JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
PAGES 1617–1650
2002
Water-deficient Calc-alkaline Plutonic Rocks
of Northeastern Superior Province, Canada:
Significance of Charnockitic Magmatism
JOHN A. PERCIVAL1∗ AND JAMES K. MORTENSEN2
1
GEOLOGICAL SURVEY OF CANADA, 601 BOOTH STREET, OTTAWA, ONTARIO, K1A 0E8, CANADA
2
DEPARTMENT OF EARTH AND OCEAN SCIENCES, UNIVERSITY OF BRITISH COLUMBIA, 6339 STORES ROAD,
VANCOUVER, B.C., V6T 1Z4, CANADA
RECEIVED JULY 24, 2001; REVISED TYPESCRIPT ACCEPTED MARCH 1, 2002
Calc-alkaline batholiths of the Archaean Minto block, northeastern
Superior Province, Canada, have pyroxene- and hornblende-bearing
mineral assemblages inferred to have crystallized from hot, waterundersaturated magmas at 2·729–2·724 Ga. A regional amphibolite- to granulite-facies tectonothermal event at 2·70 Ga
resulted in mild to negligible metamorphic effects on the dominantly
granodioritic units. Geochemical, textural and thermobarometric
studies define the crystallization history in compositions ranging
from cumulate pyroxenite through quartz diorite, granodiorite, granite,
and syn-magmatic gabbroic dykes. Early magmatic assemblages
include orthopyroxene, clinopyroxene, plagioclase, biotite, Fe–Ti
oxides and ternary feldspar, indicating crystallization from magmas
containing <2 wt % H2O at 1100–900°C. Water enrichment
in the residual melt induced hornblende crystallization at 5 ±
1 kbar, 800–600°C. Characterized by a continuum of large ion
lithophile element (LILE)-enriched, high field strength element
(HFSE)-depleted compositions, the I-type suite resembles modern
continental arc batholiths in composition and size but not primary
mineralogy. Magmatic arcs produced between 2·75 and 1·85 Ga
commonly have charnockitic components, possibly because slabderived fluids interacted with mantle wedges at ambient temperatures
higher by >100°C than at present, producing large volumes of
water-deficient magma.
INTRODUCTION
granitoid rocks; igneous pyroxenes; water-undersaturated
magma; charnockite
Since Holland’s (1900) initial description of south Indian
charnockites as orthopyroxene-bearing rocks of granitic
composition, the term charnockite has been applied to
rocks of widely divergent origin: granitic rocks metamorphosed to the granulite facies (metamorphic charnockites); and rocks whose pyroxene crystallized
directly from a magma (igneous charnockites). Unmetamorphosed, discrete plutons such as the Barrington Tops
batholith (Eggins & Hensen, 1987), Kleivan granite (Petersen, 1980), Mawson charnockite (Young & Ellis, 1991;
Young et al., 1997; Zhao et al., 1997) and Ballachulish
complex (Weiss & Troll, 1989) contain pyroxene of
undisputed igneous origin. However, in high-grade terranes that contain both granulite-facies metamorphic rocks
and pyroxene-bearing granites, the origin of pyroxene in
granites may be less evident (Newton, 1992; Percival,
1994). By virtue of their large abundance, these units
can be inferred to have played a significant role in the
evolution of cratonic crust.
Several complementary approaches can be used to
distinguish igneous and metamorphic origins of pyroxenebearing granitoids. Field observations such as dykes of
pyroxene-bearing granite cutting amphibolite-facies rocks
(Frost & Frost, 1987; Bohlender et al., 1992) argue for
hot, dry magmas. Similarly, textures that indicate the
relative crystallization order of pyroxenes and hydrous
ferromagnesian phases serve to distinguish prograde
metamorphic from igneous histories. Thermobarometry
can provide clues to origin, for example where relict
∗Corresponding author. E-mail: [email protected]
 Oxford University Press 2002
KEY WORDS:
JOURNAL OF PETROLOGY
VOLUME 43
phases preserve temperatures >1000°C and must be
considered in the light of possible igneous processes (e.g.
Bohlen & Essene, 1978; Rollinson, 1982). Geochronology
may establish different generations of mineral growth
related to igneous and metamorphic events.
Pyroxene-bearing granitoid rocks and their hydrous
equivalents constitute a large part of the 500 km ×
500 km Minto block of the northeastern Superior Province of Canada (Fig. 1). Pyroxene was initially considered
to be of metamorphic origin based on reconnaissance
investigations (Stevenson, 1968). Herd (1978) supported
this conclusion, regarding pyroxenes to have formed
during a granulite-facies M1 metamorphism and hornblende and biotite during a retrogressive M2 event. In
contrast, recent studies have concluded that many of
these rocks are essentially unmetamorphosed, containing
early igneous pyroxene and late igneous amphibole (Percival et al., 1990, 1992, 2001; Shore, 1991; Bégin &
Pattison, 1994; Stern et al., 1994). In this paper we
document the age, geochemical and mineral chemical
characteristics of pyroxene-bearing granitoid suites of the
Minto block and discuss implications for magma genesis.
REGIONAL GEOLOGICAL SETTING
The Superior Province consists of two distinct regions
(Fig. 1): a southern block made up of alternating, easttrending, relatively low-grade greenstone and metasedimentary subprovinces; and a northeastern (Minto)
block, consisting essentially of granitoid and high-grade
metamorphic rocks with northerly structural and aeromagnetic trends (Card, 1990; Percival et al., 1992, 1996,
2001; Labbé et al., 1998; Pilkington & Percival, 1999).
In the south, the distribution of belts has been attributed
to successive lateral accretion of juvenile oceanic terranes,
microcontinents and collisional sedimentary prisms in
the interval 2·72–2·70 Ga (Card, 1990; Corfu & Davis,
1992; Williams et al., 1992). (All quoted ages are U–Pb
zircon dates unless otherwise indicated.)
Recent work in the Minto block has recognized a series
of orogenic events between 2·81 and 2·70 Ga, providing
linkages to areas to the south and west (Percival &
Skulski, 2000). Scattered remnants of ancient (2·9–3·0
Ga; Percival et al., 2001) crust occur in the east, in
the Goudalie and Douglas Harbour domains, possibly
representing an orogenic foreland during subsequent
tectonic events. Volcanic and associated rocks (2·84–2·83
Ga) in the Qalluviartuuq domain have juvenile Nd isotopic signatures (Skulski et al., 1996), suggesting oceanfloor and arc settings. Early (>2·81 Ga) shear zones
bounding distinct tectonostratigraphic packages may represent intraoceanic accretionary structures (Percival &
Skulski, 2000). The early collage was cut by a suite of
calc-alkaline plutons associated with minor volcanic rocks,
NUMBER 9
SEPTEMBER 2002
dated at 2775 Ma (Skulski et al., 1996), and the composite
2·84–2·77 Ga basement overlain unconformably by
<2748 Ma sedimentary rocks including conglomerate,
iron formation and greywacke (Percival & Skulski, 2000).
The metasedimentary units are probably correlative with
high-grade schists and paragneisses of the Lake Minto
domain to the west (Fig. 1). Calc-alkaline plutonic rocks
of 2730–2720 Ma age are widespread throughout the
Utsalik and Lake Minto domains (Percival et al., 2001)
and volcanic rocks of similar age occur in the Vizien belt
(Percival et al., 1994), representing voluminous continental
arc magmatism (Stern et al., 1994). In >2725 Ma granodiorites of the Utsalik domain, Nd values of +1·1 to
−0·5 indicate significant involvement of older crust,
whereas in the Lake Minto domain to the west, rocks
of similar age and composition are relatively juvenile
( 2725Nd = +0·1 to +1·3). The youngest supracrustal
rocks are conglomerate and greywacke of the Vizien belt,
with clasts <2718 Ma (Percival & Card, 1994; Lin et al.,
1995; Skulski & Percival, 1996). Calc-alkaline plutonic
rocks of the Tikkerutuk domain (2712–2702 Ma; Percival
et al., 2001), represent renewed arc magmatism. Eastward
thrusting of this active arc over the Lake Minto domain
resulted in burial, deformation, and metamorphism to
the amphibolite and granulite facies (3·5–10 kbar,
575–900°C, 2702 Ma; Bégin & Pattison, 1994; Percival
& Skulski, 2000), as well as production and emplacement
of crustally derived plutons including diatexites
(2696–2693 Ma). Still younger magmatism included
2688 Ma orthopyroxene granite (Stern et al., 1994; Percival et al., 2001), 2675 Ma granite (Percival & Skulski,
2000), and 2660 Ma pegmatites (Percival & Card, 1994).
Late-stage growth of monazite in supracrustal belts
(2688–2628 Ma; Percival & Skulski, 2000) and zircon
overgrowths (Percival et al., 2001) may reflect circulation
of metamorphic, magmatic and hydrothermal fluids.
ANALYTICAL TECHNIQUES
All data were acquired in laboratories of the Geological
Survey of Canada (Ottawa). Geochemical data were
obtained from whole-rock powders. Major elements were
analysed on fused discs by X-ray fluorescence (XRF),
and trace elements, including rare earth elements, by
inductively coupled plasma mass spectrometry (ICP-MS).
Errors are estimated at ±5% for major elements and
±10% for trace elements.
Zircon concentrates were prepared from 20 kg samples
using conventional crushing, grinding, Wilfley table,
heavy liquids and Frantz magnetic separator techniques.
The techniques for zircon grain selection, abrasion, dissolution, geochemical preparation and mass spectrometry
have been described by Parrish et al. (1987). All zircon
1618
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 1. Generalized geological map showing distribution of major rock units and domains in the central Minto block. Zircon dates (Percival et
al., 1992; Percival & Card, 1994; Skulski & Percival, 1996) are from igneous grains with errors generally <±3 Ma. Inset map (lower right) shows
Superior Province with subprovince boundaries and location of study area in the NE.
fractions were air abraded (Krogh, 1982) before dissolution to minimize the effects of post-crystallization Pb
loss. Procedural blanks were 10–22 pg for Pb and from
1 to 3 pg for U.
Errors assigned to individual analyses were calculated
using the numerical error propagation method of Roddick
(1987). Decay constants used are those recommended
by Steiger & Jäger (1975), and compositions for initial
common Pb were taken from the model of Cumming &
Richards (1975). Regressions of discordia arrays were
carried out using the model of Davis (1982). All errors
are given at the 2 level.
