybdenite mineralised quartz veins
(V1) and formed the earliest secondary inclusions in
granite quartz. This ¯uid is more abundant in the west of
the batholith, corresponding to a decrease in emplacement depth. Within veins, and to the east, this ¯uid was
trapped homogeneously, but in granite quartz in the
west it unmixed at 305±390 °C and 0.7±1.8 kbar. Homogeneous quartz d18O across the batholith (9.5 ‹
0.4& n = 12) suggests V1 precipitation at high temperatures (perhaps 600 °C) and pressures (1±3 kbar)
from magmatic ¯uids. Microthermometric data for V1
indicate lower temperatures, suggesting inclusion volumes re-equilibrated during cooling. The second ¯uid
was a H2O-NaCl-KCl, low-moderate salinity (0±10 wt%
NaCl eq.), moderate temperature (270±340 °C), high dD
()18 ‹ 2&), low d18O (0.5±2.0&) ¯uid of meteoric
origin. This ¯uid penetrated the batholith via quartz
C. O'Reilly á M. Feely1 (&)
Department of Geology, University College Galway, Ireland
G.R.T. Jenkin á A.E. Fallick
Isotope Geosciences Unit, Scottish Universities Research
and Reactor Centre, Rankine Avenue,
Scottish Enterprise Technology Park,
East Kilbride, G75 0QF, Scotland, UK
D.H.M. Alderton
Royal Holloway, University of London,
Egham, Surrey, England, UK
1
Present address:
Department of Geology and Geophysics,
University of Edinburgh, West Mains Road,
Edinburgh, EH9 3JW, Scotland, UK
Editorial responsibility: I. Parsons
veins (V2) which in®ll faults active during post-consolidation uplift of the batholith. It forms the most common
inclusion type in granite quartz throughout the batholith
and is responsible for widespread retrograde alteration
involving chloritization of biotite and hornblende,
sericitization and saussuritization of plagioclase, and
reddening of K-feldspar. The salinity was generated by
¯uid-rock interactions within the granite. Within granite
quartz this ¯uid was trapped at 0.5±2.3 kbar, having
become overpressured. This ¯uid probably in®ltrated
the Granite in a meteoric-convection system during
cooling after intrusion, but a later age cannot be ruled
out. The ®nal ¯uid to enter the Granite and its host
rocks was a H2O-NaCl-CaCl2-KCl ¯uid with variable
salinity (8±28 wt% NaCl eq.), temperature (125±
205 °C), dD ()17 to )45&), d18O ()3 to + 1.2&),
d13CCO2 ()19 to 0&) and d34Ssulphate (13±23&) that
deposited veins containing quartz, ¯uorite, calcite, barite, galena, chalcopyrite sphalerite and pyrite (V3).
Correlations of salinity, temperature, dD and d18O are
interpreted as the result of mixing of two ¯uid endmembers, one a high-dD ()17 to )8&), moderate-d18O
(1.2±2.5&), high-d13CCO2 (> )4&), low-d34Ssulphate
(13&), high-temperature (205±230 °C), moderate-salinity (8±12 wt% NaCl eq.) ¯uid, the other a low-dD ()61
to )45&), low-d18O ()5.4 to )3&), low-d13C (<)10&),
high-d34Ssulphate (20±23&) low-temperature (80±125 °C),
high-salinity (21±28 wt% NaCl eq.) ¯uid. Geochronological evidence suggests V3 veins are late Triassic; the
high-dD end-member is interpreted as a contemporaneous surface ¯uid, probably mixed meteoric water and
evaporated seawater and/or dissolved evaporites,
whereas the low-dD end-member is interpreted as a
basinal brine derived from the adjacent Carboniferous
sequence. This study demonstrates that the Galway
Granite was a locus for repeated ¯uid events for a variety of reasons; from expulsion of magmatic ¯uids
during the ®nal stages of crystallisation, through a meteoric convection system, probably driven by waning
magmatic heat, to much later mineralisation, concentrated in its vicinity due to thermal, tectonic and
121
compositional properties of granite batholiths which
encourage mineralisation long after magmatic heat has
abated.
Introduction
Granite plutons intruded into the upper crust represent
thermal anomalies which are predicted to cause ¯uid
convection (Norton and Knight 1977). This provides a
ready explanation for the widespread alteration, veining
and mineralisation frequently observed in granites and
their host rocks (e.g. Criss and Taylor 1986). However,
¯uid circulation driven by magmatic heat is expected
within only a short period of time (<106 years) following
intrusion (Cathles 1981; Criss and Taylor 1986) and,
while many examples of alteration, veining and mineralisation have ages indistinguishable from the age of the
associated intrusion (e.g. Skinner 1979), an increasing
number of examples occur that are spatially related to
granites, yet signi®cantly younger (Missouri: Wenner
and Taylor 1976; Cornubia: Jackson et al. 1982; Chesley
et al. 1993; Massif Central: Turpin et al. 1990; Munoz
et al. 1994; Harz: Behr et al. 1987; Spain: Halliday and
Mitchell 1984; Canals and Cardellach 1993). Although
magmatic heat can be augmented by radioelement heat
production (Willis-Richards and Jackson 1989) this may
be insucient to account for all these examples of later
¯uid activity and other causes of ¯uid in®ltration need to
be sought. These observations suggest that in some cases
the relationship between alteration and mineralisation
and the host granite needs re-evaluation.
The Connemara region has experienced numerous
¯uid-rock interaction events. The oldest (Yardley et al.
1983, 1991) was associated with regional metamorphism
(490±465 Ma, Tanner 1990; Cli€ et al. 1996), the most
recent was Tertiary (G.R.T. Jenkin, P. Mohr,
J.G. Mitchell & A.E. Fallick, in revision). While some
localised alteration and mineralisation within the 400
Ma (Caledonian) Galway Granite batholith is related to
magmatic ¯uids (Derham and Feely 1988; Gallagher
et al. 1992), other alteration is caused by later nonmagmatic ¯uids (Menuge et al. 1997; G.R.T. Jenkin,
P. Mohr, J.G. Mitchell & A.E. Fallick, in revision).
Moreover, the Galway Granite and its host rocks display widespread retrograde alteration. Jenkin et al.
(1992) examined the retrograde alteration of south-west
Connemara and concluded that the granite and its host
rocks were altered by the same ¯uid. However, these
authors could not resolve whether there had been two
phases of alteration involving two ¯uids with similar
characteristics, one just after granite intrusion, the other
more than 90 Ma later, or whether all the alteration took
place in a single event >90 Ma after intrusion of the
Galway Granite. The latter possibility implies that much
alteration in the batholith (and country rocks) may be
unrelated to magmatic heat. This has implications for
genetic models of the mineralisation and alteration
spatially associated with the Galway Granite. In addi-
tion, either possibility allows a further opportunity to
study alteration and mineralisation that are spatially,
but not temporally, related to an upper crustal granite.
Here we characterise the physical conditions, composition, possible sources and timing of ¯uids that have affected the Galway Granite in order to understand better
the relationship of the di€erent ¯uid phases to the
Granite.
Geological setting
The Connemara complex is a Caledonian inlier in western Ireland
(Fig. 1a). Metasedimentary and metavolcanic rocks of late-Precambrian (Dalradian) age (Rogers et al. 1989) were intruded at
490 Ma by the Meta-Gabbro Suite (MGS), comprising ultrama®c
to granitic intrusions (Leake 1989). The Galway Granite batholith
was intruded at 400 Ma (Leggo et al. 1966; Pidgeon 1969) into the
Dalradian and MGS, and to the south into the metasediments and
metavolcanics of the Lower Ordovician South Connemara Group
(Williams et al. 1988). To the east the granite has a faulted contact
with Lower Carboniferous limestones. Gravity and aeromagnetic
data show that the batholith extends for several km to the south
and east under the Lower Carboniferous rocks of Galway Bay
(Murphy 1952; Madden 1987), where the contact with the limestones is both faulted and unconformable (O'Reilly et al., in press).
The granite and its host rocks are cut by the Teach DoÂite suite of
Upper Carboniferous dolerite dikes (Mitchell and Mohr 1987;
G.R.T. Jenkin, P. Mohr, J.G. Mitchell & A.E. Fallick, in revision).
Granite lithologies examined here range from granodiorite
(Carna Granite) through adamellite (Errisbeg Townland Granite;
ETG) to alkali-feldspar leucogranite (Murvey Granites; MG). The
latter, which are the youngest intrusions, form six bodies, in the
centre (Costelloe Murvey Granite = CMG), and at the margins of
the batholith (Fig. 1a), and were probably intruded at shallow
depth, since the CMG has sub-volcanic features. Between the
Shannawona and Barna Faults, the central block of the batholith
(predominantly a ma®c, megacrystic variety of ETG) is shown by
gravity studies to be 3±4 km thinner (Madden 1987). Hornblende
geobarometry (Leake and Ahmed-Said 1993) suggests the central
block granites crystallised at greater depth than adjacent portions,
and were uplifted by late-magmatic di€erential slip against marginal granite, and upward faulting between the Shannawona and
Barna Faults, probably before intrusion of the CMG at shallower
depth.
Fluid activity in the Galway Granite:
veins and retrograde alteration
A large number of veins have been identi®ed throughout the
1500 km2 of the batholith (O'Reilly et al., in press), although
much of the central block is poorly exposed. A vein classi®cation
scheme is established here based on paragenesis, structural setting
and cross-cutting relationships.
V1 veins
These are associated with molybdenite mineralisation and appear
to be localised to small areas within the batholith. So far, these
veins have been identi®ed at Murvey, Mace Head, Kilkieran,
Costelloe and Barna. Molybdenite also occurs in pegmatites, aplites, in greisened Murvey Granite (Gallagher et al. 1992), and joint
coatings. At Mace Head (Derham and Feely 1988; Gallagher et al.
1992), V1 veins cut, and are cut by, aplite and pegmatite dikes. The
V1 veins occur in stockworks and are 0.4±50 cm thick, have sharp,
planar walls, and comprise massive, milky-clear quartz, with
122
Fig. 1a Geology of the study
area b Sample locations; sample
numbers are pre®xed by R in
the text. One granite sample
(R9) and four V3 veins (R3, R8,
R15 and R16) were from the
Costelloe Quarry (Q on b). Two
V1 veins (Ma-13 and MH-19)
from Mace Head are described
by Gallagher et al. (1992)
muscovite ‹ K-feldspar ‹ pyrite ‹ molybdenite ‹ chalcopyrite
‹ wolframite. Deformation is evidenced by undulose extinction in
quartz. The wallrock is generally neither brecciated nor sheared. A
weak alteration halo extending several cm from the vein contains
pyrite ‹ muscovite ‹ K-feldspar ‹ molybdenite. Derham (1993)
classi®ed the Mace Head deposit as a calc-alkaline (plutonic-type)
stockwork Mo-deposit (Sutherland-Brown 1976).
V2 veins
The V2 veins occur throughout the batholith, in®lling steeply dipping faults and fractures ®rst active during post-consolidation uplift (Leake and Ahmed-Said 1993). They are typically 10±25 m
thick, but up to 50 m in the Shannawona fault, and comprise
quartz ‹ hematite ‹ sericite ‹ chlorite ‹ calcite‹epidote. The
two most common orientations are NNE±SSW (parallel to the
Shannawona Fault) and east±west (e.g. sample R70 which in®lls a
fault contact between the granite and South Connemara Group).
Sulphides have not been observed to be associated genetically with
these veins, but minor autunite on joint faces appears to be
(O'Reilly et al., in press). Recurrent shearing and brecciation of
wallrock and veins are evident. The oldest quartz in a V2 vein is
typically massive and shows undulose extinction; the youngest
generations are unstrained, occurring in cross-cutting microveins of
drusy comb-quartz with vug-®lling chalcedony. Wallrock alteration
may extend tens of metres from larger veins and comprises hematization, silici®cation and alteration of plagioclase to sericite and
biotite to chlorite. The V2 veins cut aplites and pegmatites but were
not observed to cut, or be cut by, other vein types or the Upper
Carboniferous dikes.
V3 veins
These are mostly east±west trending, vertical±steeply dipping veins,
<22 cm thick, ®lling late fractures and strike-slip faults throughout the granite. They contain quartz ‹ ¯uorite ‹ calcite ‹ barite
‹ galena ‹ chalcopyrite‹sphalerite‹pyrite with minor chlorite/
clay. Wallrock alteration occurs up to 1 m from the veins and is
characterised by sericitization of plagioclase, strain-free quartz,
muscovite, ¯uorite, pyrite, galena and chalcopyrite. Brecciated
fragments of wallrock occur in several veins and veins exhibit
comb-quartz and vugs of several cm. Drusy quartz crystals often
contain growth zones de®ned by primary ¯uid inclusions. Veins
typically show two generations of mineral deposition, with galena
and chalcopyrite usually associated with, or later than, the second,
vug-in®lling generation. Fluorite and barite joint coatings cut
several V1 veins and overgrow hematite coatings of possible V2 age.