Mineral chemical analyses were obtained on a Camebax electron microprobe equipped with wavelength-dispersive spectrometers. Accelerating voltage was 15 kV,
with specimen current varying from 10 to 30 nA, depending on mineral type. Counting times varied from
10 to 30 s per element, for total count times of >100 s
per spot. Standards include a variety of minerals, oxides
and metals, and data reduction was performed using
routines provided by Pouchou & Pichoir (1984). Analytical reproducibility is of the order of ±5% for most
elements. Structural formulae were calculated using programs developed by G. J. Pringle.
MINERALOGY AND GEOCHEMISTRY
OF PYROXENE-BEARING PLUTONS
Four suites of coarse-grained, pyroxene-bearing granodioritic rocks have been identified in the Minto block,
through reconnaissance and detailed petrological and
geochronological studies (Leclair et al., 2001; Percival et
al., 2001). Many bodies vary internally from pyroxenedominated to hornblende-, biotite-dominated mafic
mineral assemblages and therefore compositional
terms with mineral modifiers are used (e.g. orthopyroxene–clinopyroxene–biotite granodiorite), rather
than nomenclature specific to pyroxene-bearing rocks
(e.g. charnockite, mangerite, opdalite, etc.). The term
‘charnockitic’ is used in this paper for general reference
to igneous pyroxene-bearing granitic rocks.
A suite of >2·78 Ga calc-alkaline plutons in northern
Lake Minto domain represents the oldest pyroxenebearing granodiorites recognized to date (Skulski et al.,
1996). A younger group of calc-alkaline, pyroxene- and
hornblende-bearing granodiorites of the Lake Minto and
Utsalik domains forms the main focus of this study. Based
on common ages of 2725 ± 5 Ma, these rocks were
grouped as the ‘Leaf River suite’ (Stern et al., 1994).
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JOURNAL OF PETROLOGY
VOLUME 43
The linear Tikkerutuk domain of pyroxene-bearing calcalkaline granodiorite and granite has ages in the range
2710–2693 Ma (Percival et al., 2001). Orthopyroxenebearing peraluminous granodiorite (diatexite) in the Lake
Minto domain (2696 Ma) may have originated through
crustal melting resulting from collisional tectonism at
2700 Ma (Percival & Skulski 2000).
Utsalik domain plutons
The Utsalik domain is characterized by north-striking
sheets of granodiorite and monzogranite on the 1–10 km
scale (Percival & Card 1994). Pods and enclaves include
pyroxenite, gabbro and diorite in bodies of 10 cm to
30 m scale. Pyroxenites consist of assemblages of orthopyroxene, clinopyroxene and biotite, with accessory
olivine, spinel, Fe–Ti oxides, hornblende and plagioclase. Gabbro and diorite have common clinopyroxene–biotite–plagioclase ± orthopyroxene assemblages.
Granodiorite is medium to coarse grained, homogeneous and massive to weakly foliated, with several
mineral facies (Fig. 2a–c). Mafic mineral assemblages of
orthopyroxene, clinopyroxene and biotite (Fig. 3a), all
of apparent igneous origin, grade through facies with
clinopyroxene cores in hornblende (Fig. 2b), to massive,
homogeneous rocks containing hornblende and biotite
(Fig. 2c). Hornblende-bearing enclaves commonly have
centimetre-scale orthopyroxene-rich margins where in
contact with pyroxene-bearing granodiorite (Fig. 2d).
Pyroxene-rich selvages are also observed where dykes of
pyroxene-bearing granodiorite transect mafic rocks (Fig.
2e; Ross, 1991). Sheets and plutons of granite with
similar mineral assemblages to those in granodiorite have
common accessory allanite. Gabbro and diorite occur
as 1–5 m sheets or dykes in granodiorite and granite.
The dykes generally consist of assemblages of
clinopyroxene–biotite–plagioclase, are straight walled,
massive to weakly foliated and boudinaged (Percival et
al., 1992). Some dykes have cuspate margins (Fig. 2f )
and are back-veined by granodiorite, suggesting synmagmatic emplacement (Shore, 1991; see Blundy &
Sparks, 1992).
In quartz diorite, granodiorite and quartz monzodiorite, pyroxenes occur as randomly oriented subhedral crystals up to 5 mm, along with plagioclase
(An11–48), quartz, alkali feldspar, ilmenite, magnetite and
accessory apatite, zircon and rare monazite; biotite occurs
as fox-red grains (Fig. 3a and b). Clinopyroxene rarely
has 001 lamellae of possible pigeonite (Fig. 3c) (see
Ollila et al., 1988). Orthopyroxene rarely exhibits internal
exsolution whereas clinopyroxene commonly has fine
(5–20 m) 100 exsolution lamellae in the cores of some
grains, with unexsolved rims (Fig. 3d–f ). Orthopyroxene
NUMBER 9
SEPTEMBER 2002
and ilmenite form the principal exsolved phases (Fig.
3f ). With increasing development of hydrous matrix
phases, patches of hornblende and biotite develop as
overgrowths and along 100 planes in pyroxenes. The
granodiorites grade through zones with pyroxene cores in
hornblende (Fig. 4a–d) to common hornblende-, perthitebearing assemblages in granodiorites and granites. Magnetite is generally abundant, giving these rocks high
magnetic susceptibility and in turn producing an intense
regional aeromagnetic anomaly (Pilkington & Percival,
1999, 2001).
New chemical analyses of plutonic rocks of the Utsalik
domain (Fig. 5) are reported in Table 1 and presented
with complementary data from Stern et al. (1994) in Figs
6–10. Sample locations and mineral assemblages are
listed in the Appendix. The suite can be divided into
eight sub-units based on mode of occurrence and mineral
assemblages. Mafic rocks include gabbroic and dioritic
enclaves, as well as two chemically distinct types of synplutonic dykes. Five suites of main-phase plutonic rock
are distinguished on the basis of their mafic mineral
assemblage: (1) orthopyroxene–clinopyroxene–biotite ±
hornblende diorite to granodiorite; (2) clinopyroxene–
biotite diorite to granodiorite; (3) clinopyroxene–
hornblende–biotite diorite to granodiorite; (4)
hornblende–biotite granodiorite; (5) biotite granodiorite
to granite. Together, the Utsalik units represent silica
contents in the 45–73 wt % SiO2 range and form coherent
major and trace element trends on Harker diagrams (Fig.
6). They are dominantly calc-alkalic, with a few calcic
and alkali–calcic compositions using the classification of
Frost et al. (2001).
Gabbroic enclaves are characterized by variable compatible-element (Mg, Cr, Ni, V) and alumina contents
within a narrow (49–52%) silica range (Fig. 6). Some
rocks have low Al2O3 and high CaO contents suggesting
the presence of cumulate clinopyroxene. A second suite
of enclaves, of gabbroic to dioritic composition (49·1–
58·6 wt % SiO2), forms collinear and continuous trends
with more evolved rocks on variation diagrams (Fig.
6). The enclaves have orthopyroxene–clinopyroxene–
biotite–plagioclase assemblages in common with granodiorites and may represent early crystallized units approaching parental compositions.
Mafic dykes also plot on compositional trends defined
by the more evolved rocks. Some group 1 dykes display
widely variable compatible-element contents, suggesting
fractionation before emplacement. The second group of
dykes has relatively low MgO contents (<5 wt %) and
high K2O, P2O5, TiO2, Zr and Ba (Table 1; Fig. 6).
Three of five dykes have slight negative Eu anomalies.
On a primitive-mantle-normalized extended-element
profile (Fig. 7), the mafic rocks have patterns characterized by enrichment in light rare earth elements
(LREE) (La/Ybn = 1·2–43) and large ion lithophile
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PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 2. Field photographs of pyroxene-bearing and associated rocks: (a) orthopyroxene–clinopyroxene–biotite granodiorite showing coarse grain
size and homogeneous, massive character; (b) transitional granodiorite: pale patches are rich in perthite, with dark hornblende; matrix is
orthopyroxene–clinopyroxene–plagioclase–quartz; (c) hornblende–biotite granite with cores of clinopyroxene (pale) in hornblende; (d) enclaves
of hornblende-bearing gabbro in coarse orthopyroxene–clinopyroxene granodiorite have orthopyroxene-rich selvages; (e) dyke of pyroxenebearing granite cutting hornblende-bearing mafic gneiss (note orthopyroxene-rich reaction selvage); (f ) syn-plutonic clinopyroxene–biotite diorite
dyke in orthopyroxene–clinopyroxene–biotite granodiorite (note boudinaged geometry).
elements (LILE), as well as negative Nb, Zr and Ti
anomalies. The compositions correspond well to those
of calc-alkaline basalt and andesite formed in modern
arc environments (see Rollinson, 1993). The high LILE
contents and element depletions in these relatively primitive members of the suite, which also characterize the
more evolved compositions, suggest that the trace element
patterns may reflect a mantle wedge signature.
The Utsalik plutonic domain is dominated by rocks of
granodioritic and granitic composition. These form linear
trends on Harker plots of compatible elements (MgO,
Fe2O3, V, Ni, Cr, etc.), that decrease systematically with
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JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
SEPTEMBER 2002
Fig. 3. Photomicrographs of pyroxene-bearing plutonic rocks: (a) medium-grained diorite showing random orientation of orthopyroxene (O),
clinopyroxene (C), magnetite (M) and plagioclase (P); minor biotite (B) is not associated with pyroxenes (sample U-12); (b) medium-grained
quartz diorite showing randomly oriented pyroxenes (O, C), plagioclase (P) and biotite (B) in textural equilibium (PBA91-58); (c) clinopyroxene
(C) in granodiorite showing 001 lamellae of possible pigeonite (indicated by arrows) and faint 100 lamellae (orthopyroxene?); matrix phases are
biotite (B), plagioclase (P) and quartz (Q ) (PBA91-6); (d) clinopyroxene (C) in granodiorite showing 100 exsolution lamellae of orthopyroxene in
grain core (formerly augite) and unexsolved rims (diopside) (U-18); (e) granodiorite showing clinopyroxene (C) overgrown by biotite (B) (note
concentration of exsolution lamellae in core of grain and unexsolved rims) (U-3); (f ) detail of relationship between exsolved (E) region in augitic
core of clinopyroxene grain and clear (C) diopsidic rim (U-1). High-temperature igneous crystallization followed by orthopyroxene exsolution in
sub-calcic portions is implied from textures illustrated in (d), (e) and (f ). All photomicrographs are taken in plane-polarized light except (d)
(crossed nicols).
silica (Fig. 6). Plots of incompatible elements show more
scatter (e.g. K2O, Rb, Th, Ba) but generally increase
with silica. Both K2O and Al2O3 have inflection points
in the 60% SiO2 range. Alumina contents decline slightly
with silica above >60% SiO2, whereas potash levels rise.
These compositional changes appear to correspond to
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PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 4. Photomicrographs of hornblende-bearing plutonic rocks: (a) partial hornblende (H) rims around clinopyroxene; orthopyroxene (O) is
unaltered (PBA91-143); (b) complete rims of hornblende (H) around clinopyroxene (C) and orthopyroxene (O) in a plagioclase-rich (P) matrix
(U-13); (c) large twinned hornblende (H) grain with inclusions of clinopyroxene (C), plagioclase (P) and magneite (M) (U-14); (d) hornblende
(H)–perthite (Pe) association in granite and granodiorite; quartz (Q ) and plagioclase (P) are also present (GS90-1). All photomicrographs are
taken in plane-polarized light except (d) (crossed nicols).
the appearance of hornblende in the mafic mineral
assemblage, in addition to orthopyroxene, clinopyroxene
and biotite.
Rare earth patterns (averages shown in Fig. 8) vary
with bulk composition and mineral assemblage. All have
sinusoidal patterns with concave-down LREE, concaveup heavy rare earth element (HREE) profiles (La/Ybn =
13–197) and negligible Eu anomalies. However, the levels
of LREE enrichment and HREE depletion increase
systematically
from
orthopyroxene–clinopyroxene,
through clinopyroxene–hornblende and hornblende–
biotite, to biotite-only assemblages (Fig. 8), probably also
reflecting higher silica contents (respectively 61·2, 65·4,
67·4 and 71·5 wt % SiO2).
On primitive-mantle-normalized extended-element
plots, the intermediate and felsic rocks show strong enrichment in LILE as well as in LREE (Figs 7 and 9).
Pyroxene-bearing rocks have comparable concentrations
of both compatible and incompatible trace elements to
hornblende-, biotite-bearing units (Fig. 9). Both suites
have negative Nb and Ti anomalies and slight Zr enrichment. Thorium is systematically higher in the hornblende-bearing and biotite-only rocks, consistent with its
partitioning into hydrous magma (e.g. Wood & Blundy,
2001).
The geochemical features are consistent with fractional
crystallization of a basaltic calc-alkaline magma resembling the mafic end-members of the suite. Some
crustal contamination is indicated by 2725Nd values in the
range −0·5 to +1·1 (Stern et al., 1994). Using known
ages of potential contaminants (2·8–3·0 Ga; Stern et al.,
1994; Percival et al., 2001), the volume of older crust
assimilated before fractionation is in the range 10–20%.
Mass-balance considerations would predict a large mass
of complementary cumulates.
Lake Minto domain plutons
Plutons of Lake Minto domain occur as kilometre-scale
sheets, separated by screens of supracrustal rock and
1623
50·9
0·40
16·4
8·6
0·14
8·86
10·90
2·4
0·54
0·07
99·2
210
180
26
110
5·3
17
160
370
2·2
1·9
28
11
2·1
0·27
11
22
11
2·1
0·72
2·1
1·7
0·38
1·0
0·17
1·1
53·8
0·84
17·5
9·6
0·13
3·86
6·65
3·9
2·41
0·27
99·0
21
25
16
170
21
75
890
780
9·4
3·2
120
23
7·8
0·36
39
83
40
7·1
1·4
5·5
4·0
0·83
2·1
0·35
2·0
49·1
1·19
18·6
12·0
0·13
4·07
7·65
4·0
1·85
0·47
99·1
20
32
18
180
12
93
630
970
13
3·3
120
33
2·8
0·86
40
98
58
11·0
2·3
8·5
5·9
1·1
3·0
0·53
2·9
U-28
gbr
encl
OCB
U-12
dte
encl
OCB
U-30
dte
encl
H
49·1
50·5 58·6
1·06
1·37 0·65
16·5
17·1 16·0
11·2
11·9
8·5
0·16
0·16 0·11
6·33
4·55 3·86
9·41
7·85 5·56
3·6
3·7
4·2
1·67
1·64 1·79
0·29
0·55 0·15
99·3
99·3 99·4
92
22
36
52
23
48
29
21
13
200
180
82
11
8·2 15
75
52 110
580
1000 350
570
770 410
10
12
12
3·2
5·4
3·9
110
260 110
22
22
18
1·3
0·69 10
0·32
0·23 2·9
34
45
28
77
90
55
35
45
23
5·6
7·2
4·0
1·6
2·2
1·1
4·5
6·1
3·4
3·7
4·1
2·8
0·75
0·83 0·6
2·0
2·1
1·6
0·35
0·34 0·26
2·1
1·8
1·7
U-29
gbr
encl
CB
U-32
gdi
main
OCHB
U-35
gdi
main
OCHB
58·0
62·5 57·1
0·70
0·46 1·48
16·6
17·0 16·6
7·3
5·3 10·0
0·09
0·05 0·14
3·56
1·85 2·42
5·72
3·54 6·27
4·4
4·2
4·5
1·55
3·59 1·03
0·21
0·18 0·51
98·1
98·7 100·1
52
32
12
33
17
12
15
5·4 17
120
81
92
15
22
15
67
120
17
460
1000 410
620
460 580
9·4
6·5 20
4·1
5·2 10
160
220 530
17
11
29
1·7
8·7
2·8
0·26
0·58 0·51
39
42
48
79
77 110
33
28
56
5·4
4·1 10·0
1·1
1·1
2·6
3·8
2·7
8·4
2·9
1·8
5·6
0·58
0·37 1·1
1·4
0·82 2·6
0·26
0·15 0·39
1·6
0·79 2·1
U-31
gdi
main
OCB
U-11
gdi
main
OCHB
U-37
gdi
main
OCHB
U-38
gdi
main
CB
U-43
gdi
main
CHB
56·8
62·5 56·1
66·4 69·1
0·82
0·95 1·54
0·36 0·43
16·1
15·8 14·0
15·2 14·6
6·9
6·3 12·5
4·1
4·5
0·09
0·08 0·19
0·05 0·04
4·18
1·34 3·08
1·65 1·60
5·76
3·83 6·67
3·07 3·58
4·3
3·6
3·9
3·4
4·3
1·55
3·77 0·82
4·48 1·44
0·15
0·26 1·07
0·15 0·11
96·7
98·4 99·9
98·9 99·7
62
10 <10
25
24
29
<10 <10
13
15
20
12
20
7·9
4·1
110
63
89
58
70
16
21
12
21
28
67
73
16
120
81
340
1400 250
1300 220
490
310 460
480 460
12
14
17
3·5
6·9
3·8
7·2
3
3·8
3·8
100
320 110
150 120
24
17
43
12
6·1
6·8
2·4
2·8
4·5 68
0·97
0·81 0·48
0·33 0·6
32
37
78
33
52
62
68 160
66 100
29
32
83
26
27
5·9
5·5 15·0
4·1
2·6
0·82
1·7
2·1
0·93 0·52
5·6
4·5 12·0
2·9
1·0
4·4
3·1
8·1
2·2
0·81
0·9
0·61 1·7
0·4
0·18
2·4
1·5
4·0
1·0
0·49
0·4
0·23 0·66
0·17 0·12
2·1
1·3
3·4
1·0
0·71
U-36
gdi
main
OCHB
66·4
0·47
14·5
5·0
0·07
1·96
3·30
4·3
2·66
0·15
98·8
29
18
7·7
56
24
88
500
470
10
6·4
290
26
27
0·8
80
150
62
9·4
1·5
6·0
4·4
0·88
2·2
0·38
2·3
U-44
gdi
main
CHB
U-45
gdi
main
CHB
60·8
66·8
0·72
0·51
16·3
15·5
6·9
4·5
0·11
0·06
3·14
1·92
4·93
3·93
4·1
4·2
2·48
1·44
0·20
0·18
99·7
99·0
39
40
24
21
15
4·4
89
33
25
15
120
40
380
460
420
540
14
4·2
5·1
5·1
200
210
19
9·3
32
1·5
1·1
0·7
71
31
130
52
49
20
6·8
3·1
1·3
0·89
4·4
2·4
3·3
1·7
0·65
0·33
1·7
0·81
0·29
0·12
1·7
0·65
U-7
gdi
main
CB
U-51
gdi
main
HB
U-52
grnt
main
B
69·1 63·3 71·5
0·54 0·99 0·32
14·6 15·7 13·6
3·9
6·6 2·2
0·08 0·13 0·02
0·71 1·25 0·97
2·18 3·50 1·52
4·2
4·7 2·8
3·62 2·95 5·39
0·14 0·23 0·10
99·1 99·4 98·4
<10
<10
15
<10
<10
14
7·4 12
1·6
11
23
27
18
35
33
84
110 160
2300 1400 940
380
470 400
12
13
5·9
3·7
4·7 3·4
140
270 120
11
27
6·1
1·8
8·2 56
0·42 1·2 0·63
20
43 110
37
86 190
17
40
60
3·1
7·2 5·3
2·6
2·6 0·85
2·7
6·4 1·6
2·1
5·0 1·1
0·42 0·97 0·19
1·0
2·7 0·38
0·19 0·45 0·08
1·1
2·6 0·4
U-4
gdi
main
OCHB
VOLUME 43
1624
NUMBER 9
Lithology and mineral abbreviations as in the Appendix.