The R16 V3 vein in the CMG quarry cuts a dolerite dike of the
Teach DoÂite suite. O'Connor et al. (1993) obtained a clinopyroxene
40
Ar/39Ar age of 231 ‹ 4 Ma from this dike, which they interpreted as representing the intrusion age. However, the dike is
similar in mineralogy, chemistry and orientation to other Teach
DoÂite dolerites which Mitchell and Mohr (1987) and G.R.T. Jenkin, P. Mohr, J.G. Mitchell & A.E. Fallick (in revision), believe are
Upper Carboniferous (305 Ma). The sample dated by O'Connor
et al. (1993) was collected by C. O'Reilly only 2 m from the R16 V3
vein. Therefore, the age determined by O'Connor et al. (1993) may
represent not the intrusion age, but the age of alteration of the dike
by ¯uid from the R16 vein. Such an alteration age is consistent with
Sm/Nd data for V3 vein ¯uorites from the CMG Quarry (Menuge
et al. 1997).
Abundance: 7%
F: 0.50±0.80 (21)
TMCO2: )56.8 to )56.5 (18)
TMclathrate: 4.9±7.2 (18)
(º 9.2 to 5.4 eq. wt% NaCl)
THCO2 to V: 27.3±31.2 (9)
ÐÐÐ
ÐÐÐ
TH(L): 306±411 (14)
ÐÐÐ
ÐÐÐ
TD: 297±341 (5)
Mol% CO2: 3±11 (9)
qtot: 0.704±0.881 (9)
Absent
Abundance: 100%
F: 0.70±0.98 (270)
TFM: )33.4 to )19.4 (23)
TM: )7.4 to )0.3 (298)
(º 11.0 to 0.5 eq. wt% NaCl)
TH(L): 153.3±317 (305)
qtot: 0.681±0.930 (270)
Abundance 50%
F: 0.88±0.99 (37)
TFM: )65.3 to )41.2 (11)
TMh: )26.0 to )23.4 (7)
TM: )13.7 to )4.7 (44)
(º 17.5 to 7.4 eq. wt% NaCl)
TH(L): 137.8±229 (38)
qtot: 0.904±1.029 (35)
d
Abundance: 36%
F: 0.10±0.75 (75)c
TMCO2: )57.1 to )56.3 (66)
TMclathrate: 4.2±8.2 (56)
(º 10.3 to 3.6 eq. wt% NaCl)
THCO2 to V: 28.4±31.3 (35)
THCO2 by mf: 30.7±31.2 (8)
THCO2 to L: 21.5±31.3 (10)
TH(L): 305±368 (8)
TH(Critical): 332±358 (4)
TH(V): 306±385 (16)
TD: 263±391 (34)
Mol% CO2: 5±59 (37)
qtot: 0.454±0.890 (37)
Abundance: 18%
TMCO2: )57.1 to )56.6 (15)
THCO2 to V: 25.0±31.4 (7)
THCO2 by mf: 30.0±31.1 (3)
THCO2 to L: 28.6±31.2 (3)
qtot: 0.243±0.642 (13)
Abudance: 100%
F: 0.70±0.97 (336)
TFM: )32.6 to )20.3 (39)
TM: )6.7 to )0.2 (294)
(º 10.1 to 0.3 eq. wt% NaCl)
TH(L): 150.8±302 (296)
qtot: 0.686±0.954 (246)
Abundance: 82%
F: 0.87±0.99 (65)
TFM: )61.1 to )49.5 (23)
TMh: )27.9 to )23.0 (35)
TM: )25.2 to )5.7 (63)
(º 25.7 to 8.8 eq. wt% NaCl)
TH(L): 117.3±212 (60)
qtot: 0.947±1.123 (51)
Type 1: 5±40 lm;
LH2 O ‡LCO2 ‡ V CO2 .
In vein quartz isolated
and clusters; in granite
quartz clusters and
planar arrays;
()) crystal, sub ())
crystal, elongate and
irregular shapes
Type 2: 5±15 lm;
LCO2 ‡ VCO2 : occur in
granite quartz in planar
arrays; ()) crystal,
sub ()) crystal and
irregular shapes
Type 3: 1±<30 lm;
LH2 O ‡ VH2 O ; colourless ice on freezing;
planes and trails; ())
crystal, sub ()) crystal,
elongate, ¯attened and
irregular shapes
Type 4: 2±200 lm;
LH 2 O ‡ V H 2 O ; pale-dark
brown ice on freezing;
planes and trails;
()) crystal, sub ())
crystal, tubular and
irregular shapes
Abundance: 33%
F: 0.93±0.98 (12)
TFM: )65.6 to )58.7 (9)
TMh: )25.2 to )23.9 (10)
TM: )16.1 to )15.6 (10)
(º 19.5 to 19.1 eq. wt% NaCl)
TH(L): 133.1±169.1 (10)
qtot: 1.047±1.074 (8)
Abundance: 100%
F: 0.75±0.98 (103)
TFM: )33.8 to )19.6 (7)
TM: )5.5 to )0.5 (132)
(º 8.5 to 0.8 eq. wt% NaCl)
TH(L): 147.2±302 (127)
qtot: 0.712±0.967 (112)
Absent
Absent
Granite east
of BF [3]
b
SF Shannawona Fault, BF Barna Fault
[11] = number of samples
c
(15) = number of measurements
d
Abundance: percentage of class of samples containing a given type of ¯uid inclusion, e.g. 4 of the 11 granite samples (36%) west of the SF contain type 1 inclusions
e
Inclusion contents at room temperature (LH 2 O aqueous liquid phase, LCO2 CO2 -rich liquid phase, VCO2 CO2 -rich vapour phase, VH 2 O water-rich vapour phase)
a
e
e
e
e
Granite between
SF and BFa [14]
Granite west
of SFd [11]b
Fluid
inclusion type
Table 1 Fluid inclusion microthermometry data for granite, V1, V2 and V3 veins. [F degree of ®ll = Vaq /Vtot, where Vaq is the volume of aqueous liquid and Vtot is the total volume
of the inclusion; measured at 40 °C for type 1 inclusions, otherwise at room temperature. qtot bulk inclusion density (g cm)3)]. Type 1 inclusion aqueous ¯uid equivalent salinities
calculated from TMclathrate measured in the presence of liquid and vapour CO2 using Collins (1979). Type 3 and 4 inclusion ¯uid equivalent salinities calculated from TM
using Hall et al. (1988). Mol% CO2 values of type 1 inclusions in the system H2O-CO2-NaCl calculated from salinity, THCO2 and estimated volumetric proportion of CO2 using
FLINCOR (Brown 1989). Bold text indicates that some or all of these inclusions occur as primary inclusions in their host mineral; see text for details
123
Absent
Abundance: 29%
F: 0.45±0.85 (35)
TMCO2: )57.3 to )56.6 (31)
TMclathrate: 5.1±8.2 (35)
(º 8.9 to 3.6 eq. wt% NaCl)
THCO2 to V: 28.5±31.2 (17)
THCO2 by mf: 30.5±31.0 (5)
ÐÐÐ
TH(L): 303±322 (4)
ÐÐÐ
ÐÐÐ
TD: 286±357 (27)
Mol% CO2: 3±11 (19)
qtot: 0.672±0.946 (19)
Absent
Abundance: 100%
F: 0.70±0.98 (174)
TFM: )26.0 to )19.8 (10)
TM: )9.8 to )0.4 (186)
(º 13.7 to 0.7 eq. wt% NaCl)
TH(L): 180.0±318 (214)
qtot: 0.722±0.929 (175)
Abundance 29%
F: 0.90±0.97 (12)
TFM: )53.3 to )52.3 (3)
TMh: )23.8 and )22.6 (2)
TM: )14.1 to )4.7 (13)
(º 17.9 to 7.4 eq. wt% NaCl)
TH(L): 120.8±214 (7)
qtot: 0.947±1.042 (6)
Abundance: 67%
F: 0.40±0.80 (14)
TMCO2: )57.1 to )56.3 (9)
TMclathrate: 6.3±7.9 (13)
(º 7.0 to 4.1 eq. wt% NaCl)
THCO2 to V: 29.1±31.1 (6)
THCO2 by mf: 31.3 (1)
THCO2 to L: 30.1±31.2 (6)
TH(L): 247±346 (6)
TH(Critical): 306 (1)
ÐÐÐ
TD: 259±326 (9)
Mol% CO2: 4±19 (12)
qtot: 0.615±0.921 (12)
Not found in this study,
but reported by Gallagher
et al. (1992) in one V1
vein from Mace Head
Abundance 100%
F: 0.80±0.97 (94)
TFM: )33.5 to )21.0 (7)
TM: )6.2 to )0.4 (92)
(º 9.5 to 0.7 eq. wt% NaCl)
TH(L): 156.8±258 (100)
qtot: 0.831±0.935 (72)
Abundance: 66%
F: 0.91±0.99 (4)
TFM: )56.5 to )50.1 (5)
TMh: )23.5 to )23.3 (3)
TM: )19.1 to )17.9 (3)
(º 21.7 to 20.9 eq. wt% NaCl)
TH(L): 63.4±134.1 (3)
qtot: 1.086±1.133 (3)
CO2 homogenisation temperatures
THCO2 to V: to vapour
THCO2 by mf: by meniscus fading
THCO2 to L: to liquid
Melting temperatures
TMCO2: CO2 ice melting
TMclathrate: clathrate melting
TFM: ®rst ice melting
TM: last ice melting
TMh: hydrate melting
Microthermometric abbreviations [all temperatures are given in degrees Celsius (°C)]
Abundance 100%
F: 0.85±0.99 (421)
TFM: )77.6 to )44.4 (158)
TMh: )32.9 to )12.8 (222)
TM: )27.3 to )4.9 (520)
(º 27.0 to 7.7 eq. wt% NaCl)
TH(L): 81.6±228 (476)
qtot: 0.908±1.145 (443)
Absent
Absent
Absent
V3 veins [9]
TH(L): to liquid
TH(Critical): critical homogenisation
TH(V): to vapour
Bulk homogenisation temperatures
Abundance: 33%
F: 0.90±0.98 (27)
TFM: )51.2 and )50.0 (2)
TMh: )28.5 to )22.9 (4)
TM: )12.1 to)7.3 (27)
(º 16.1 to 10.9 eq. wt% NaCl)
TH(L): 131.9±213 (27)
qtot: 0.945±1.035 (25)
Abundance: 100%
F: 0.80±0.96 (95)
TFM: )23.3 to )21.6 (7)
TM: )2.3 to 0.0 (90)
(º 3.8 to 0.0 eq. wt% NaCl)
TH(L): 170.3±268 (97)
qtot: 0.775±0.922 (85)
Absent
V2 veins [3]
V1 veins between
SF and BF [7]
V1 veins west
of SF [3]
Table 1 (Contd.)
TD: decrepitation temperature
PD: pressure at decreptitation
PH: pressure at homogenisation
TT: trapping pressure
PT: trapping pressure
Others
Quartz and calcite hosted FIs
F: 0.89±0.99 (45)
TFM: )62.9 to )48.1 (16)
TMh: )27.0 to )20.5 (33)
TM: )28.0 to )6.5 (45)
(º 27.4 to 9.8 eq. wt% NaCl)
TH(L): 53.4±219 (44)
qtot: 0.934±1.177 (42)
Absent
Absent
Absent
Glengowla
vein [1]
124
125
The Glengowla vein (R93)
Kennan et al. (1988) studied late, low-temperature, replacement
and vein base metal deposits in the Connemara Dalradian which
are often fault hosted, unrelated to igneous activity and distinct
from earlier stratabound and skarn hosted mineralisation. One
such deposit, Glengowla, an east±west trending vuggy lode with
calcite-barite-quartz-galena-pyrite-chalcopyrite-sphalerite in the
Middle Dalradian Lakes Marble Formation, was sampled for
comparison with the very similar V3 veins in the Galway Granite.
Concordant K-Ar dates at 212 ‹ 4 Ma for Glengowla vein material may re¯ect either the time of ore formation or ``remobilisation of ¯uids'' (Halliday and Mitchell 1983).
¯uorite samples. After pre-¯uorination, the samples were rapidly
decrepitated within the reaction vessels of a silicate oxygen preparation line at 250±550 °C. Released water was reacted with BrF5
at 60±300 °C and the resulting oxygen converted to CO2 for isotopic analysis. Insucient material remained to duplicate these
analyses, so their precision and accuracy are unde®ned. Repeat
analyses during the course of this work gave an average d18O of
9.66& (n = 4) for NBS#28 and dD of )0.5& (n = 3) for VSMOW. Analytical precision is estimated to be better than 0.2&
for quartz and barite d18O and sulphide and barite d34S, 0.1& for
calcite d18O and d13C, and 2& for ¯uid inclusion CO2 d13C and
H2O dD. All errors are one standard deviation (r), unless stated
otherwise.