49·4
45·6 49·8
0·75
1·49 2·37
16·7
7·6 15·8
12·1
16·0 14·2
0·13
0·23 0·15
7·04
16·90 4·70
8·36
6·89 7·40
3·1
0·6
3·8
1·48
2·28 1·23
0·17
0·10 0·56
99·2
97·7 100·0
74
1300
35
85
810
44
16
41
16
210
190 150
8·2
4·8
9·3
59
160
27
510
210 470
910
65 680
2·9
32
27
1·9
3·7
5·6
52
92 240
8·4
98
22
2·8
1·9
0·23
1·2
0·57 0·11
17
16
37
37
53
81
17
47
46
3·3
17·0
9·0
0·96
1·1
2·5
2·5
24·0
7·5
1·5
24·0
4·6
0·28
4·9
0·85
0·65
12·0
2·0
0·1
2·0
0·31
0·63
9·7
1·5
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
Cr
Ni
Sc
V
Pb
Rb
Ba
Sr
Nb
Hf
Zr
Y
Th
U
La
Ce
Nd
Sm
Eu
Gd
Dy
Ho
Er
Tm
Yb
U-18 U-19 U-23 U-27
gbr
gbr
gbr
dte
dyke 1 dyke 1 dyke 1 encl
B
OCH OCH OCB
U-15
gbr
encl
H
Sample:
Lithology:∗
Setting:
Min. ass.:
Table 1: Chemical compositions of Utsalik plutonic rocks
JOURNAL OF PETROLOGY
SEPTEMBER 2002
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 5. Locations of samples analysed in this study, along with pyroxene occurrences based on foot traverses and spot-checked petrographically.
Owing to the coarse grain size of all units, field identification of assemblages is generally reliable.
younger plutonic sheets. Pyroxenite, hornblendite, gabbro, and diorite occur as pods up to 40 m in size.
Boudinaged dykes of diorite and gabbro are also common.
Rare tonalite gneiss enclaves have yielded ages >3·1 Ga.
Granodiorite, the main plutonic phase, is medium to
coarse grained, homogeneous and weakly foliated to
massive. Mafic mineral assemblages range from orthopyroxene–clinopyroxene–biotite with or without late
hornblende, to hornblende–biotite.
No new geochemical results are reported here; however, analyses of Lake Minto domain rocks reported by
Stern et al. (1994) are shown for comparative purposes
in Fig. 10. In comparison with the coeval (2725 Ma)
Utsalik domain plutonic suite, granodiorites of the Lake
Minto domain have systematically lower Th and U
contents, and are more depleted in the HREE (Fig. 10a).
The fractionated HREE profile suggests the presence of
residual garnet in the source (Stern et al., 1994). Lake
Minto granodiorites show several geochemical features
comparable with those of modern adakites, such as high
Sr/Y, low Y (Fig. 10b) and La/Yb up to 260, which are
commonly considered to result from the presence of a
slab melt component (Drummond & Defant, 1990;
Martin, 1994; Schiano et al., 1995). In contrast, Utsalik
plutonic rocks fall mainly within fields defined by ‘normal’
arcs (see Castillo et al., 1999; Fig. 10b). In view of the
gradation from adakitic to normal arc signatures (Fig.
10b), it is possible that the Lake Minto and Utsalik suites
represent magmas derived from a single mantle wedge,
fluxed respectively by slab melts and slab-derived fluids.
GEOCHRONOLOGY
U–Pb data are listed in Table 2 and plotted on conventional U–Pb concordia diagrams in Fig. 11.
Utsalik domain
Sample U-34 (orthopyroxene–clinopyroxene–biotite
granodiorite)
Four fractions of zircon and two of monazite were
analysed (Fig. 11a). The zircons form stubby euhedral
prisms with faint internal growth zoning and no evidence of metamorphic overgrowth. One fraction (B) is
1625
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
SEPTEMBER 2002
Fig. 6. Harker diagrams showing variation of compatible and incompatible element contents with silica. Naney granodiorite and granite from
Naney’s (1983) experimental runs; average I-type (n = 991) and S-type (n = 578) from Whalen et al. (1987): C-type (n = 12) from Kilpatrick
& Ellis (1992); M-type (n = 8) from Eggins & Hensen (1987).
concordant with a 207Pb/206Pb age of 2724·3 ± 4·5 Ma,
and a regression line through all four zircon analyses
gives calculated upper and lower intercept ages of 2725·6
+7·4/–3·8 Ma and 1472 Ma, respectively. The crystallization age of the rock is given by the 207Pb/206Pb age
of fraction B, and the lower intercept indicates that
mainly Mesoproterozoic Pb loss has affected the zircons.
The two monazite fractions give considerably younger
ages, at 2704 and 2695 Ma. These results may indicate
either prolonged growth of metamorphic monazite in
this unit, or partial and variable Pb loss from igneous
monazite, or some combination of the two.
Sample U-54 (clinopyroxene–biotite granodiorite)
Zircons in this sample form relatively coarse, medium
brown, weakly zoned, stubby euhedral prisms. Four
fractions were analysed (Table 2). Three of these are
2·1–9·1% discordant, and define a linear discordia array
with a relatively imprecise upper intercept age of 2737·0
1626
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 7. Primitive-mantle-normalized (Sun & McDonough, 1989) extended-element profiles showing average compositions for Utsalik domain
plutonic rocks.
Fig. 8. Average rare-earth element plots (chondrite normalized; Sun & McDonough, 1989) for Utsalik domain plutonic rock types.
+10·5/–5·3 Ma (Fig. 11b). This regression is strongly
controlled by the most discordant analysis (DA). The
three most concordant analyses give 207Pb/206Pb ages in
the range of 2723–2729·5 Ma (Table 2). The calculated
upper intercept is interpreted as a maximum crystallization age for the sample, and the 207Pb/206Pb age of
the most concordant fraction (AB, 2729·5 Ma) gives a
minimum possible crystallization age.
1627
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
SEPTEMBER 2002
Fig. 9. Primitive-mantle-normalized extended-element profiles comparing hornblende-, biotite-bearing and pyroxene-bearing Utsalik granodiorites. Symbols correspond to rock types listed in Figure 6.
Lake Minto domain
Sample M-13 (clinopyroxene–hornblende–biotite
granodiorite)
intercept is interpreted as the crystallization age of the
rock unit, and the pattern of discordance indicates that
both Mesoproterozoic and more recent Pb loss affected
the zircons.
Zircons recovered from this sample form stubby euhedral
prisms that range from clear and colourless to translucent
and dark brown. Six single- and multi-grain fractions of
abraded zircons were analysed, including two brown
grains (AA and AB) and four fractions of relatively clear
zircon (Fig. 11c). A regression through the three most
concordant fractions (0·7–4·8% discordant; including two
clear fractions and one brown fraction) gives calculated
upper and lower intercept ages of 2724·3 +2·9/–2·0 Ma
and 1057 Ma, respectively. The more discordant fractions
plot both above and below this regression line. The upper
Sample M-14 (clinopyroxene–hornblende–biotite
granodiorite)
Zircons from this sample occur as stubby euhedral prisms
with faint growth zonation and no evidence of metamorphic rims. Five abraded single zircon grains define
a linear array with individual fractions ranging from 1·6
to 34·9% discordant (Fig. 11d). Calculated upper and
lower intercept ages are 2724·8 ± 3·3 Ma and 1483 Ma.
The upper intercept age gives the crystallization age of
the rock unit and the lower intercept indicates that mainly
1628
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 10. Geochemical comparison of Utsalik domain plutonic rocks and Lake Minto granodiorites. (a) Primitive-mantle-normalized extendedelement profiles comparing the field for Utsalik plutons (Fig. 7) with Lake Minto analyses from Stern et al. (1994). (b) Sr/Y vs Y plot for Lake
Minto and Utsalik domain plutons. For scaling reasons, two Lake Minto rocks (Sr/Y 1113 and 6010; Y 0·7 and 0·1) and one Utsalik rock (0·66,
98) were omitted from the diagram. Adakite and ‘typical arc’ fields from Castillo et al. (1999). The data show compositional gradation and
suggest the influence of a slab melt component in the generation of plutons in the Lake Minto domain.
Mesoproterozoic Pb loss has affected the zircons. Five
abraded monazite fractions were also analysed; the analyses cluster on or near concordia with 207Pb/206Pb ages
ranging from 2704 to 2712 Ma (Table 2). The reason
for the range of ages for monazite is not certain, and
may reflect either prolonged growth of metamorphic
monazite or partial resetting of igneous monazite, or
both.
1629
0·006
0·003
0·006
0·002
0·001
0·005
0·004
0·009
0·017
0·009
0·004
0·004
0·003
0·006
A: N2,+134,1
B: N2,+134,1
C: N2,+134,1
1630
D: N2,+134,1
E: N2,+134,1
M1,1
M2,1
M3,1
M4,1
M5,1
Sample U-54
AA: N1,105-149,1
AB: N1,105-149,4
CA: N1,105-149,1
DA: N1,105-149,1
103
303
216
341
65
180
134
207
17500
12050
10350
13260
15070
1024
551
526
94
436
157
113
1550
952
695
695
21530
26130
362
1278
6026
2043
753
5652
5652
10800
580
2054
6493
34760
2083
346
785
380
1011
850
2014
4222
13
41
41
24
38
23
495
108
47
62
45
35
24
15
46
12
139
17
15
319
33
41
21
17
17
33
18·6
10·8
13·3
12·8
95·7
91·2
94·7
96·1
96·7
2·0
3·6
3·0
3·1
2·1
6·6
10·5
7·6
14·4
84
9·5
13·9
11·3
13·1
12·8
98·9
98·6
0·50327(0·41)
0·51715(0·11)
0·52312(0·15)
0·51866(0·13)
0·52072(0·10)
0·52305(0·09)
0·44834(0·09)
0·52075(0·10)
0·52119(0·11)
0·44834(0·09)
0·44834(0·09)
0·43285(0·11)
0·43285(0·11)
0·49800(0·09)
0·31151(0·13)
0·50933(0·16)
0·489737(0·11)
0·52013(0·10)
0·52357(0·16)
0·23911(0·12)
0·51839(0·15)
0·52545(0·29)
0·52473(0·14)
0·51488(0·16)
0·52300(0·11)
0·51803(0·09)
(±% 1)
Pb†
208
Pb/238U§
206
%
Pb/235U§
12·827(0·41)
13·444(0·14)
13·577(0·18)
13·430(0·14)
13·345(0·11)
13·456(0·10)
13·502(0·26)
13·328(0·12)
13·362(0·12)
10·622(0·10)
11·632(0·15)
10·045(0·12)
13·466(0·21)
12·554(0·10)
7·907(0·19)
13·070(0·17)
12·109(0·12)
13·425(0·11)
13·559(0·17)
5·989(0·28)
13·322(0·16)
13·617(0·29)
13·600(0·15)
13·227(0·16)
13·386(0·12)
13·184(0·10)
(±% 1)
207
Pb/206Pb§
0·18485(0·05)
0·18854(0·06)
0·18823(0·09)
0·18780(0·06)
0·18588(0·03)
0·18659(0·03)
0·18614(0·19)
0·18562(0·05)
0·18593(0·03)
0·17182(0·04)
0·17804(0·07)
0·16830(0·03)
0·18721(0·04)
0·18721(0·04)
0·18410(0·12)
0·18612(0·04)
0·18020(0·03)
0·18720(0·03)
0·18782(0·04)
0·18167(0·22)
0·18639(0·07)
0·18795(0·14)
0·18797(0·06)
0·18632(0·07)
0·18562(0·04)
0·18458(0·04)
(±% 1)
207
Pb/206Pb
2696·9(1·7)
2729·5(2·1)
2726·8(2·9)
2723·0(1·9)
2706·0(0·9)
2712·3(0·9)
2708·3(6·4)
2703·7(1·7)
2706·5(1·0)
2575·5(1·3)
2634·7(2·2)
2540·8(1·0)
2717·8(1·4)
2678·7(0·9)
2690·2(4·0)
2708·2(1·4)
2654·7(1·0)
2717·7(1·0)
2723·2(1·3)
2682·0(7·3)
2710·6(2·4)
2724·3(4·5)
2724·4(2·0)
2709·9(2·3)
2703·8(1·5)
2694·5(1·2)
( Ma;±% 2)
age
207
NUMBER 9
1407
1965
1007
950
927
2222
1109
1173
170
843
461
195
433
412
270
213
164
92
387
270
9124
8287
Pb (pg)
common
Pb/204Pb Total
(meas.)‡
206
VOLUME 43
∗N1, N2, non-magnetic at given degrees side slope on Frantz isodynamic magnetic separator; grain size given in microns; u, unabraded; M, monazite; last digit
is number of grains analysed.
†Radiogenic Pb; corrected for blank, initial common Pb, and spike.
‡Corrected for spike and fractionation.
§Corrected for blank Pb and U, and common Pb, Stacey & Kramer (1975).
0·003
0·004
807
664
461
794
0·037
0·027
0·002
0·009
BB: N5,+134,1
BC: N5,+134,6
Sample M-14
266
151
627
447
185
212
content
(ppm)
content
Pb†
(ppm)
U
0·003
0·003
0·001
0·001
0·015
0·020
(mg)
description∗
Sample U-34
A: N2,+105,1
B: N2,+105,1
C: N2,+105,1
D: N2,+105,4
M1,u,1
M2,u,1
Sample M-13
AA: N5,+134,1
AB: N5,+134,1
AC: N5,+134,1
AD: N5,+134,3
Wt
Sample
Table 2: U–Pb analytical data for zircon and monazite
JOURNAL OF PETROLOGY
SEPTEMBER 2002
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 11. U–Pb concordia diagrams for zircon and monazite from four samples. POF, probability of fit of regression line.
MINERAL COMPOSITIONS AND
CRYSTALLIZATION CONDITIONS
Analyses of minerals of Lake Minto and Utsalik domain
plutons (Fig. 5) are reported in Tables 3–10 and derived
P–T conditions in Tables 11 and 12. Mineral compositions are similar in both domains, as are derived
crystallization conditions.
Clinopyroxene compositions straddle the diopside–
augite boundary, with Wo contents of 42–47 mol %
[Table 3; Lindsley (1983) parameters]. The mg-number
[Mg/(Mg + Fe2+)] is generally in the range 0·62–0·76;
Al2O3 constitutes <2 wt % and Na2O <0·55 wt %.
Exsolution lamellae (100) of orthopyroxene on a scale of
2–10 m in clinopyroxene cores (Fig. 3d–f ) are too
fine for microprobe resolution; the internal zones were
analysed with a defocused beam and yielded minimum
Wo contents of 40 mol % (reintegrated).
Orthopyroxene (Table 4) has generally low Wo contents (<6 mol %), lower mg-numbers than coexisting clinopyroxene (0·17–0·65), as well as lower Al2O3 contents
(<1·07 wt %).
Plagioclase compositions (Table 5) fall in the range
An12–56, with Or contents up to 3 mol %, although most
common rock types have plagioclase between An20 and
An30. Substantial amounts of normal zonation (An56–35)
are present within single grains in the more calcic rocks,
whereas variation is limited to a few mol % in the An20–30
range. Coarse antiperthite (exsolved Or blebs and strings
up to 10 by 20 m), particularly in the central parts
of grains, suggests slightly higher Or contents before
exsolution. BaO is present in amounts up to 0·3 wt %.
Alkali feldspar (Table 5) is mainly orthoclase, with
albite components as high as 43 mol %. Minor BaO
(<1·2 wt %) is present in most analyses, with SrO contents
below 0·1 wt %. Microcline occurs in both granites
and granodiorites. Cryptoperthitic alkali feldspars with a
ternary component occur in several pyroxene-bearing
rocks and three have been examined in detail with the
microprobe. The feldspars are internally heterogeneous
with strings and blebs of exsolved plagioclase on the 2–20
by 100–200 m scale, in proportions up to 30 vol. %.
The exsolved plagioclase compositions correspond closely
to those of matrix plagioclase, with maximum An content
of 29. Reintegration of the ternary feldspar compositions
yield An8·6Ab21·4Or70 to An2·5Ab22·5Or75 (Table 5).
Amphibole generally plots in the edenitic and ferroedenitic hornblende fields defined by Leake (1978). Its
mg-number is generally in the range 45–65 with a few
anomalous values (Table 6). Calculated Fe3+/(Fe3+ +
1631
0·28
MnO
99·89
Total
99·13
0·34
97·98
0·47
21·7
13·8
0·38
6·37
2·40
0·05
1·41
0·21
51·2
M-6
1632
0·89
0·02
0·04
0·31
0·01
0·67
0·91
0·03
0·72
Fe3+
Fe2+
Mn
Mg
Ca
Na
mg-no.
0·09
0·17
Fs
0·42
0·14
0·43
0·12
0·45
0·44
0·76
0·03
0·89
0·78
0·01
0·20
0·07
0·07
0·12
0·44
0·44
0·75
0·05
0·87
0·77
0·01
0·21
0·07
0·00
1·97
0·15
0·40
0·45
0·77
0·03
0·92
0·71
0·01
0·27
0·00
0·00
0·08
0·01
0·16
0·38
0·45
0·76
0·04
0·92
0·69
0·02
0·29
0·00
0·00
0·06
0·00
1·99
99·21
0·53
22·6
12·2
0·52
9·26
0·00
0·00
1·32
0·16
52·6
M-12
0·13
0·43
0·44
0·76
0·04
0·89
0·75
0·02
0·22
0·06
0·00
0·08
0·01
1·94
98·43
0·52
21·8
13·3
0·52
6·89
2·04
0·03
1·81
0·29
51·2
U-1
0·13
0·43
0·44
0·76
0·04
0·88
0·75
0·01
0·23
0·06
0·00
0·08
0·00
1·95
98·32
0·55
21·7
13·3
0·41
7·22
1·99
0·00
1·74
0·12
51·36
U-2
0·21
0·36
0·44
0·68
0·03
0·87
0·65
0·02
0·37
0·04
0·00
0·05
0·00
1·97
100·8
0·44
21·4
11·6
0·49
11·9
1·37
0·00
1·09
0·12
52·3
U-3
0·31
0·26
0·43
0·62
0·04
0·87
0·44
0·03
0·54
0·00
0·00
0·04
0·00
2·02
100·9
0·51
21·2
7·8
0·85
16·7
0·00
0·00
0·86
0·08
52·9
U-4
0·43
0·13
0·43
0·53
0·04
0·87
0·24
0·03
0·78
0·00
0·00
0·02
0·00
2·01
99·8
0·46
20·2
4·0
0·90
23·3
0·00
0·00
0·47
0·07
50·3
U-5
0·16
0·38
0·47
0·77
0·04
0·94
0·67
0·02
0·28
0·00
0·00
0·04
0·00
2·00
99·6
0·55
23·3
12·0
0·58
8·80
0·00
0·03
0·93
0·08
53·3
U-7
Weight per cent Fe2O3 calculated from charge-balanced structural formula; mg-number = Mg/(Mg + Fe2+ + Fe3+).
0·46
0·37
Wo
En
Lindsley (1983) end-members
0·06
0·00
1·95
0·01
99·44
0·43
23·0
12·7
0·29
8·55
0·00
0·10
1·86
0·22
52·4
M-8
0·16
0·42
0·43
0·77
0·03
0·89
0·70
0·01
0·27
0·00
0·00
0·08
0·01
1·99
101·0
0·43
22·6
12·7
0·44
8·66
0·00
0·06
1·74
0·20
54·1
U-8
0·24
0·32
0·44
0·67
0·03
0·89
0·57
0·01
0·43
0·00
0·00
0·05
0·00
2·00
98·8
0·41
21·6
9·9
0·44
13·4
0·00
0·00
0·99
0·15
51·8
U-11
0·19
0·39
0·42
0·72
0·03
0·86
0·69
0·01
0·34
0·00
0·00
0·06
0·01
2·00
97·5
0·43
20·9
12·0
0·37
10·4
0·00
0·00
1·31
0·16
51·9
U-12
0·15
0·40
0·45
0·77
0·04
0·92
0·72
0·01
0·26
0·00
0·00
0·07
0·00
1·98
98·9
0·52
22·7
12·7
0·38
8·35
0·11
0·05
1·54
0·14
52·4
U-13
NUMBER 9
0·79
0·74
0·01
0·24
0·00
1·95
0·01
98·91
0·67
21·5
13·7
0·18
6·79
2·53
0·02
1·61
0·23
51·7
M-7
VOLUME 43
0·01
0·06
0·00
Al
Cr
1·97
0·01
1·96
0·01
Si
Ti
Cations normalized for 6 oxygens
0·46
Na2O
CaO
0·21
7·71
22·3
9·96
FeO
0·00
22·4
1·32
Fe2O3
0·29
13·2
0·00
Cr2O3
2·03
0·36
52·7
M-5
11·9
1·25
Al2O3
MgO
0·20
52·0
TiO2
SiO2
M-1
Table 3: Electron microprobe analyses of clinopyroxene
JOURNAL OF PETROLOGY
SEPTEMBER 2002
0·00
0·00
Cr2O3
Fe2O3
97·3
0·03
0·68
17·6
0·74
26·1
0·00
0·00
0·62
0·07
51·4
M-1
99·8
0·00
0·49
22·5
0·46
21·8
0·00
0·12
1·07
0·11
53·2
M-5
1633
0·03
0·00
0·00
1·32
0·05
0·56
0·05
0·00
0·30
Fe3+
Fe2+
Mn
Mg
Ca
Na
mg-no.