Retrograde alteration
Fluid inclusion petrography and chronology
Jenkin et al. (1992) note that the retrograde alteration seen in the
MGS and Dalradian rocks, (alteration of plagioclase to sericite,
epidote and calcite, alteration of biotite and hornblende to chlorite ‹ epidote ‹ calcite and reddening in K-feldspars), also affected the northern margin of the Galway Granite. Our observations
of the Galway Granite show that, in addition to the localised veinspeci®c wallrock alterations described above, it is almost pervasively a€ected by this retrograde alteration. This alteration is particularly intense adjacent to major faults and V2 veins. At a smaller
scale it is associated with transgranular microveins <250 lm wide
(quartz ‹ sericite ‹ calcite ‹ chlorite ‹ epidote) and cross-cutting planes of ¯uid inclusions. The similarity of the V2 mineral
assemblage to the retrograde alteration of the granite suggests a
similar origin.
Four types of ¯uid inclusion were identi®ed in granite
quartz and veins from the Galway Granite (Table 1).
Their essential features are:
Analytical methods
Vein samples for analysis (Fig. 1b) were selected from the bestdeveloped examples of each vein type and represent a subset of the
total population of veins in the granite. Petrographic and microthermometric studies were carried out on vein minerals (quartz,
¯uorite, calcite and barite) and granite quartz. The petrography of
104 samples was examined using thin sections, doubly polished
wafers and crystals. Inclusion abundances were determined by
counting in 25 ®elds of view (each 62500 lm2) in 30 lm sections.
Microthermometry (2300 inclusions from 49 samples) was carried
out on 200 lm doubly polished wafers using a calibrated Linkam
THM600 heating-freezing stage (MacDonald and Spooner 1981).
Microthermometric data are reduced to salinities and mol% CO2
values, as appropriate, as described in Table 1.
High-purity samples for chemical and isotopic analysis were
obtained by a combination of mineral separation, hand picking
and acid washing (Jenkin et al. 1994). Crush-leach ICP-AES
(Shepherd et al. 1985) was used to determine bulk ¯uid inclusion
chemistry. Duplicate analyses of three samples demonstrate that
within-sample variation is large ( ‹ 20%) compared with the
accuracy of the chemical analyses (<5%) estimated from analysis
of standard solutions (n = 4). Our results are therefore only semiquantitative.
Inclusion ¯uids were extracted by decrepitation under vacuum
at 1100±1200 °C (quartz), 650 °C (calcite), 550 °C (¯uorite) or 400±
450 °C (barite). Cryogenic separation and volumetric measurement
of gas fractions and measurement of d13CCO2 and dDH2O follow
Jenkin et al. (1994). A mercury-piston inlet system to the mass
spectrometer enabled routine dD measurement of the 30±200 lmol
of hydrogen produced by reduction of water with hot uranium. The
volume and d13C of CO2 released from calcite were not measured
since thermal decomposition of calcite also releases CO2. Barite
d18O was measured using a method similar to that of Sakai (1977).
Standard techniques were used for the analysis of d34S in sulphides
and barite (Robinson and Kusakabe 1975; Coleman and Moore
1978). We attempted to measure d18OH2O of the ¯uids in three
Type 1. Rare H2O-CO2 inclusions of moderate salinity
found in some V1 veins and granite quartz, including
some granite samples from areas with no known V1
mineralisation.
Type 2. Rare CO2 inclusions, with no visible water,
found only in granite quartz west of the Shannawona
Fault. Gallagher et al. (1992) found similar inclusions in
a V1 vein from Mace Head.
Type 3. Aqueous inclusions of low±moderate salinity
found in every granite, V1 and V2 sample and especially
abundant around V2-vein in®lled fault zones.
Type 4. Aqueous inclusions of moderate±high salinity
often found in granite quartz, V1 and V2 veins, and
ubiquitous in V3 and Glengowla samples. Their generally higher degree of ®ll, greater size and often irregular
shapes help distinguish them from type 3 inclusions in
granite and V1 and V2 veins.
The relative ages of these ¯uid inclusions are deduced
from textural observations:
1. Type 1 and 2 inclusions are absent in V2, V3 and
Glengowla (R93) veins. Type 3 inclusions are absent in
V3 or Glengowla veins. Type 4 inclusions occur in all V3
veins and in many granites and V1 and V2 veins (Table 1). In granite quartz, type 4 inclusions are most
abundant around ¯uorite- or barite-coated joints and V3
veins.
2. Primary (using the criteria of Roedder 1984) type 1
inclusions occur only in V1 veins. Primary type 3 inclusions occur only in V2 veins. Primary type 4 inclusions occur only in V3 veins and the Glengowla (R93)
vein. Primary inclusions are often accompanied by secondary and pseudosecondary inclusions of the same
type. All inclusion types in granite quartz are secondary,
occurring in planes which cross-cut the granite quartz
crystals.
3. Type 2 inclusions in granite quartz west of the
Shannawona Fault are always spatially associated with
type 1 inclusions, often in the same plane.
4. Planes of Type 1 and 2 inclusions in granite quartz
and V1 quartz are cut by planes of secondary type 3
126
Type 1 inclusions
Fig. 2 Abundance of aqueous (type 3 plus type 4) inclusions versus
degree of alteration in the Galway Granite; 90±100% of the
aqueous inclusions in all samples are type 3
Melting of CO2 ice ()56.8 ‹ 0.2 °C) in type 1 inclusions
occurs close to the eutectic temperature of CO2
()56.6 °C) indicating very pure CO2 (Table 1). The
TMclathrate indicate aqueous phase salinities of 10.3±
3.6 wt% NaCl eq. (Table 1), with granite-hosted type 1
inclusions tending to be more saline. Type 1 inclusions in
V1 veins and in the single sample of granite east of the
Shannawona fault that contained them (R32), have a
restricted range of CO2 contents (3±19 mol%: Table 1)
and all but one show total homogenisation to the
aqueous liquid, or decrepitated while doing so. Such
features suggest these type 1 inclusions were trapped
from a homogeneous ¯uid. In contrast, type 1 inclusions
in granite quartz west of the Shannawona fault (Table 1)
have a wide range in CO2 contents (5±59 mol%) and
show total homogenization to the aqueous liquid, to the
vapour phase, critically, or else decrepitated.
Type 2 inclusions
inclusions (often in three distinct planes) but never vice
versa. Planes of secondary type 4 inclusions cut planes of
all other inclusion types in granite quartz, V1 and V2
veins.
5. The V3 veins are seen to cut V1 veins (see above).
Thus, an aqueous-carbonic ¯uid (type 1 and 2 inclusions) which deposited V1 veins was the earliest ¯uid
in the Galway Granite. Subsequently an in¯ux of low±
moderate salinity type 3 aqueous ¯uid deposited V2
veins. The higher salinity type 4 ¯uid, which deposited
V3 veins, is the youngest ¯uid to have in®ltrated the
batholith and its envelope.
Every granite sample examined from the batholith
has been a€ected by retrograde alteration and evidence
suggests that this was caused by type 3 ¯uid: these inclusions are present in every granite sample examined
and are often the only inclusion type present, constituting 85±100% by number. Furthermore, type 3 inclusion abundance in granite quartz is roughly proportional
to the degree of retrograde alteration (Fig. 2). In contrast, in granites with a high degree of retrograde alteration and high abundance of type 3 inclusions there are
few or no type 1 and type 2 inclusions. Wallrock alteration to the V1 veins is restricted to within a few cm of
the vein and is texturally earlier than the retrograde alteration. Type 4 ¯uids cannot have caused widespread
retrograde alteration as they are absent in many of the
granites examined, being restricted to regions around V3
veins or barite- or ¯uorite-coated joints.
Fluid inclusion microthermometry
Results are summarised in Table 1, together with all
microthermometric abbreviations used in the text.
The TMCO2 ()56.8 ‹ 0.2 °C) again indicates very pure
CO2 (Table 1). No clathrate or ice melting was observed in any type 2 inclusion, but small quantities of
water could be present since it preferentially wets inclusion walls, and at low concentrations may form an
invisible ®lm around the walls of the inclusion (Kreulen
1987).
The coexisting type 1 and 2 inclusions in granite
quartz west of the Shannawona Fault show features
indicating a common origin from a parent ¯uid which
experienced immiscibility; both types occur in the same
plane and the type 1 inclusions show a large range in
CO2 content and homogenize both to liquid and to vapour, as well as critically. Type 2 inclusions therefore
represent the CO2-rich phase of the immiscible ¯uid,
type 1 inclusions the water-rich phase. Figure 3 shows
that most type 1 inclusions lie on the relevant solvus at
about 1 kbar, suggesting trapping at that pressure and
300±400 °C. However, ¯uids which homogenize to the
vapour should plot to the CO2-rich side of the solvus
crest; i.e. they should have more CO2 than they are
calculated to have. This may be due to systematic errors
in estimating the volumetric proportion of CO2 in CO2rich inclusions and may also re¯ect the near-critical behaviour often seen in inclusions around the crest of the
¯at-topped solvus in this system at pressures >1 kbar
(Shepherd et al. 1985).
Alternatively the wide range in CO2 contents of coexisting type 1 and 2 inclusions could result from variable water leakage from initially homogeneous trails of
H2O-CO2 inclusions (Bakker and Jansen 1990). However we consider that post-trapping modi®cation of inclusion compositions is insigni®cant: type 1 and 2
inclusions are present in roughly the same proportions in
three granite samples showing lightly strained quartz as
in a granite sample showing highly strained quartz.
127
Fig. 3 T versus mol% CO2 plot for type 1 and 2 inclusions in
granite quartz west of the Shannawona Fault. The H2O-CO2-6
wt% NaCl eq. solvus of Bowers and Helgeson (1983) is shown. An
error of ‹10 mol% CO2, due mainly to errors in the determination
of the volumetric proportion of CO2, is assumed for all mol% CO2
determinations for type 1 inclusions. Type 2 trapping temperatures
are assumed to equal to those for type 1 inclusions as indicated by
TH and maximum TD (Table 1). As the quantity of water in type 2
inclusions cannot be determined, a range of CO2 content is shown
Type 3 inclusions
Mean TFM ()24.2 ‹ 4.2 °C; Fig. 4) is close to the eutectic temperature of the H2O-NaCl-KCl system,
()23 °C; Shepherd et al. 1985). The TM indicates salinities of 0.0±13.7 (3.8 ‹ 2.5) wt% NaCl eq. (Table 1,
Figs. 4, 5). Very low salinity type 3 inclusions are found
in all sample types, but the more saline inclusions
(>4 wt% NaCl eq.) are restricted in occurrence to
granite and V1 samples (Fig. 5). The V2-hosted type 3
Fig. 4 TM versus TFM plot showing the di€erent chemistry and
salinities of type 3 and 4 ¯uid inclusions. (GG Galway Granitehosted samples, Gw Glengowla veins). Top axis shows equivalent
salinities in the NaCl-H2O system (Hall et al. 1988) corresponding
to TM. Terms eutectic and halite are respectively the salinities of
eutectic melting and halite saturation in the NaCl-H2O system
Fig. 5 TH versus salinity plot for granite (GG) and vein-hosted
¯uid inclusions in the Galway Granite and at Glengowla (Gw)
inclusions are therefore less saline on average
(1.5 ‹ 0.9 wt% NaCl eq.) than other type 3 inclusions
(4.0 ‹ 2.5 wt% NaCl eq.) and also have TH at the low
end, and TFM at the high end, of the type 3 range although showing similar degrees of ®ll (Table 1,
Figs. 4, 5). Furthermore, primary V2-hosted type 3 inclusions tend to be less saline than the secondary ones
(ranges are 0.0±1.8 and 0.3±3.8 wt% NaCl eq. respectively) indicating a progression to higher salinity with
vein development. Type 3 inclusions must have less than
1 mol% CO2 (Hedenquist and Henley 1985) as neither
liquid CO2, nor clathrates, were seen during microthermometry.