0·03
0·68
Fs
0·01
0·45
0·54
1·98
0·35
0·64
0·01
0·65
0·00
0·02
1·26
0·01
0·68
0·00
0·00
0·05
0·00
0·37
0·61
0·01
0·62
0·00
0·02
1·20
0·13
0·73
0·01
0·00
0·03
0·00
1·97
99·1
0·00
0·55
21·3
0·82
23·1
0·21
0·00
0·77
0·13
52·2
M-6
0·32
0·63
0·05
0·64
0·00
0·09
1·22
0·02
0·61
0·09
0·00
0·04
0·00
1·94
100·3
0·00
2·15
21·9
0·64
19·7
3·02
0·04
0·80
0·08
52·0
M-7
0·42
0·57
0·01
0·57
0·00
0·02
1·13
0·03
0·82
0·01
0·00
0·03
0·00
1·98
100·9
0·00
0·44
20·1
0·83
25·7
0·42
0·04
0·74
0·06
52·5
M-8
0·45
0·54
0·00
0·54
0·00
0·01
1·01
0·02
0·85
0·00
0·00
0·17
0·00
1·93
101·5
0·00
0·17
18·2
0·49
27·2
0·00
0·02
3·74
0·13
51·7
M-10
0·45
0·53
0·02
0·54
0·00
0·03
1·04
0·04
0·87
0·00
0·00
0·03
0·00
1·99
99·6
0·02
0·72
18·1
1·12
27·1
0·00
0·06
0·69
0·09
51·7
M-12
0·50
0·49
0·01
0·49
0·00
0·03
0·94
0·04
0·97
0·00
0·00
0·03
0·00
1·99
100·4
0·02
0·68
16·3
1·34
29·9
0·00
0·03
0·57
0·04
51·6
U-3
0·61
0·34
0·05
0·35
0·01
0·10
0·64
0·06
1·15
0·05
0·00
0·02
0·01
1·97
102·1
0·08
2·37
10·9
1·89
34·9
1·71
0·00
0·34
0·16
49·8
U-4
0·81
0·17
0·02
0·17
0·00
0·04
0·32
0·07
1·54
0·00
0·00
0·01
0·00
2·01
99·9
0·02
0·82
5·13
1·99
43·9
0·00
0·01
0·13
0·09
47·8
U-5
0·78
0·20
0·02
0·20
0·01
0·03
0·38
0·03
1·51
0·00
0·00
0·04
0·00
1·99
101·4
0·07
0·67
6·28
0·89
44·0
0·00
0·06
0·86
0·08
48·5
U-6
Weight per cent Fe2O3 calculated from charge-balanced structural formula; mg-number = Mg/(Mg + Fe2+ + Fe3+).
0·03
0·29
Wo
En
Lindsley (1983) end-members
0·55
1·03
0·02
0·85
0·00
0·00
0·02
0·00
Al
Cr
2·01
0·00
1·99
0·00
Si
Ti
Cations normalized for 6 oxygens
99·6
0·04
Total
1·23
CaO
Na2O
1·40
9·11
MnO
MgO
38·7
0·46
Al2O3
FeO
0·10
48·6
TiO2
SiO2
T-1
Table 4: Electron microprobe analyses of orthopyroxene
0·36
0·62
0·01
0·63
0·00
0·02
1·19
0·03
0·69
0·00
0·00
0·04
0·00
2·00
100·9
0·05
0·60
21·6
0·91
22·4
0·00
0·00
0·91
0·06
54·4
U-8
0·72
0·27
0·01
0·27
0·00
0·03
0·52
0·02
1·39
0·00
0·00
0·02
0·00
2·00
99·1
0·03
0·59
8·46
0·56
40·3
0·00
0·00
0·46
0·10
48·6
U-10
0·54
0·40
0·06
0·42
0·00
0·11
0·78
0·03
1·05
0·02
0·00
0·02
0·00
1·98
100·8
0·01
2·66
13·3
0·99
32·1
0·71
0·07
0·49
0·11
50·3
U-11
0·44
0·54
0·01
0·54
0·00
0·03
1·06
0·03
0·86
0·03
0·00
0·03
0·00
1·97
97·2
0·01
0·64
18·0
0·99
26·1
0·91
0·01
0·61
0·11
50·0
U-12
0·36
0·60
0·04
0·61
0·00
0·07
1·16
0·02
0·69
0·06
0·00
0·04
0·00
1·95
101·4
0·00
1·87
20·9
0·73
22·3
2·03
0·01
0·98
0·14
52·4
U-13
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
623
491
449
516
468
448
456
456
698
647
TOr
394
467
492
620
616
650
443
274
457
1061
616
492
473
442
442
602
744
478
445
451
447
447
442
282
653
654
1613
595
745
747
492
275
438
407
860
TAn
387
466
444
280
357
VOLUME 43
TAb
451
77·16
70·00
63·13
94·45
91·73
96·02
95·05
55·78
XOr
Temperatures (°C) at 5 kbar from Elkins & Grove (1990)
4·80
80·00
87·16
91·85
75·00
90·55
91·18
89·71
95·05
95·71
15·20
12·84
0·00
0·00
8·15
22·50
2·50
0·00
9·45
4·29
0·00
8·82
0·00
4·95
0·00
0·00
10·29
22·09
0·75
8·60
21·40
36·67
0·20
0·00
5·55
8·25
0·02
0·83
XAn
3·87
4·95
0·00
43·39
XAb
0·10
2·19
1·03
1·92
1·90
1·33
2·03
1·16
1·71
XOr
Alkali feldspar
32·87
2·81
1·17
3·01
0·93
1·48
2·43
1·96
0·60
2·71
17·86
80·97
68·91
28·08
11·94
87·13
80·29
18·23
32·30
24·03
20·13
64·99
73·55
79·27
23·69
74·36
69·56
28·25
30·43
68·54
76·47
21·61
23·55
74·54
75·30
24·60
27·65
71·76
74·24
70·65
XAb
XAn
Plagioclase
23·36
U-9
U-7
U-5
U-4
U-1
G-2
G-1
M-13
M-11
M-10
M-9
M-3
M-2
T-3
T-2
T-1
Table 5: Coexisting feldspar compositions and derived temperatures
26·21
U-11
64·32
JOURNAL OF PETROLOGY
NUMBER 9
SEPTEMBER 2002
Fe2+) ratios vary from zero to 0·35. Al2O3 contents range
from 7·5 to 11·5 wt % and TiO2 from 0·57 to 2·3 wt %.
Halide contents are negligible.
Biotite has mg-numbers of 11–69 with most in the
range 40–65. TiO2 contents are generally 4–5 wt % with
exceptional values as high as 6·27 wt %. The micas
contain variable quantities of F (0·1–1·45 wt %) and
negligible Cl (0–0·29 wt %) (Table 7).
Ilmenite (Table 8) generally has low haematite contents
(0–4 mol %). All analyses indicate substantial MnO components (2–13·5 wt %). Coexisting magnetite (Table
9) contains up to 2·3 wt % TiO2 and minor Al2O3
(<1·9 wt %).
Garnet occurs in three samples, along with orthopyroxene, hornblende, biotite, plagioclase and quartz.
Compositions are in the range Alm61–75, Py5–29, Gro5–16,
Sp4–5 (Table 10).
A variety of minerals and assemblages provides information on pressures and temperatures of equilibration.
From these, crystallization conditions have been inferred.
Barometry
The aluminium-in-igneous hornblende barometer has
been calibrated empirically and experimentally, although
the thermodynamic basis is not well established. Many
calc-alkaline granodiorites of the Lake Minto and Utsalik
domains contain the requisite buffering assemblage (hornblende, biotite, titanite, magnetite, plagioclase, alkali feldspar, quartz), with the additional phases orthopyroxene
and clinopyroxene. Owing to their low Al contents, the
pyroxenes are unlikely to have influenced the Al budget
significantly. In several independent study areas, the Alin-hornblende barometer has provided results for pyroxene-bearing plutons consistent (within 1 kbar) with
pressures derived from their contact metamorphic
aureoles (e.g. Weiss & Troll, 1989; Lahti, 1995).
Results of the Al-in-hornblende barometer [Schmidt
(1992) calibration] are listed in Table 11 and plotted in
Fig. 12. Apparent pressures range from 3·8 to 6·5 kbar,
with most values between 4·5 and 5·5 kbar. Samples
located within <5 km of each other generally yield values
within 1 kbar. The compositional restrictions of Anderson
& Smith (1995) eliminate over half of the samples. Their
temperature correction results in a pressure range of
2·2–6·1 kbar and a lower average pressure (3·9 kbar).