Type 4 inclusions
The TFM shows a wider range (Fig. 4) than in the type 3
inclusions. The highest TFM values correspond to the
eutectic temperature of the H2O-CaCl2-NaCl system
()55, )52 °C; Shepherd et al. 1985). The brown-coloured ice (Table 1) is also diagnostic of the presence of
CaCl2 (Shepherd et al. 1985). The lower TFM values
could result from the presence of additional components, such as LiCl, which depress the H2O-CaCl2-NaCl
eutectic (Borisenko 1982) or misidenti®cation of subsolidus recrystallisation as ®rst melting. The latter can be
ruled out, as appreciable quantities of liquid were present in the low-TFM inclusions by )60 °C. Hydrate
melting ()32.9 to )12.8 °C: Table 1) in almost all inclusions examined occurred before ice melting. The TM
indicates a wide range in salinities extending to much
higher values (8±28 wt% NaCl eq.) compared to type 3
inclusions (Figs. 4, 5). The TM values below the NaClH2O eutectic, and even below halite saturation in this
system when no halite is observed in these inclusions,
emphasise the multicomponent nature of this ¯uid
(Fig. 4). For inclusions with paired hydrate (TMh) and
ice-melting (TM) data, compositions estimated in the
H2O-CaCl2-NaCl system (Oakes et al. 1990) mostly
cluster around a point with 13 wt% salinity and XNaCl
128
[ = NaCl/(NaCl + CaCl2) weight ratio] » 0.7 (range:
0.48±0.92). A second group of more saline inclusions
cluster around 22 wt% salinity, XNaCl » 0.65 (range
0.48±0.8). The mol% CO2 for type 4 inclusions, like that
for type 3 inclusions, is inferred to be £1 mol%, for the
same reasons.
On a TH-salinity plot (Fig. 5) the type 4 inclusions
display what is usually interpreted (Shepherd et al. 1985)
as a mixing trend between ¯uids of di€erent salinities but
similar TH. That ¯uids of di€erent salinities have ¯owed
through V3 veins is better demonstrated by considering a
single vein (Fig. 6). As with most V3 veins, this vein
exhibits a two-stage mineral paragenesis. Early combquartz is coated by vug-®lling quartz crystals, galena,
chalcopyrite and ¯uorite. Primary inclusions in the early
comb-quartz showed relatively low salinities, whereas
planes of secondary inclusions in the same quartz are
appreciably more saline. Primary inclusions in a growth
zone in the vug quartz were even more saline, whereas
planes of secondary inclusions in the vug quartz were
more dilute again, almost as dilute as the primary inclusions in the comb-quartz (Fig. 6). All inclusions, in
both early and late quartz show similar TFM, TMh and
TH. Thus, in just one V3 vein there is, at minimum, a
temporal progression from low- to high-salinity ¯uids
and back to moderate-salinity ¯uids. No correlation was
found between type 4 inclusion composition and location within the batholith however.
Fig. 6 TM versus TH plot showing that both high- and low-salinity
type 4 ¯uids have ¯owed through the R8 V3 vein. Top axis as in
Fig. 4
Type 4 inclusions in quartz and calcite from Glengowla show great similarity to the type 4 ¯uids in the V3
veins (Table 1, Figs. 4, 5) and are hereafter regarded as
being formed in the same event. Quartz-hosted inclusions show a bimodal salinity distribution between
25±27 and 12±17 wt% NaCl eq.; only the moderate-
Table 2 Bulk ¯uid inclusion chemistries by crush-leacha ICP-AES
Molalitiesd
Sample
number
Sample
typeb
Host
mineralc
Inclusion
type
Atomic ratios
K/Na
Ca/Na
NaCl
KCl
CaCl2
R82 {i}
R82 {ii}
Average
GG
GG
Qz
Qz
3 > 1,2,4
3 > 1,2,4
0.10
0.15
0.13
0.03
0.05
0.04
0.64
0.60
0.62
0.06
0.09
0.08
0.02
0.03
0.02
R24
R43 {i}
R43 {ii}
Average
V1
V1
V1
Qz
Qz
Qz
3 > 1,2,4
3 > 1,2,4
3 > 1,2,4
0.31
0.21
0.16
0.19
0.02
0.06
0.13
0.10
0.68
1.20
1.15
1.18
0.21
0.25
0.18
0.22
0.01
0.07
0.15
0.11
R75 {i}
R75 {ii}
Average
V1
V1
Qz
Qz
3 > 1,2,4
3 > 1,2,4
0.16
0.12
0.14
0.03
0.02
0.03
1.06
1.11
1.17
0.17
0.13
0.15
0.03
0.02
0.03
R76
Ma-13
MH-19
V1
V1
V1
Qz
Qz
Qz
3>4
3 > 1,2,4
3>4
0.19
0.35
0.16
0.26
0.09
0.23
0.42
0.75
0.81
0.08
0.27
0.13
0.11
0.06
0.19
R46
V2
Qz
3
0.22
0.06
0.26
0.06
0.01
R8
R28
R33
R67
V3
V3
V3
V3
Qz2
Qz1
Qz1
B
4
4
4
4
0.16
0.24
0.30
0.07
0.16
0.24
0.09
0.32
3.75
2.05
1.91
2.20
0.60
0.50
0.58
0.15
0.59
0.48
0.18
0.69
R93a
Gw
Qz1
4
0.37
0.15
2.36
0.87
0.35
a
Crushing carried out on 0.5±2 mm fractions, except for granite quartz (0.8±1 mm)
(GG Galway Granite, V1 V1 vein, V2 V2 vein, V3 V3 vein, Gw Glengowla)
c
(Qz quartz, Qz1 V3 early quartz, Qz2 V3 late quartz, B Barite)
d
Molalities calculated using the method of Weisbrod and Poty (1975), using mean sample type 3 inclusion salinities for granite, V1 and V2,
and mean sample type 4 inclusion salinities for V3
b
129
salinity population was found in calcite. From paired
TMh-TM data, ¯uid compositions (Oakes et al. 1990)
range from a cluster around 11 wt% salinity,
XNaCl » 0.85 (range: 0.62±0.95) in quartz and calcite, to
a group with 25 wt% salinity and XNaCl » 0.58±0.8 in
the quartz.
Chemistry of ¯uid inclusion leachates
Samples chosen for bulk chemical analysis of ¯uid inclusion leachates were dominated by a single inclusion
type (Table 2). The granite quartz and V1 veins contained >85% by number of type 3 inclusions. Since
type 4 inclusions tend to be larger and more saline than
type 3 inclusions, the small proportion of type 4 inclusions in R76 and MH-19 could signi®cantly a€ect the
composition of the bulk leachate.
Analyses of type 3 ¯uid-dominated samples (excluding R76 and MH-19) con®rm type 3 ¯uid is dominated
by Na and K (Table 2, Fig. 7), in line with the TFM data
(Fig. 4). The K/Na (0.22 ‹ 0.09) and Ca/Na (0.06
‹ 0.03) atomic ratios for these samples are similar to
those found in geothermal waters (Ellis 1979). Analyses
of V3 and Glengowla samples con®rm type 4 ¯uid is
enriched in Ca compared to type 3 (Table 2, Fig. 7), but
the K/Na (0.23 ‹ 0.12) of type 4 ¯uid spans a similar
range. The two V1 vein samples which bear signi®cant
levels of type 4 ¯uids, (R76 and MH-19), have bulk
leachate compositions which appear to be dominated
chemically by type 4 ¯uid (Fig. 7, Table 2). Type 4 ¯uid
de®nes a ®eld between oil®eld brines and type 3 ¯uid
(Fig. 7).
(1982) criteria for ¯uid immiscibility. They are spatially
associated, often trapped in the same planes. Type 1
inclusions that do not decrepitate, homogenize to different states over the same temperature range. Both inclusion types exhibit a wide range of compositions
around the 6 wt% NaCl solvus crest (Fig. 3). Therefore,
TT (305±390 °C; Fig. 3) and PT (¯uid pressure at TT) for
both inclusion types in granite quartz west of the
Shannawona Fault are given by type 1 inclusion TH and
PH (¯uid pressure at TH) respectively. The TD and PD
(¯uid pressure at TD) provide minima. The determination of PH using an equation of state is compromised by
available equations being intended for generally higher
temperatures (>350 °C; Bowers and Helgeson 1983) or
higher densities (>1.0 g cm)3; Brown and Lamb 1989).
Since many type 1 inclusions in granite quartz west of
the Shannawona Fault have salinities close to 6 wt% we
suggest it is more reliable to read pressures directly from
a solvus plot based solely on experimental data (Fig. 3);
this gives pressures of 0.8±1.8 kbar (Fig. 8), similar to
Fluid PT conditions
Type 1 and 2 ¯uids
West of the Shannawona Fault, type 1 and 2 inclusions
in granite quartz appear to satisfy the Ramboz et al.
Fig. 7 Na-Ca-K composition (wt%) of ¯uid inclusion leachates
(Table 2). Oil®eld brines from Carpenter et al. (1974)
Fig. 8 Pressure-temperature diagram showing PT constraints for
di€erent inclusion types. Field for type 1 and 2 inclusions in quartz
west of the Shannawona Fault determined by TH/TD range for
type 1 inclusions and pressures taken from the solvus diagram
(Fig. 3). All type 1 isochores calculated using the Bowers and
Helgeson (1983) equation of state. Bold-rimmed box (V1) shows
possible PT ®eld for type 1 ¯uids in V1 veins suggested by d18O
data. Field for type 3 inclusions in Galway Granite and V1 veins
formed by intersection of isochores and chlorite temperatures
(mean ‹ 1rn)1 for all chlorite analyses) for the Galway Granite
calculated with the Cathelineau (1988) geothermometer. Filled
circles show intersection of mean isochores and mean chlorite
temperatures for the ®ve Granite samples on which chlorite
geothermometry was carried out; an individual uncertainty box
(‹rn)1) is only shown for R40, which lies outwith the range of
most chlorite temperatures. Trapping temperatures of type 3
inclusions in V2 veins are unconstrained, so only a range of
isochores are shown (between isochore for mean TH+1rn with
mean salinity+1rn and an isochore for the mean TH)1rn with
mean salinity)1rn). Type 4 inclusion envelope corresponds to the
region covered by isochores for mean data for each V3 phase
investigated. Also shown are mean isochores for R16 quartz and
calcite and the pressure drop required if they were deposited at the
same temperature. All isochores constructed using FLINCOR
(Brown 1989)
130
that obtained (0.7±1.5 kbar) using the recent equation of
state of Duan et al. (1995; Z. Duan, personal communication). These ranges are consistent with a pressure of
<1.5 ‹ 1.0 kbar estimated for ®nal crystallisation of
granite hornblende west of the Shannawona Fault
(Leake and Ahmed-Said 1993).
In contrast to the inclusions in the granite quartz,
type 1 inclusions in the R75 V1 vein from Kilkieran west
of the Shannawona fault show no evidence of phase
separation. This suggests that ¯uid P and/or T conditions were high enough to prevent immiscibility, perhaps
because this vein formed at an early stage of uplift. An
isochore for mean primary type 1 inclusions indicates
this vein could have formed at similar, or higher, P-T to
the inclusions in the host granite quartz (Fig. 8), maximum pressure being 1.5 ‹ 1 kbar (Leake and AhmedSaid 1993). Gallagher et al. (1992) determined an oxygen
isotope temperature of 360±450 °C for intergrown V1
muscovite and quartz at Mace Head, and from a type 1
inclusion isochore estimated pressures of 1.3±2.1 kbar
(Fig. 8). However, their PT estimate may be unreliable
because they state ``some phase separation may have
taken place at the time of trapping'' (their p. 321), in
which case TH and PH would provide better estimates of
the PT at trapping. These values, the origin of their
isochore, are compatible with the PT range for the inclusions in the host-granite quartz (Fig. 8).
To the east of the Shannawona fault, none of the
type 1 inclusions in the two V1 veins and the single
sample of granite which contain them show evidence
that they result from phase separation. Minimum pres-
sures are given by the solvus (1.5 kbar), maximum
pressures are given by Mn-rich garnets in CMG pegmatites and aplites cut by V1 veins and molybdenite
mineralised joints, which suggest pressures <3 kbar
(Whitworth and Feely 1994). Using this range, TT for
type 1 inclusions east of the Shannawona fault is 366±
567 °C (Fig. 8).
The higher pressures and temperatures for type 1
inclusions in granite quartz east of the Shannawona
fault compared with those in granite quartz to the west,
may be a consequence of greater uplift of the central
block. Final juxtaposition of the central block with the
western part took place by movement along faults such
as the Shannawona Fault (Leake and Ahmed-Said
1993), i.e. syn-V2. Higher depths and pressures may
account for the scarcity of CO2-bearing ¯uids east of the
Shannawona fault, since they will tend to concentrate at
higher levels within a batholith.
Type 3 ¯uid
Chloritization of biotite relates to in®ltration of type 3
¯uid, so chlorite chemistry geothermometers were used
to estimate type 3 ¯uid temperatures for ®ve granite
samples for which microthermometric data were also
available (Table 3). Chlorite replacing biotite may inherit some chemical features if alteration is incomplete
(Cathelineau 1988), therefore only completely chloritised
biotites were analysed. The chemical homogeneity of
chlorite from throughout the batholith suggests that it
Table 3 Average chlorite chemistries by electron microprobe and thermometry results. (MG Murvey Granites, ETG Errisbeg Townland
Granite, METG Matic Errisbeg Townland Granite)
Sample number
R 7a
R 34a
R 40a
R 41a
R 42a
GJ 003b
GJ 213b
GJ 220b
Lithology
MG
ETG
METG
Leuco-granite
METG
ETG
MG
SiO2
TiO2
Al2O3
FeO(T)
MnO
MgO
CaO
K2O
Na2O
total
nc
T Walshe(°C)d,e
TCath(°C)d
Type 3 TH(°C)
27.89
0.04
19.93
26.60
0.90
15.69
0.08
0.05
0.59
91.77
10
251 ‹ 15
311 ‹ 11
247 ‹ 26
27.10
0.13
20.89
25.51
0.70
16.67
0.05
0.09
0.54
91.68
9
291 ‹ 22
340 ‹ 13
241 ‹ 35
28.79
0.28
17.85
24.88
0.58
16.94
0.24
0.09
0.55
90.20
5
204 ‹ 10
269 ‹ 8
210 ‹ 27
26.70
0.07
19.39
26.88
1.00
14.66
0.11
0.05
0.42
89.28
10
263 ‹ 26
320 ‹ 20
230 ‹ 38
27.27
0.11
20.20
25.34
0.66
16.86
0.14
0.08
0.67
91.33
7
280 ‹ 25
330 ‹ 17
227 ‹ 27
Roundstone
Granite
26.86
0.00
18.54
22.80
0.60
17.68
0.07
0.00
±
86.55
11
264 ‹ 10
312 ‹ 8
±
26.30
0.12
18.72
22.49
0.68
16.17
0.01
0.05
±
84.54
16
253 ‹ 28
308 ‹ 21
±
24.29
0.06
18.86
33.82
1.26
7.90
0.10
0.05
±
86.34
13
302f
336 ‹ 13
±
a
Samples E of Shannawona fault analysed by C.O'R. on probe sections (2 lm beam, 15 kV)
Samples W of Shannawona fault analysed by G.R.T.J. on mounts of chlorite separates (2 lm beam, 20 kV). TCath values for these
analyses represent a revision of those in Jenkin et al. (1992) which were based an earlier version of this geothermometer
c
Oxide values are averages of n analyses. Standard deviations are generally > 1 wt%, the biggest variation being in Mg and Fe,
representing Mg/Fe variability
d
Temperatures were calculated for individual analyses and the mean and rn±1 calculated from the results, the mean will therefore di€er
slightly from that calculated from the mean oxide values
e
Calculated for PH2 O ˆ 1 kbar; this geothermometer is insensitive to pressure
f
Only one analysis converged to a solution
b
131
formed at similar temperatures, in accord with the small
range of all type 3 inclusion TH (232 ‹ 29 °C; Table 3,
Fig. 5).
Temperatures given by the Cathelineau (1988) geothermometer (TCath) are higher than those determined
with the Walshe (1986) geothermometer (TWalshe). The
TWalshe appear to be low estimates of type 3 ¯uid temperatures, given there is signi®cant overlap between
TWalshe and TH for R7 and R40 (Table 3), implying
pressures around the liquid-vapour curve, yet no boiling
is observed in these samples. The TCath are consistent
with the lack of boiling since they are all higher than TH
for coexisting type 3 inclusions, although De Caritat
et al. (1993) note that this geothermometer can overestimate temperature by up to 200 °C when applied to
mineral assemblages di€erent to those in which it was
calibrated. The Galway Granite is however, compositionally similar to the acid volcanics at the Los Azufres
system where the Cathelineau (1988) geothermometer
was calibrated, and furthermore, mean TCath for all
granite chlorites (318 ‹ 23 °C rn)1) is within error of
the robust quartz-epidote oxygen isotope temperature
(297 ‹ 13 °C rn)1) estimated by Jenkin et al. (1992) for
retrograde alteration in the country rocks. Therefore,
TCath are taken as reasonable estimates of the type 3
¯uid temperatures and suggest trapping pressures of
0.5±2.3 kbar (Fig. 8) except in R40 where rather lower
temperatures (270 °C) and pressures (0.4±1.6 kbar)
applied.
The TT for the V2 type 3 inclusions is unconstrained.
Stable isotope evidence (see below) indicates that V2
veins were major conduits which carried surface-derived
¯uid into the granite. Thus, PT may be lower than for
the same ¯uid in the granite, perhaps similar to that for
R40, which, in contrast to most granite samples, contains type 3 inclusions of comparable salinity to those in
V2 veins.
Type 4 ¯uid
The vuggy textures of the Glengowla and V3 veins
combined with the brittle nature of the host rock are
taken to indicate a hydrostatic pressure regime for type 4
¯uid. As type 4 inclusions show no evidence of boiling,
¯uid pressures are assumed to be ³PH. Maximum ¯uid
pressures are constrained by the maximum depth of
burial; earlier type 1 inclusions in granite quartz west of
the Shannawona fault, near the roof of the batholith
were trapped at pressures £2 kbar, corresponding to a
depth of <7.5 km for lithostatic pressures. If the pressure regime in the V3 veins entirely hydrostatic then for
a depth of 7.5 km, maximum P¯uid is 0.7 kbar. For this
pressure range type 4 inclusions were trapped at 112±
258 °C (Fig. 8, Table 4).
In three of the V3 veins (R16, R67, R93) the secondformed mineral has a signi®cantly higher estimated
formation temperature than the ®rst-formed (Table 4).
For R67 this can be attributed to the down-temperature
re-equilibration of the barite-hosted inclusions, since the
majority of inclusions in the barite have necked to
monophase liquid. However, in R16 and R93, the ®rstformed mineral is quartz, in which inclusions are signi®cantly less susceptible to re-equilibration. Stretching
of inclusions in the second-formed calcite during microthermometry could have arti®cially increased their TH,
but no evidence of this was observed. Di€erences in
trapping pressure could allow ¯uids with di€erent TH to
have been trapped at similar temperatures; however a
pressure drop between quartz and calcite deposition of
1.2±1.5 kbar would be required (e.g. arrow A on Fig. 8),
implying that the ¯uid that deposited the quartz would
have to have been partly lithostatically pressured. While
some pressure ¯uctuation cannot be discounted, it seems
more likely that there was signi®cant increase in temperature during formation of some V3 veins. In R16, the
last-formed ¯uorite was deposited at similar temperature
to the ®rst-formed quartz, suggesting cyclical temperature variations.
Stable isotope results
V1 veins
The V1 quartz ¯uid inclusion dD clusters around )20&
(rn)1 = 4&) (Table 4), comparable to previous values
at Mace Head (Gallagher et al. 1992), which range from
)18 to )20&, with an outlier at )29&. As type 3 inclusions account for 85±100% by number of inclusions
in V1 quartz, bulk dD, rather than re¯ecting the dD of
the ¯uid that deposited the vein, re¯ects the type 3 ¯uid
dD (see below). The lower dD of R76 and MH24(C)
()29&; Gallagher et al. 1992) and also perhaps F4 JD
(no microthermometric determinations) relate to the
additional presence of lower dD type 4 ¯uids (see below).
Low mol% CO2 values of 0.5±3% for V1 bulk volatiles (Table 4), re¯ect the dominance of CO2-poor
type 3 ‹ 4 ¯uids in the V1 veins. The new d13 CCO2
values ()8.0 ‹ 2.0&; Table 4) for samples where immiscibility was not observed, may re¯ect the d13 CCO2 of
the ¯uid that deposited the veins. The wider range obtained by Gallagher et al. (1992) from the MH-19 vein
(d13 CCO2 ˆ ÿ14:2 to )4.3&) may relate to the immiscibility that occurred in that sample. The d13 CCO2 of )8&
are consistent with a magmatic carbon source in type 1
and 2 ¯uids, although this value could also conceivably
be generated by a mixture of marine and organic carbon.
Six V1 quartz samples from throughout the batholith
(Fig. 1, Table 4) yielded a d18O range of 9.1 to 9.9&
(with one outlier at 10.7&), overlapping closely that
determined by Gallagher et al. (1992) for six V1 quartzes
from Mace Head (9.0 to 10.3& excluding three marginal
samples a€ected by later low-d18O ¯uids). This narrow
V1 quartz d18O range suggests that, although the veins
are dominated by type 3 ‹ 4 ¯uid inclusions, most of
them have not exchanged oxygen isotopes with these
later, low-temperature, low- d18O ¯uids, which would
132
result in more scattered quartz d18O (Table 4). The homogeneity in quartz d18O also implies mixing with other
¯uids was not important in V1 deposition. In order to
obtain such homogeneous V1 quartz d18O (9.5 ‹ 0.4&,
n = 12) over the large (250 °C) temperature range
indicated by microthermometry for these veins (Tables 1,4) would require that the ¯uid d18O varied by up
to 6& and, coincidentally, varied with temperature with
the same slope as the quartz-water fractionation curve.
This might seem improbable, but could come about if
the ¯uid depositing the quartz was in equilibrium over
this temperature range with a large volume of granite
homogeneous in d18O, i.e. the ¯uid was a magmatic/
magmatic-metamorphic ¯uid (Sheppard 1986). Evidence
suggests the batholith when intruded was relatively homogeneous in d18O; granite from the west of the batholith little a€ected by later alteration has d18O = 8.9
‹ 0.4&, n = 6, and quartzes from these samples are
Table 4 Stable isotope data (ND not determined, Gw Glengowla).
Measured (Meas.) ¯uid d values are a bulk value for the total ¯uid
inclusion population. Calculated (Calc.) ¯uid d values are derived
from the bulk mineral values by means of fractionation factors at
the appropriate temperature. Oxygen isotope fractionation factors:
quartz-water from Matsuhisa et al. (1979); calcite-water from
O'Neil et al. (1969) as modi®ed by Friedman and O'Neil (1977);
barite-water from Kusukabe and Robinson (1977). Salt e€ects on
mineral-¯uid oxygen isotope fractionations calculated from mean
equivalent salinities of primary FIs assuming that they are ap-
proximated by the salt e€ects for pure NaCl solutions (Horita et al.
1995). Fluid CO2 d13C calculated assuming all carbon is present in
the ¯uid as H2CO3 (app) = H2CO3 + CO2 (aq.) and that
d13 H2 CO3 …app† d13 CCO2 (gas) (cf. Jenkin et al. 1994). Carbon
isotope fractionation factor between calcite and CO2 (gas) from
Chacko et al. (1991). V3 and Glengowla d34 SCDT values (&) : R8
pyrite 3.5, R16 chalcopyrite 0.3, R67 barite 20.4, R93a pyrite 12.7,
R93b barite 21.9, R93c barite 20.9, R93c galena )5.1 R93c chalcopyrite )3.7, R93c sphalerite )0.7
Mineral
Mean
salinity of d18 OSMOW
primary
(&)
¯uid
inclusions
(wt%
NaCl eq.)
Calcite
d13CPDB
(&)
Mol%
CO2 of
¯uid
inclusion
volatilesb
Calc./Meas.
¯uid H2O
d18OSMOW
(&)
Meas.
Bulk
¯uid
H2O
dDSMOW
(&)
Calc./Meas.
¯uid CO2
d13CPDB
(&)
366±567c
366±567c
412±567f
318±466h
318±466h
318±466h
318±466h
?6.0d
?6.0d
6.6g
5.5g
?6.0d
?6.0d
?6.0d
10.7
9.9
9.7
9.1
9.2
9.6
9.7
±
±
±
±
±
±
±
1.32
0.48
1.16
2.97
0.82
ND
1.38
6.1±9.7e
5.3±8.9e
6.3±8.7e
3.0±6.7e
3.1±6.8e
3.5±7.2e
3.6±7.3e
)19
)19
)14
)18
)24
)24
)18i
)9.2
ND
)9.0
)5.7
ND
ND
)9.4i
3
3=4
295±341j
295±341j
1.9
0.6
7.5
15.3
±
)7.1
0.34
ND
0.5±2.0
9.6±10.8
)17
)24
ND
)4.3 to )3.8
4
4
4
4
4
4
4
4
4
143±241k
180±255k
131±252k
105±176k
178±244k
85±172k
139±232k
118±156k
215±262k
23.5
13.2
16.7
20.6
23.0
19.8
15.9
13.1
13.9
12.9
12.2
±
14.3
11.4
±
12.3
10.7
±
±
±
±
±
)4.0
±
±
±
±
0.74
ND
ND
ND
ND
3.00
0.56
ND
ND
)2.9±4.0
)0.6±3.8
()7.3 to 5.3)1
)5.7 to 1.3
0.9±4.3
)3.5 to )1.5
)4.0 to 2.8
)1.6 to 1.7
)0.2 to 1.8
)38
)29
)26
ND
)17
)41
)45
)31
)18
)18.7
ND
ND
ND
)4.0 to )2.1
)9.5
ND
ND
ND
R93a Gw Quartz (1st) 4
R93a Gw Calcite (2nd) 4
118±211k
195±256k
15.6
11.3
15.9
±
3.08
13.6; (15.1)m )2.0; ()0.8)mND
)2.5 to 5.2
3.9 to 6.8
)32
)24
)16.8
)1.4 to 0.2
R93b Gw Barite (?)