No systematic variation in hornblende crystallization
pressure is evident either within or between the Lake
Minto (3·8–5·7 kbar; average 4·82 kbar) and Utsalik domains (3·9–6·5 kbar; average 4·95 kbar). The observation
of garnet–orthopyroxene–plagioclase–quartz pressures of
5·5–10·1 kbar in metamorphic enclaves of the Lake Minto
domain (Bégin & Pattison, 1994) is not supported by the
hornblende barometry for proximal samples, which yields
1634
0·00
0·00
Cr2O3
Fe2O3
1·67
0·58
0·09
K 2O
F
Cl
1635
0·21
13·5
2·56
0·03
9·21
1·74
43·2
M-1
95·6
0·16
0·13
1·27
1·25
11·4
95·5
0·07
0·27
1·08
1·36
11·5
9·75 10·8
0·48
13·7
5·24
0·00
8·89
1·13
42·2
T-2
1·31
42·3
M-3
96·4
0·06
0·50
1·21
1·50
11·7
12·0
0·50
11·4
4·14
0·04
97·3
0·13
0·46
1·45
1·33
11·7
11·0
0·30
12·0
5·26
0·02
8·77 10·0
1·08
43·6
M-2
0·36
0·49
0·04
0·38
6·62
0·55
0·02
0·13
0·21
0·41
1·89
2·46
0·03
1·73
0·30
0·00
1·66
0·20
6·62
0·59
0·02
0·24
0·23
0·44
1·90
2·71
0·06
1·44
0·47
0·01
1·57
0·12
6·41
0·54
0·03
0·22
0·28
0·39
1·91
2·49
0·04
1·52
0·60
0·00
1·79
0·15
0·52
0·01
0·00
0·24
0·30
1·89
2·41
0·08
1·56
0·67
0·00
1·56
0·17
6·56
95·0
0·04
0·01
1·19
1·00
11·4
10·4
0·59
12·1
5·79
0·00
8·58
1·46
42·4
M-4
0·68
0·05
0·30
0·29
0·34
1·95
2·83
0·02
1·36
0·00
0·12
1·73
0·26
6·61
96·4
0·19
0·63
1·52
1·18
12·1
12·6
0·15
10·8
0·00
1·02
9·80
2·32
44·0
M-5
2·09
43·3
M-8
0·66
0·01
0·15
0·21
0·36
1·88
3·02
0·02
1·07
0·49
0·01
1·60
0·19
6·61
96·3
0·03
0·32
1·12
1·23
11·7
13·6
0·17
8·6
4·40
0·04
0·61
0·02
0·30
0·26
0·47
1·94
2·69
0·01
1·63
0·10
0·03
1·84
0·24
6·48
97·8
0·06
0·64
1·37
1·60
12·1
12·1
0·10
13·0
0·88
0·24
9·09 10·4
1·65
44·3
M-7
0·69
0·01
0·14
0·16
0·32
1·81
3·27
0·03
0·72
0·76
0·00
1·32
0·06
6·84
96·0
0·02
0·29
0·87
1·12
11·5
14·9
0·20
5·8
6·83
0·01
7·58
0·57
46·4
M-9
0·39
13·1
1·76
0·00
9·23
1·86
42·7
0·44
0·03
0·14
0·29
0·46
1·86
2·01
0·05
2·01
0·56
0·01
1·76
0·25
6·36
98·6
0·10
0·29
1·49
1·54
11·4
0·58
0·02
0·33
0·29
0·46
1·87
2·57
0·05
1·69
0·20
0·00
1·68
0·22
6·59
95·2
0·09
0·67
1·47
1·54
11·3
8·88 11·2
0·38
15·8
4·86
0·04
9·82
2·18
41·8
0·62
43·5
0·55
0·02
0·18
0·21
0·38
1·91
2·55
0·09
1·59
0·49
0·00
1·48
0·09
6·72
97·2
0·08
0·37
1·11
1·29
11·8
11·3
0·67
12·6
4·29
0·00
0·28
9·8
6·05
0·05
9·26
1·85
43·3
U-1
0·45
0·00
0·13
0·23
0·39
1·89
2·01
0·06
1·92
0·52
0·00
1·89
0·07
6·53
98·7
0·00
0·27
1·22
1·33
11·7
0·59
0·01
0·07
0·25
0·41
1·80
2·70
0·04
1·23
0·68
0·01
1·64
0·21
6·50
96·7
0·03
0·14
1·30
1·40
11·2
8·98 12·1
0·49
15·3
4·58
0·01
8·34 10·7
0·81
44·5
M-11 M-12 M-13 G-2
0·64
0·03
0·05
0·25
0·32
1·89
2·86
0·03
1·19
0·44
0·00
1·75
0·20
6·54
96·6
0·12
0·10
1·32
1·10
11·8
12·9
0·23
9·5
3·88
0·00
9·98
1·77
43·9
U-2
0·52
0·02
0·27
0·29
0·43
1·81
2·29
0·04
1·88
0·27
0·00
1·67
0·26
6·60
99·4
0·07
0·57
1·52
1·49
11·4
10·3
0·28
15·1
2·42
0·03
9·55
2·31
44·4
U-3
0·30
0·02
0·09
0·24
0·51
1·80
1·36
0·06
2·79
0·32
0·00
1·59
0·22
6·66
98·0
0·07
0·18
1·20
1·70
10·8
5·89
0·49
21·5
2·77
0·00
8·68
1·87
42·9
U-4
0·21
0·03
0·28
0·27
0·58
1·80
0·95
0·06
3·32
0·30
0·00
1·50
0·19
6·68
99·3
0·10
0·56
1·34
1·91
10·7
4·04
0·47
25·3
2·56
0·00
8·11
1·64
42·6
U-5
Weight per cent Fe2O3 calculated from charge-balanced structural formula; mg-number = Mg/(Mg + Fe2+ + Fe3+).
no.
mg-
0·06
0·29
0·02
F
Cl
0·25
0·38
0·34
Na
K
2·25
1·90
1·54
0·06
1·93
0·04
Mn
1·77
0·61
Mg
2·73
Fe2+
Ca
0·00
Fe3+
0·00
1·64
1·81
0·00
Al
Cr
6·53
0·13
6·66
0·18
Si
Ti
Cations normalized for 23 oxygens
Total 96·1
1·23
11·4
Na2O
CaO
0·32
6·56
MnO
MgO
20·7
9·76
Al2O3
FeO
1·53
42·3
TiO2
SiO2
T-1
Table 6: Electron microprobe analyses of hornblende
U-7
8·27
0·21 0·55
0·00 0·03
0·22 0·04
0·33 0·24
0·44 0·40
1·79 1·90
0·94 2·48
0·02 0·06
2·92 1·87
0·53 0·13
0·01 0·00
2·06 1·49
0·17 0·15
6·37 6·82
100·3 95·9
0·01 0·11
0·45 0·09
1·67 1·25
1·49 1·34
10·9 11·6
4·0810·9
0·16 0·46
22·7 14·7
4·57 1·17
0·04 0·00
11·4
1·44 1·34
41·4 44·7
U-6
0·62
0·03
0·09
0·24
0·38
1·84
2·70
0·03
1·51
0·16
0·01
1·67
0·22
6·71
98·0
0·11
0·20
1·26
1·34
11·7
12·3
0·21
12·3
1·40
0·04
9·64
2·01
45·6
U-8
1·29
40·7
0·49
0·01
0·16
0·27
0·45
1·82
2·22
0·09
1·89
0·42
0·00
1·53
0·11
6·74
97·8
0·04
0·33
1·40
1·52
11·2
9·83
0·70
15·0
3·69
0·00
0·25
0·02
0·06
0·20
0·52
1·71
1·07
0·02
2·61
0·65
0·00
2·10
0·15
6·40
96·3
0·08
0·11
1·01
1·70
10·1
4·56
0·15
19·8
5·45
0·01
0·22
10·6
2·88
0·12
10·1
0·75
45·1
0·37
0·02
0·04
0·32
0·28
1·90
1·60
0·03
2·43
0·24
0·00
2·08
0·16
6·48
96·3
0·09
0·09
1·61
0·94
11·4
0·63
0·04
0·13
0·25
0·36
1·91
2·83
0·03
1·31
0·32
0·01
1·76
0·08
6·66
97·8
0·16
0·27
1·33
1·25
12·1
6·88 12·9
0·19
18·7
2·08
0·00
11·3
1·36
41·7
0·52
0·00
0·01
0·19
0·39
1·84
2·31
0·03
1·97
0·19
0·00
1·45
0·15
6·89
99·8
0·00
0·02
1·04
1·39
11·7
10·6
0·26
16·1
1·70
0·03
8·43
1·40
47·1
U-10 U-11 U-13 U-14
8·60 11·3
0·95
44·5
U-9
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
SiO2
0·04
1636
96·9
94·8
0·11
0·53
9·73
0·13
0·05
0·02
12·6
0·25
17·0
0·04
14·1
2·96
37·2
T-3
93·2
0·11
0·37
9·58
0·10
0·42
0·03
11·5
0·03
16·1
0·07
13·6
4·78
36·5
M-1
96·4
0·13
0·92
9·77
0·05
0·17
0·03
17·1
0·06
11·1
0·45
13·5
4·23
38·9
M-5
0·00
2·10
0·56
0·03
0·18
1·90
0·03
0·03
0·01
2·68
0·00
2·63
0·73
0·03
0·43
1·83
0·01
0·01
0·01
3·74
0·01
1·37
0·05
5·85
0·68
0·01
0·30
1·78
0·04
0·00
0·04
3·53
0·01
1·69
0·00
2·19
0·42
mg-number = Mg/[Mg + Fe(total)].