118±256k,n 13.5
11.5
)0.8 to 8.3
)27
ND
Sample
Host
minerala
Inclusion Mineral
type
formation
temperature
(°C)
R23 V1
R24 V1
R43 V1
R75 V1
R76 V1
F4 JD V1
Ma-13 V1
Quartz
Quartz
Quartz
Quartz
Quartz
Quartz
Quartz
3
3
3»1
3»1
3>4 » 1
?
3
R46 V2
R68 V2
Quartz
Calcite
R8 V3
R15 V3
R15 V3
R16 V3
R16 V3
R16 V3
R28 V3
R67 V3
R67 V3
Quartz
Quartz (1st)
Fluorite (2nd)
Quartz (1st)
Calcite (2nd)
Fluorite (3rd)
Quartz
Barite (1st)
Fluorite (2nd)
a
(E)
(E)
(E)
(W)
(W)
(W)
(W)
4
±
ND
E and W in parentheses refer to position relative to Shannawona Fault for V1 samples, rankings for V3 samples show relative position in
vein paragenesis
Volatiles are dominantly H2O and CO2
c
As no primary ¯uid inclusions were found in these samples, the estimated trapping temperature range of type 1 inclusions in the Central
Block (see text) was used
d
No salinity measurements possible
e
The bulk ¯uid would have a higher d18O, by virtue of also containing CO2 (cf. Jenkin et al. 1994)
f
Estimated using mean R43 type 1 inclusion TH and PT of 1.5±3.0 kbar
g
From clathrate melting temperature (Collins 1979) of primary type 1 inclusions
h
Temperatures for R75 between TH and T at P = 2.5 kbar
i
From Gallagher et al. (1992)
j
Chlorite formation temperature
k
Estimated to lie between [mean] TH)1rn)1 (TH) and [mean TH + 1rn)1 (TH)] pressure corrected to 0.7 kbar using isochore for
mean salinity +1rn)1 (S) using data for primary inclusions only. Isochores for H2O-NaCl-CaCl2 system (Zhang and Frantz 1987)
l
Believed to be erroneous since lies to 18O-poor side of MWL
m
Kennan et al. (1988)
n
All inclusions have necked to monophase liquid, so the combined range for R93 quartz and calcite is used
b
133
9.8 ‹ 0.2&, n = 3, identical to the values in the V1
veins (Jenkin 1988; Jenkin et al. 1992).
V2 veins
Fluid inclusions in R46 are exclusively type 3 so the bulk
water dD of )17& (Table 4) provides a reliable estimate
for type 3 ¯uid. The R68 calcite contains 50% type 4
inclusions, which probably cause its lower dD. Taking
the dD values for those V1 and V2 samples from this
study and Gallagher et al. (1992) which are almost exclusively populated by type 3 inclusions, gives a type 3
¯uid dD range of )14 to )20& ()18 ‹ 2, n = 7). The
type 3 ¯uid dD is probably the same as its source, since
¯uid dD is very resistant to change during water-rock
interaction except when ¯uid/rock ratios are very low
(Sheppard 1986). High ¯uid/rock is indicated by the
high abundance of type 3 inclusions in granite quartz
and the narrow range in type 3 ¯uid dD. The CO2-poor
Fig. 9 d18O-dD plot of type 3 ¯uid in V2 veins (black symbols and
horizontal open arrow) and type 4 ¯uid in V3 and Glengowla veins
(open symbols). Range in d18O for V2 samples corresponds to
maximum and minimum temperature (Table 4). Error bars for single
V3 ¯uorite ¯uid d18O values correspond to estimated error on direct
measurement. Symbols for other V3 minerals are at d18O corresponding to mean primary TH (lower d18O) and mean T at 0.7 kbar, and
error bars on these points are ‹1rn)1 in TH. EM1 and EM2 are
possible end-member compositions of type 4 ¯uids. The ®elds of other
natural waters (Sheppard 1986) and the seawater evaporation curves
of Holser (1979) (H) and Pierre et al. (1984) (P) are shown for
comparison. The curve of Holser (1979) has been extended to 45X
following Knauth and Beeunas (1986). Multipliers against evaporation curves (e.g. 4X) are evaporation ratios (weight of H2O in original
seawater/weight of H2O in residual evaporated brine). Gypsum facies
corresponds to 3X)11X, halite facies to 11X)65X
nature of type 3 ¯uids is highlighted by the CO2 content
of only 0.34 mol% of the volatiles in R46 quartz. The
d13C of the R68 calcite is compatible with a derivation of
type 3 ¯uid carbon from the granite.
The high dD of type 3 ¯uid and, within the Shannawona fault, its low d18O (0.5±2.0&; R46 quartz) and
lower than seawater salinity (Table 4) suggest type 3
¯uid was of meteoric origin (Fig. 9). If the dD of type 3
¯uid ()18&) is taken to be that of the original meteoric
¯uid, then this had d18O » )3.5&. The higher d18O of
type 3 ¯uid in R46 is undoubtedly the result of oxygen
isotope exchange with the Galway Granite. The ¯uid
d18O for the R68 calcite (9.6±10.8&) is considerably
higher than that for R46 quartz, cannot be generated
solely by exchange with the granite at type 3 ¯uid temperatures and could be interpreted as a metamorphic
¯uid (Fig. 9). In fact it probably re¯ects an initially
greater degree of equilibration of type 3 ¯uid with the
granitic wallrock in this unusually thin (10 cm) vein,
followed by low-temperature oxygen isotope exchange
between the calcite and later type 4 ¯uid which also
occurs in it (c.f. Jenkin et al. 1994).
At the time of granite intrusion (400 Ma), Connemara was beneath a subaerially exposed depositional
basin in a tropical climate (27°S; Torsvik 1985; Anderton et al. 1979). The Earth was not glaciated (Frakes
1979) so seawater dD and d18O would have been lower
than at present, perhaps similar to estimates for nonglacial seawater (Sheppard 1986). We estimate therefore,
that at 400 Ma meteoric ¯uid in Connemara would have
had dD » )22 to )30& (Fig. 9; cf. Fallick et al. 1985),
comparable to our estimate for the type 3 ¯uid source.
However, meteoric ¯uids with similar dD are also
available later during the evolution of Connemara (see
below).
134
Type 4 ¯uid shows a wide range of bulk dD ()45 to
)17&), d18O ()5.7 to 8.3&), mean salinity (11.3±23.5;
wt% NaCl eq. of primary inclusions) and deposition
temperature (85±215 °C or 156±262 °C, depending on
pressure) (Table 4). All of these parameters vary, not
only between di€erent vein samples, but even between
minerals within the same sample (Fig. 6). Furthermore,
no consistent trend in any of these parameters is observed with the paragenetic sequence when di€erent vein
samples are compared (Table 4). Type 4 ¯uid therefore
varied spatially and temporally.
Type 4 ¯uid appears to be highly variable, but the
variations are all well correlated (Figs. 9, 10), even
though dD and salinity are for bulk ¯uids, whereas d18O
and temperature are for primary ¯uids. The dD, d18O
and temperature are all positively correlated, whereas
salinity is negatively correlated with these three parameters. The best correlation is seen for dD versus d18O
estimated from TH; a line can be drawn (dD/d18O » 7;
Fig. 9) that passes through the error bars of all but two
samples, for which special explanations apply. The
correlation between dD and d18O estimated from temperatures at 0.7 kbar is less good, because the ¯uorite
¯uid d18O, being directly measured, do not move to
higher values with increasing temperature as the other
samples do. We take this to imply that temperatures
were close to TH, and pressure corrections are
> 0.7 kbar. Of the few data points lying away from the
d18O-dD trend, the ¯uid d18O for R15 ¯uorite, on the
low-d18O side of the meteoric water line, is probably a
spurious result from the direct d18O measurement
method. The ¯uid d18O for R93 calcite appears too high,
but can be attributed to oxygen isotope exchange be-
tween the hot ¯uid which deposited this calcite and the
high-d18O marble wallrock at Glengowla (note that such
exchange need not have taken place with the lower
temperature ¯uid that deposited the quartz).
The correlations of dD, d18O and temperature in
type 4 ¯uid are qualitatively similar to those in formation waters from sedimentary basins (Sheppard 1986
and references therein). However, in formation waters
salinities increase with d18O, whereas in type 4 ¯uid the
opposite is the case. In addition, formation water temperatures up to 200 °C are extreme, and formation water
slopes on a d18O-dD plot are lower (dD/d18O » 4.8±1.2).
A simple formation water origin for type 4 ¯uid is
therefore excluded.
The covariation of dD, d18O, temperature and salinity
of type 4 ¯uid is best explained as the result of mixing of
two ¯uids, as suggested by the microthermometry data.
Assuming temperatures near to TH, one possible endmember (EM1) was a high-dD ()17&), high-d18O
(1.2&), high-T (205 °C) ¯uid of moderate salinity
(bulk = 12 wt%), the other (EM2) had low dD ()45&),
low d18O ()3&), low T (125 °C), but high salinity
(bulk = 21 wt%) (Figs 9, 10). Of course, the true endmembers may be more extreme in composition and
temperature, since all bulk values measured may themselves be mixtures. More extreme salinities and TH are
measured on individual type 4 inclusions, and, projecting
to extreme TH (80 and 230 °C), would suggest EM2 has
dD = )61&, d18O = )5.4& and EM1 has dD = )8&,
d18O = 2.5& (Fig. 9). Salinities are 25 and 10 wt%
respectively, very close to the measured salinity range of
type 4 ¯uid (Fig. 5). Because most of the variation in type
4 ¯uid d18O can be explained by mixing between two
¯uids, this implies that, apart from the ¯uid that deposited R93 calcite, the d18O of type 4 ¯uid was a€ected
Fig. 10a,b (V3 and Glengowla mineral (type 4 ¯uid) bulk inclusion
¯uid dD versus: a Bulk ¯uid inclusion salinity (mean of all inclusion
measurements, both primary and secondary). Error bars for
salinity are standard error. b Mean primary TH. Error bars are
‹1rn)1 TH. EM1 and EM2 are possible end-member compositions
of type 4 ¯uids
Glengowla and V3 veins
135
little by ¯uid-rock interaction during V3 deposition. This
is despite V3 veins being relatively narrow and causing
sericitic alteration.
The few d13 CCO2 measured for type 4 ¯uid also show
a rough correlation with dD suggesting that the endmembers may have di€erent sources of CO2; ¯uids with
dD >)25& have d13 CCO2 of )4.0 to 0.2&, whereas
those with dD <)32& have d13 CCO2 of )9.5 to )18.7&
(Table 4). While the high d13C of the ¯uid that deposited
R93a calcite probably re¯ects carbon derivation from
the host marbles, such an explanation is not available
for the ¯uid that deposited R16 calcite in the centre of
the granite. The low d13 CCO2 for EM2 suggests an origin
involving organic material, the high d13 CCO2 for EM1
suggests an origin from marine carbonate.
The d34S for R67 and Glengowla barites (Table 4)
falls within the small range (20.4±23.4&, n = 4) determined by Kennan et al. (1988) for late veins in Connemara outside the Galway Granite (including one from
Glengowla). This homogeneity in barite d34S, together
with the dominance of barite over sulphides in these
veins suggests that ¯uid sulphate d34S was little a€ected
by reduction to sulphide. Sulphate d34S may therefore be
inherited from the source. Stratabound mineralisation in
Connemara has d34S <10& (Kennan et al. 1988), suggesting that the Dalradian sediments were not the sulphur source. An alternative is that the sulphate is from
seawater. Since intrusion of the Galway Granite in the
Lower Devonian, the only time when seawater d34S was
as high as 20 to 24& was the Upper Devonian±Lower
Carboniferous 380±330 Ma (Upper Devonian; Claypool et al. 1980, Lower Carboniferous estimates from
Lower Carboniferous hosted base-metal deposits in
Ireland; Caul®eld et al. 1986 and references therein;
Anderson 1990).
The V3 sulphides (Table 4) could be derived from
magmatic sulphide in the granite, which has similar d34S
(Laouar et al. 1990). Although there is no textural evidence that barite and sulphides grew in equilibrium, the
approach to isotopic equilibrium between base-metal
sulphide and sulphate (262±304 °C ‹ 20 °C; fractionation factors of Ohmoto and Rye 1979) suggests an origin
for the sulphides by equilibrium with sulphate in an
oxidising ¯uid. Pyrite at Glengowla (d34S = 12.7&; this
study, and )23.7&; Kennan et al. 1988) shows disequilibrium with other phases and must relate to some
other process, probably reduction of sulphate.
The low d18O ()3 to )5.4&) estimated for EM2 indicates that the water is ultimately of meteoric origin.
However, its dD of )45 to )60& is distinctive, being
lower than any dD we have estimated for meteoric waters in Connemara later than intrusion of the Galway
Granite, except in the Tertiary. Speci®cally, at
210 Ma, during the Upper Triassic±Lower Lias (Harland et al. 1990), the best estimate of the age of the V3
veins, Connemara was 30°N (Smith et al. 1981) and
undergoing the transition from low-relief subaerial
maritime conditions to semi-marine lagoons in a hot dry
climate (Anderton et al. 1979; Wilson 1981; Ziegler
1990). During the Upper Triassic the continents were
not glaciated (Frakes 1979). We estimate that at this
time, surface ¯uids in Connemara may have ¯uctuated
between meteoric ¯uid with high dD ()21 to +3&) and
evaporated seawater with dD ranging from )7& up to
+13& and down to )27& or even lower (Fig. 9).
Clearly EM2 was not derived from Upper Triassic meteoric water, but EM1 has a dD compatible with such a
source. An origin for EM1 solely from evaporated seawater can be excluded since evaporated seawater with
d18O and dD near EM1 will be close to halite saturation
(26.3 wt%; Holser 1979), whereas EM1 contains only
9±12 wt% salinity. The EM1 could represent Upper
Triassic meteoric water or unevaporated seawater that
underwent a positive 18O shift by ¯uid-rock interaction
before mixing with EM2. However, in this case an additional source of salinity is required. This is unlikely to
be ¯uid-rock interaction in situ because earlier, hotter,
type 3 ¯uids, attained at most only 8 wt% salinity by
¯uid-rock interaction (see below). The EM1 may have
gained salinity by solution of evaporites or perhaps EM1
itself represents a mixture of Upper Triassic meteoric
water with evaporated seawater.
Discussion
Origin of type 1 and 2 ¯uids
The remarkable homogeneity of V1 quartz d18O across
the batholith is interpreted as due to equilibration of V1
¯uids with the granite at low ¯uid/rock ratios. However,
it is questionable whether the V1 ¯uid would have
equilibrated oxygen isotopes with the granite at temperatures as low as 320 °C (Table 4). Furthermore,
small ¯uctuations in the temperature of quartz precipitation from a ¯uid equilibrated with the granite at
320 °C would result in a larger scatter in quartz d18O,
e.g. ¯uctuations of only 20 °C would produce a 1.3&
spread in quartz d18O. The stable isotope data therefore
demand quartz precipitation at higher temperature. V1
temperatures of 600 °C are also indicated by the latemagmatic nature of V1 veining and the 568±655 °C
oxygen isotope temperatures for quartz-magnetite pods
related to molybdenite mineralisation (Gallagher et al.
1992). In addition, the recent fractionation factor of
Zheng (1993), gives revised temperatures of 520±640 °C
for the V1 quartz-muscovite oxygen isotope data of
Gallagher et al. (1992). Such temperatures are higher
than those inferred for other calc-alkaline (plutonictype) stockwork Mo deposits, although they are reached
in other types of Mo-mineralisation (Westra and Keith
1981). Most type 1 inclusions may record lower temperatures because, despite being texturally primary, they
have re-equilibrated volumes to lower temperatures
(increased density) during cooling (Sterner and Bodnar
1989; Barker 1995). Maximum pressure of V1 vein
deposition is <3 kbar (garnets in CMG aplites)
and probably less than 1.5 ‹ 1 kbar (hornblende
136
geobarometry); however pressures must have been high
enough (>1 kbar at T >600 °C) to preclude immiscibility. The V1 temperatures and pressures may therefore
have been higher and lower respectively compared to
those indicated by ¯uid inclusion data (Fig. 8) raising
the possibility that such resetting may have occurred in
other similar deposits.
Carbonic ¯uids in granite quartz west of the Shannawona fault were trapped at lower temperatures
(<400 °C) than the V1 veins, indicating continued, but
less focused, release of carbonic ¯uids with declining
temperature.
On the basis of the temperature, the C and O isotope
data, and the S isotope data of Gallagher et al. (1992), an
origin wholly within the granite (``magmatic'') is implied
for V1 ¯uid. In contrast, the bulk dD values for V1 veins
are not representative of the primary ¯uids (type
1 ‹ type 2 inclusions), as type 3 inclusions (and to a
lesser extent type 4 inclusions) are always volumetrically
far more important. This ``overprint'' of the dD signature
by later ¯uids, which leave the quartz d18O intact, provides a salient reminder that it cannot be assumed that
bulk inclusion dD of vein material relates to the vein
forming event (Pickthorn et al. 1987) without careful
examination of the inclusions within a regional context.
Origin of type 3 ¯uid
Stable isotope and salinity data indicate that type 3 ¯uid
was of meteoric origin. The presence of hematite, epidote and autunite, and absence of sulphides in V2 veins
suggests type 3 ¯uid was oxidising, consistent with a
meteoric origin.
Microthermometry indicates type 3 ¯uid was generally more dilute within V2 veins than in the granite and
V1 veins (Fig. 5), and became more saline during V2
development. Type 3 ¯uid in grain boundaries, fractures
and joints within the granite undoubtedly experienced
lower ¯uid/rock ratios than when channelled in veins.
The inference therefore is that the meteoric ¯uid gained
salinity by ¯uid-rock interaction within the granite (e.g.
Kamineni 1987; Ferry 1985). The V2 veins probably
represent the main conduits by which the meteoric ¯uid
penetrated the granite.
Type 3 ¯uid PT within granite quartz (0.5±2.3 kbar;
Fig. 8) exceeds the 0.7 kbar limit for hydrostatically
pressured ¯uids in the Galway Granite, indicating that
type 3 ¯uid is partially lithostatically pressured (overpressured) in the granite. That type 3 ¯uid is surface
derived therefore presents a paradox, since this requires
surface ¯uids to ¯ow up a pressure gradient. The most
common mechanism by which meteoric ¯uids can penetrate the crust is in a buoyancy-driven convection system (Criss and Taylor 1986). While meteoric-convection
systems require hydrostatic pressure to operate, surface
¯uids within them can become lithostatically pressured
(Fournier 1991; Jenkin et al. 1994). Hydrostatic ¯uids
passing through ®ne cracks in the granite (secondary
inclusion trails) could have become overpressured
around the time of trapping, as a result of a combination
of tectonic movements and mineral precipitation.
We have linked the retrograde alteration within the
Galway Granite to type 3 ¯uid. Previously Jenkin et al.
(1992) correlated the alteration within the granite with
that in the host rocks and showed that the alteration in
both was caused by a high-dD, low-d18O ¯uid indistinguishable from type 3 ¯uid (c.f. Kennan et al. 1988).
Thus type 3 ¯uid is inferred to have also ¯owed through
the host-rocks to the granite. Yardley et al. (1983) report
secondary aqueous ¯uid inclusions with similar TM and
TH to type 3 and 4 inclusions in metamorphic veins up to
18 km north of the granite. The probable extent of in®ltration of type 3 ¯uid is thus several thousand km2;
comparable with the fossil geothermal systems seen
around subvolcanic batholiths (Criss and Taylor 1986).
Kennan et al. (1988) suggest that meteoric waters caused
the late base-metal mineralisation in Connemara and
mineralisation is often associated with meteoric-convection systems (Criss and Taylor 1986). However, apart
from minor hematite and autunite, mineralisation is
never spatially associated with V2 veins, except adjacent
to mineralised V1 or V3 veins. Although the vast quantities of type 3 ¯uid which circulated throughout the
Galway Granite and Connemara may have produced
near-surface mineralisation, it has produced none of
signi®cance at the present erosion level.
Timing and cause of type 3 ¯uid in®ltration
and retrograde alteration
The timing and cause of type 3 ¯uid in®ltration remain
enigmatic. The most straightforward explanation is that
type 3 ¯uid in®ltration resulted from convection through
the heat anomaly of the Galway Granite soon after intrusion at 400 Ma (Jenkin et al. 1992). However, the
Upper Carboniferous Teach DoÂite (TD) dikes, both in
and around the granite, show widespread isotopic and
chemical alteration (Mitchell and Mohr 1987; G.R.T.
Jenkin, P. Mohr, J.G. Mitchell & A.E. Fallick in revision) which the latter conclude relates to an Upper
Triassic hydrothermal event. While this is the estimated
age of type 4 ¯uid, two factors suggest that type 4 ¯uid
could not have caused most of this alteration.
1. Type 4 ¯uid is restricted to regions close to V3 veins
and joints, whereas alteration in the TD dikes is not.
2. G.R.T. Jenkin, P. Mohr, J.G. Mitchell & A.E.
Fallick (in revision) estimate that the TD dikes were altered by ¯uids with dD » )7 to )15&. This is consistent
with the dikes being altered by type 4 EM1, but rules out
involvement of the low-dD EM2. Given the intimacy of
association of these two end-members it seems unrealistic
that the dikes were only altered by one of them.
If type 4 ¯uid did not cause the alteration of the TD
dikes, then what did? Provided that we have not missed
a ¯uid event in our regional ¯uid inclusion survey of 104
137
samples from throughout the batholith, we are led to the
conclusion that the dikes must have been altered by
type 3 ¯uid. The dD of the ¯uid which altered the dikes
is indistinguishable from that of type 3 ¯uid, given the
uncertainties in the estimation of ¯uid dD. If the alteration in the TD dikes was caused by type 3 ¯uid, this
constrains type 3 ¯uid to be £305 Ma (the age of the
dikes), and, according to the interpretation of G.R.T.
Jenkin, P. Mohr, J.G. Mitchell & A.E. Fallick (in revision) of the K-Ar systematics of the dikes, to be Upper
Triassic (210 Ma). This implies that type 3 ¯uid in®ltration occurred just before type 4, perhaps as an early
phase of the same event. This option lacks appeal in that
we have diculties in explaining such high ¯uid temperatures (>300 °C) in the Upper Triassic. In contrast,
a 400 Ma age for the type 3 ¯uid has the advantage of
both a mechanism for ¯uid in®ltration (thermally induced convection), and a reason for high ¯uid temperatures (magmatic heat), but is inconsistent with the dike
data. A solution to this dilemma could be that G.R.T.
Jenkin, P. Mohr, J.G. Mitchell & A.E. Fallick (in revision) are in error, all the alteration of the TD dikes took
place at the time of their intrusion in localised hydrothermal systems, and type 3 ¯uid is indeed 400 Ma old.
At present therefore, the timing of type 3 ¯uid in®ltration and retrograde alteration remains unresolved.
The most geologically reasonable option remains that
the type 3 ¯uid in®ltrated the granite soon after intrusion in a meteoric convection system. The alternative,
that type 3 ¯uid in®ltration did not occur until
210 Ma, if found to be true, has implications for
models of retrograde alteration spatially associated with
granites.
Origin of type 4 ¯uid and base-metal mineralisation
All the available dating evidence (Halliday and Mitchell
1983; O'Connor et al. 1993; Menuge et al. 1997) points
to a mid±late Triassic age (210 Ma) for the V3 veins.
The interpretation of the K-Ar systematics of the TD
dikes by G.R.T. Jenkin, P. Mohr, J.G. Mitchell & A.E.
Fallick (in revision) also suggests a major ¯uid event in
the area at 210 Ma.
Type 4 ¯uid EM1 has dD and d18O consistent with a
surface-derived origin during the Upper Triassic, most
probably involving both seawater or meteoric water, and
evaporated seawater or evaporites. Permo-Triassic evaporites occur both in o€shore basins (Tate and Dobson
1989; Tate 1993) and to the east in Co. Cavan (Gardiner
and McArdle 1992). The high d13 CCO2 of EM1 suggests
derivation from marine carbonate and occasional dolomite and limestone bands are developed in the PermoTriassic and Rhaetian basins o€shore (Tate and Dobson
1989). Alternatively the carbonate could have been derived from Visean Carboniferous limestone, which may
have overlain the batholith, and remains to south and
east. Menuge et al. (1997) show that Sr in type 4 ¯uid
was not derived from the granite but could have been
derived from seawater, Carboniferous limestones or
younger rocks, although it is not clear whether this refers
to one or both end-members. Further evidence supporting an Upper Triassic surface-derived origin for
EM1 comes from a fault-®lling vein north-west of the
granite which was originally interpreted as containing
the ¯uid which caused retrograde alteration in the area
(GJ.196; Jenkin et al. 1992). It is now clear from the vein
mineralogy (quartz-calcite-barite) that it is a V3 vein,
unconnected with the alteration. The ¯uid inclusions in
the calcite have high dD ()22&; Jenkin et al. 1992),
indicating it is dominated by EM1. The barite d34S is
+12.6& (Jenkin 1988), distinct from that found in the
barites in this study (which probably relate to EM2±see
below), but consistent with derivation of sulphate in
EM1 from Lower Permian or Triassic seawater or evaporites (Claypool et al. 1980). Thus O, H, C and S stable
isotope data are all consistent with an Upper Triassic
surface-derived source for EM1.
The origin of EM2 is more obscure, but taken together the data indicate a basinal brine (high-salinity
formation water) origin. The high salinity and Ca content of some type 4 inclusions are distinctive features of
basinal brines (e.g. Carpenter et al. 1974). Since EM2
has salinities 2±3 times that of EM1, chemical and isotopic compositions of the saline component of type 4
¯uid will generally be dominated by EM2. Thus only
EM2 need be Ca-rich. The major chemical di€erence
between bulk type 4 ¯uid and basinal brines is the higher
K content of the former (Fig. 7; K/Na atomic ratio is
0.07±0.37 and £0.1, Carpenter et al. 1974, respectively).
This may be due to the excess K being derived from
EM1, which may have a chemistry closer to type 3 ¯uid,
or could be explained by reaction between the brine and
wallrocks. The latter must have happened before in®ltration of type 4 ¯uid at the present level, since there it
caused wallrock sericitization (K addition).
Barites analysed in this study, and by Kennan et al.
(1988), have homogeneous d34S (21.4 ‹ 1.2& n = 7),
distinct from that suggested for EM1, and since, as
noted above, the saline component in most type 4 ¯uids
will be dominated by EM2 we suggest that this corresponds to the value for EM2. This provides a strong
constraint on the age of the basin from which EM2
sulphate was derived, since if the sulphate was derived
from seawater, it can only be Upper Devonian±Lower
Carboniferous seawater. Gypsum ‹ halite are known
to have been deposited in sabkhas around the margins of
the Lower Carboniferous sea, both to the east (Anderton et al. 1979; Grennan 1992), and west (Robeson
et al. 1988). The presence of a sub-Visean unconformity
in north Connemara (Dewey and McKerrow 1963) attests to submergence during this time (c.f. Ziegler 1990).
Thus the d34S data suggest that EM2 was generated in,
or at least passed though, a Lower Carboniferous basin.
Evaporation of seawater to an evaporation ratio of
45´ could produce halite-saturated brine with d18O and
dD similar to EM2. Such an origin for the water in EM2
seems unlikely however, since the amount of evaporite in
138
the Lower Carboniferous is generally small, and only
small amounts of such a brine could be formed at such a
high evaporation ratio. Formation waters often form an
array originating from the Meteoric Water Line on a
d18O-dD plot (Fig. 9), which has been interpreted to
indicate they are dominantly of meteoric origin (e.g.
Clayton et al. 1966). However, meteoric waters with dD
as low as EM2 are not thought to have been present in
Connemara in the Lower Carboniferous or subsequently
up to the Upper Triassic. The continents were glaciated
during much of this period (Frakes 1979), and Connemara was between 7°S and 27°N (Smith et al. 1981);
minimum meteoric water dD was )22& (Fig. 9). A
number of studies have identi®ed instances in which
basinal ¯uid dD is lower than that of any available
meteoric water (Ghazban et al. 1991; Fallick et al. 1993;
Munoz et al. 1994; Wilkinson et al. 1995) and this is
generally explained by derivation of water hydrogen
from organic matter (Sheppard 1986). The low d13CCO2
in EM2 also suggests the in¯uence of organic carbon.
Lower Carboniferous shales could have been rich in
organic material, but a more likely source is Namurian
black shales, such as were deposited in the Clare delta to
the south (Anderton et al. 1979) or Westphalian coal
measures which probably covered much of the Lower
Carboniferous (Sevastopulo 1981; Anderton et al. 1979;
Ziegler 1990).
Meteoric ¯uids throughout the Upper Devonian and
Carboniferous probably had a d18O which ranged down
to the upper value of EM2 (Fig. 9). However, the
Namurian rocks of Clare (only 70 km south of the
Galway Granite) were strongly a€ected by the Variscan
orogeny (Upper Carboniferous) centred to the south,
and were heated to 350 °C (Fitzgerald et al. 1994). At
such temperatures any Carboniferous meteoric water in
these rocks would have been 18O-shifted positively by
¯uid-rock exchange to compositions more similar to
metamorphic waters (Fig. 9). Therefore either EM2 was
not derived from Carboniferous rocks to the south of
the granite, but from a more northerly area, or EM2 was
southerly derived, and the basin was recharged with
post-Variscan meteoric water. Sheppard and Langley
(1984) show that a sedimentary sequence can generate
brines more than once. Whether the meteoric water that
developed into EM2 was Carboniferous or younger,
the low d18O of EM2 suggests that minimal alteration of
its d18O took place by ¯uid-rock interaction, consistent
with its low temperature (100 °C).
The salinity in EM2 was probably derived either from
evaporated seawater, or later dissolution of evaporites
within the Lower Carboniferous rocks, or both. Concentration of brines in the Carboniferous could have
additionally been achieved by reverse osmosis through
shale bands (Graf 1982) driven by overpressures generated by the rapid deposition of the Clare delta. Brines up
to 32 wt% NaCl eq. were present in the Clare Namurian
during the Variscan (Fitzgerald et al. 1994). Brines
produced by seawater evaporation or dissolution of
halite would not have the high CaCl2/NaCl of EM2, so
some water-rock interaction is implied (Carpenter 1978),
although it is then perplexing that the oxygen isotope
shift is negligible. Some constituents of the type 4 ¯uid
could have been derived locally; ¯uorine from the ¯uorine-rich Murvey granites (Leake 1974) and lead isotope
data support derivation of V3 Pb from granite feldspar
(Reynolds 1987).
The conclusion that V3 mineralisation results from
mixing of contemporaneous surface-derived ¯uid (EM1)
with ¯uid from a Carboniferous basin (EM2) places
geometric constraints on the paths by which these ¯uids
reached the mixing zone. The EM1 cannot have simply
percolated from an Upper Triassic land surface through
an underlying Carboniferous sequence and then into the
Galway Granite, because EM1 could not then have become hotter than EM2. Rather, EM1 must have travelled a separate path allowing it to heat up more, either
by deep penetration beneath the area, or by traversing
Permo-Triassic basins o€shore, in which concurrent
volcanism took place (Tate and Dobson 1989). The
EM2 may have undergone little additional heating since
leaving its source, and was probably ejected by tectonism from adjacent or overlying Carboniferous sediments. End-Member-1 was therefore probably rising
into the mixing zone. Fluorite mineralisation similar to
V3, and probably relating to the same mineralisation
episode, occurs within Visean limestones to the south of
the Galway Granite (O'Connor et al. 1993). Fluid inclusions in these samples mostly have microthermometric properties comparable to EM2, suggesting that
the mixing zone may have risen above the granite at
times, but at higher levels was dominated by EM2.
Within these samples O'Connor et al. (1993) also ®nd
dilute (1±8 wt% NaCl + CaCl2), low-temperature
(100 °C) ¯uids in secondary inclusions which may
represent contemporaneous surface ¯uid in®ltrating the
topmost part of the mineralisation system.
The Upper Triassic hydrothermal event
in a wider context
The Upper Triassic mineralisation in Connemara is
broadly synchronous with, and has features in common
with, other Triassic±Jurassic hydrothermal activity and
mineralisation throughout the north Atlantic margins
and central and southern Europe; a hydrothermal
province ®rst identi®ed by Mitchell and Halliday (1976).
This mineralisation is characterised by base metals ‹ Ag with quartz + carbonate + ¯uorite + barite, generally involved CaCl2-rich brines of moderate±
high salinity, had ¯uid TH usually 70±200 °C, often has
features suggesting repeated mixing of ¯uids and often
occurs around the margins of Mesozoic basins, particularly in the vicinity of older granites (Mitchell and
Halliday 1976 and references therein; Halliday and
Mitchell 1984 and references therein and e.g. Behr et al.
1987; Behr and Gerler 1987; Charef and Sheppard 1988;
Canals and Cardellach 1993; Munoz et al. 1994; Wil-
139
kinson et al. 1995). For example, in Cornwall, Triassic?
(Scrivener et al. 1994) Pb-Zn mineralised ``cross-courses'' show mixtures of two ¯uids which migrated into the
area synchronously. One was a hot (150±250 °C),
low-salinity (0±6 wt% NaCl eq.) ¯uid with dD = )21 to
)25& and d18O = )0.3 to +7.4&, interpreted to be of
meteoric origin, the other was a cooler (110±150 °C),
high-salinity (23±28 wt% NaCl eq.), Ca-rich ¯uid with
dD = )50 to )80& and d18O = )0.3 to +4.7&, interpreted to be of Permo-Triassic basinal origin (Wilkinson et al. 1995).
A number of authors have attributed this pulse of
Triassic±Jurassic mineralisation to early north Atlantic
rifting (Mitchell and Halliday 1976; Gleadow 1978;
Halliday and Mitchell 1984; O'Connor et al. 1993).
Halliday and Mitchell (1984) suggest that rifting caused
rapid subsidence, thinning and fracturing of the continental crust and possible rise of mantle-derived magma.
These factors, they suggest, would increase permeability
and heat ¯ow, initiating hydrothermal activity at the
Mesozoic basin margins. The relationship between
mineralisation and older granites they attribute to ¯uid
¯ow being focused through the granites because they are
generally hotter than surrounding crust. We concur with
the latter suggestion, but add that the spatial correlation
of mineralisation with granites may also be partly because they form stable blocks adjacent to the margin of
Mesozoic basins, which in many cases (but not Connemara) were formed by post-orogenic collapse of the belt
within which the granites formed. Other factors which
will tend to localise mineralisation around older granites
may be the intense fracturing and jointing, which allow
permeability to depth, and the fact that the granite itself
can act as the source for some of the mineralising constituents. Halliday and Mitchell's (1984) rifting-related
model for Mesozoic mineralisation now needs to be
extended to explain both the involvement of saline
brines, and the mixing of basinal and surface ¯uids. The
involvement of saline brines is ultimately related to the
widespread development of evaporites in Permo-Triassic
sediments and (in Connemara) the Lower Carboniferous
due to the equatorial±tropical position of Europe during
this period. An explanation for the mixing of two ¯uids
is more challenging, since it must involve both expulsion
(compression?) of brines from basins, and, contemporaneously, tension within basement allowing deep penetration of surface ¯uids.
Concluding remarks
Three ¯uids have a€ected the Galway Granite. The
earliest was a magmatic ¯uid derived from the granite
during the ®nal stages of solidi®cation, which formed
minor molybdenite mineralisation. The second was a
meteoric ¯uid which caused the widespread retrograde
alteration previously identi®ed in the area. This ¯uid was
probably drawn into the granite in a meteoric convection system driven by magmatic heat soon after intru-
sion. However, the need to explain the alteration of
Upper Carboniferous dikes, means that this ¯uid could
have been considerably younger, either Upper Carboniferous or Upper Triassic. No signi®cant mineralisation
is associated with this ¯uid at the present level of erosion. The last ¯uid to a€ect the granite deposited
Pb ‹ Cu mineralised quartz, ¯uorite, calcite and barite
veins of late Triassic age. This ¯uid was formed by
mixing of two ¯uids within the granite, one a contemporaneous surface-derived ¯uid, the other a basinal
brine derived from the Carboniferous sequence.
This study demonstrates that the Galway Granite was
a locus for repeated ¯uid alteration and mineralisation,
ranging in age from late-magmatic to post-orogenic.
These ¯uid events di€er in cause and mechanism, but are
all related in some way to properties of the granite.
Earliest ¯uids were magmatic ¯uids generated at the ®nal stages of crystallisation, later ¯uid in®ltration was in
a convection system, most probably driven by magmatic
heat. Youngest mineralisation is unrelated to magmatic
heat, but may be localised in and around the granite by
virtue of its generally high heat ¯ow, the fact that it
forms a stable permeable block adjacent to later sedimentary basins from which mineralising solutions can be
expelled by tectonic activity, and the fact that it may act
as source rocks for some constituents of the mineralisation. This sequence of ¯uid events may be the general
case for collisional granitoids and underlines the polygenetic nature of mineralisation spatially related to
them.
Acknowledgements A. Boyce, P. Gorman and C. Ford for d34S
analyses. Z. Duan for some calculations. J. Dubessy and A. Rankin
for probing reviews. C.O'R. acknowledges a Basic Research Award
from EOLAS and a post-graduate fellowship from University
College Galway. The Isotope Geosciences Unit at SURRC is
supported by NERC and the Scottish Consortium of Universities
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