0·03
0·57
0·03
Cl
mg-no. 0·42
1·90
0·26
2·04
0·34
K
0·04
F
0·01
0·01
Ba
2·87
2·81
0·01
5·71
0·47
0·55
0·03
0·46
1·80
0·01
0·03
0·01
2·62
0·01
2·18
0·01
2·55
0·66
5·55
99·2
0·12
0·99
9·57
0·03
0·46
0·04
11·9
0·04
17·7
0·09
14·7
5·91
37·7
M-10
0·49
0·04
0·46
2·22
0·03
0·00
0·00
2·45
0·02
2·60
0·00
2·66
0·03
5·87
95·0
0·14
0·92
11·1
0·10
0·05
0·01
10·5
0·16
19·9
0·01
14·4
0·28
37·5
M-11
0·69
0·02
0·67
1·88
0·02
0·02
0·01
3·48
0·01
1·56
0·00
2·55
0·52
5·75
92·1
0·08
1·35
9·48
0·06
0·26
0·08
15·0
0·09
11·9
0·00
12·5
4·42
36·9
M-12
0·11
0·00
0·11
1·92
0·01
0·00
0·00
0·49
0·08
4·12
0·00
3·07
0·42
5·45
97·7
0·00
0·21
9·49
0·02
0·01
0·02
2·06
0·58
31·0
0·00
16·4
3·55
34·3
G-1
0·54
0·00
0·26
1·97
0·01
0·01
0·01
2·86
0·06
2·41
0·00
2·69
0·19
5·63
93·9
0·00
0·52
9·86
0·02
0·10
0·03
12·3
0·43
18·4
0·02
14·6
1·65
36·0
G-2
0·63
0·05
0·10
1·98
0·01
0·03
0·01
3·23
0·02
1·86
0·00
2·47
0·58
5·50
94·0
0·20
0·21
10·1
0·02
0·41
0·03
14·1
0·15
14·5
0·03
13·6
4·98
35·8
U-1
0·64
0·07
0·39
1·96
0·01
0·02
0·00
3·24
0·01
1·82
0·00
2·46
0·57
5·52
96·5
0·29
0·82
10·3
0·03
0·30
0·00
14·5
0·06
14·5
0·02
13·9
5·09
36·8
U-2
0·54
0·02
0·38
2·04
0·02
0·01
0·00
2·59
0·02
2·17
0·00
2·36
0·71
5·61
97·6
0·06
0·80
10·6
0·08
0·23
0·02
11·5
0·18
17·2
0·02
13·3
6·27
37·2
U-3
0·33
0·02
0·04
1·81
0·04
0·03
0·03
1·65
0·02
3·28
0·00
2·38
0·53
5·72
97·3
0·06
0·08
9·14
0·13
0·54
0·17
7·13
0·14
25·3
0·03
13·1
4·54
36·9
U-4
0·27
0·00
0·28
1·86
0·02
0·04
0·01
1·31
0·01
3·46
0·00
2·51
0·65
5·58
96·8
0·01
0·56
9·24
0·06
0·71
0·03
5·56
0·07
26·2
0·00
13·5
5·50
35·4
U-6
0·54
0·02
0·16
1·97
0·02
0·01
0·00
2·53
0·02
2·19
0·00
2·64
0·48
5·67
94·9
0·09
0·33
10·1
0·06
0·15
0·01
11·1
0·16
17·1
0·00
14·6
4·15
37·1
U-7
0·63
0·03
0·19
1·86
0·03
0·02
0·00
3·08
0·02
1·78
0·00
2·40
0·61
5·67
96·8
0·13
0·41
9·84
0·09
0·41
0·00
13·9
0·12
14·4
0·03
13·8
5·45
38·3
U-8
0·55
0·00
0·69
2·08
0·01
0·01
0·00
2·82
0·07
2·34
0·00
2·46
0·23
5·78
98·5
0·00
1·45
10·8
0·04
0·12
0·00
12·6
0·53
18·6
0·00
13·9
2·07
38·4
U-9
0·26
0·02
0·07
2·01
0·03
0·03
0·01
1·27
0·01
3·72
0·00
2·74
0·54
5·37
95·6
0·07
0·13
9·76
0·10
0·44
0·07
5·28
0·10
27·5
0·01
14·4
4·47
33·2
U-10
0·41
0·02
0·13
1·95
0·02
0·00
0·01
1·96
0·02
2·79
0·00
2·56
0·60
5·61
94·0
0·08
0·26
9·67
0·07
0·07
0·03
8·33
0·14
21·1
0·03
13·7
5·01
35·5
U-11
0·52
0·03
0·27
1·86
0·02
0·04
0·00
2·60
0·01
2·37
0·00
2·38
0·60
5·63
95·3
0·12
0·55
9·43
0·06
0·64
0·00
11·3
0·10
18·3
0·03
13·1
5·19
36·5
U-12
0·67
0·01
0·34
2·05
0·01
0·02
0·00
3·43
0·02
1·70
0·04
2·51
0·41
5·59
98·2
0·05
0·74
10·9
0·04
0·30
0·00
15·6
0·14
13·9
0·35
14·5
3·66
38·0
U-13
0·46
0·01
0·13
1·74
0·02
0·01
0·01
2·28
0·02
2·66
0·00
2·55
0·44
5·73
93·9
0·04
0·26
8·75
0·06
0·11
0·03
9·8
0·11
20·4
0·00
13·9
3·75
36·7
U-14
NUMBER 9
Na
0·00
2·15
0·01
Mg
Ca
2·18
Mn
0·03
2·91
0·01
Fe2+
0·01
2·85
2·46
0·00
Al
5·68
0·56
95·8
0·02
0·63
9·40
0·13
0·04
0·22
16·0
0·06
13·7
0·02
12·5
3·76
39·4
M-7
VOLUME 43
Cr
5·71
0·34
5·58
0·52
Si
Ti
Cations normalized for 22 oxygens
Total
0·69
0·12
F
Cl
10·3
Na2O
K 2O
0·04
0·15
CaO
BaO
0·05
9·27
MnO
22·4
0·00
13·4
4·47
36·0
MgO
FeO
Cr2O3
Al2O3
TiO2
T-1
Table 7: Electron microprobe analyses of biotite
JOURNAL OF PETROLOGY
SEPTEMBER 2002
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Table 8: Electron microprobe analyses of ilmenite
T-1
TiO2
50·1
M-11
50·3
U-3
51·4
U-4
U-8
U-9
U-11
U-12
U-13
U-14
51·8
49·8
52·3
49·3
48·5
47·7
51·1
Al2O3
0·05
0·05
0·04
0·04
0·09
0·06
0·04
0·08
0·07
Cr2O3
0·04
0·01
0·06
0·00
0·02
0·01
0·01
0·02
0·14
0·00
Fe2O3
4·48
4·43
2·67
0·00
1·36
0·24
3·56
4·20
10·79
1·56
FeO
42·9
41·8
42·3
41·4
39·6
33·2
43·0
40·3
38·4
MnO
2·01
3·36
3·81
3·47
4·95
13·52
1·30
3·16
4·28
MgO
0·01
0·00
0·04
0·00
0·00
0·00
0·05
0·04
0·02
Total
99·5
100·0
100·2
96·7
95·8
99·3
97·3
96·3
101·4
0·07
39·8
5·97
0·15
98·7
Cations normalized for 6 oxygens
Ti
1·91
1·92
1·95
2·02
1·97
1·99
1·93
1·91
1·79
1·96
Al
0·00
0·00
0·00
0·00
0·01
0·00
0·00
0·01
0·00
0·00
Cr
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·01
0·00
Fe3+
0·17
0·17
0·10
0·00
0·05
0·01
0·14
0·17
0·41
0·06
Fe2+
1·82
1·77
1·78
1·80
1·75
1·41
1·87
1·77
1·61
1·69
Mn
0·09
0·14
0·16
0·15
0·22
0·58
0·06
0·14
0·18
0·26
Mg
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
Ilm
0·01
91·2
88·6
89·1
92·2
87·5
70·6
93·5
88·6
80·4
85·0
Pyroph
4·4
7·2
8·1
7·8
11·1
29·1
2·9
7·0
9·1
12·9
Geikie
0·05
0·00
0·15
0·00
0·00
0·00
0·20
0·15
0·05
Hem
4·3
4·2
2·5
0·0
1·4
0·2
3·5
4·2
Eskol
0·10
0·00
0·10
0·00
0·05
0·00
0·00
0·05
10·2
0·30
0·55
1·5
0·00
Weight per cent Fe2O3 calculated from charge-balanced structural formula.
pressures up to 4·5 kbar lower (Fig. 12). However, pressures based on garnet–plagioclase–sillimanite–quartz barometry of pelitic enclaves correspond closely to the Alin-hornblende results. Bégin & Pattison (1994) suggested
that the garnet–orthopyroxene pressures of enclaves
could have been established before incorporation into
the calc-alkaline granodiorites. An alternative possibility
is that the hornblende compositions were not reset by
the >2700 Ma metamorphism, such that the hornblende
barometer records premetamorphic (i.e. 2725 Ma) crystallization pressures. The latter hypothesis is consistent
with 2700 Ma monazite ages of metasedimentary gneisses
(Percival & Skulski, 2000) and local, thin, 2700 Ma rims
on zircons from granodiorites (Percival et al., 2001).
Rare assemblages of garnet–orthopyroxene–plagioclase–quartz in granitoid rocks (U-6, U-10) yield independent pressure estimates (Table 11, Fig. 12). The
TWQ barometer (Berman, 1991; Berman, 1988 database) yields pressures lower (5·4, 5·5 kbar) than those
derived from hornblende from the same sample (6·9,
7 kbar) and slightly higher than Al-in-hornblende pressures of proximal samples (4·1, 4·3 kbar). The withinsample difference may reflect the sensitivity of the Al-inhornblende barometer to bulk compositions beyond the
mineralogically defined limits.
Thermometry
Analyses of 13 pyroxene pairs from Tables 3 and 4 were
used with QUILF software (Anderson et al., 1993) to
calculate temperatures. Results are reported in Table 12
and plotted in Fig. 12. They range from 625 to 1010°C,
with most between 700 and 800°C. With the exception
of the four values above 900°C, these temperatures do
not reflect primary pyroxene crystallization, but have
probably been reset during cooling or subsequent metamorphism.
Most of the feldspar compositions listed in Table 5
correspond to binary plagioclase and alkali feldspar solutions with very minor ternary components. These provide temperatures in the range 350–700°C using Elkins
& Grove’s (1990) calibration (Table 5). Most analyses
show reasonable internal consistency between albite, orthoclase and anorthite thermometers. Ternary compositions are preserved in alkali feldspar grains with
heterogeneous exsolution patterns. The fine scale of the
cryptoperthitic grain portions renders optical recognition
difficult and the ternary compositions were identified only
through microprobe analysis. Reintegrated compositions
showing An contents up to 8 mol % are plotted along
with coexisting plagioclase on 5 kbar isotherms from
1637
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
SEPTEMBER 2002
Table 9: Electron microprobe analyses of magnetite
T-1
M-11
U-3
U-4
U-8
U-9
U-11
U-12
U-13
U-14
TiO2
0·79
0·10
0·34
0·05
0·15
0·11
2·31
0·26
0·66
0·66
Al2O3
0·70
0·10
1·89
0·17
0·44
0·02
0·05
0·14
0·12
0·10
Cr2O3
0·04
0·02
0·16
0·03
0·01
0·06
0·16
0·00
0·84
0·06
Fe2O3
66·1
69·5
65·1
69·6
67·5
69·3
64·6
67·7
66·2
68·1
FeO
31·5
31·5
31·3
31·4
30·9
31·3
33·2
30·9
31·4
31·9
MnO
0·07
0·03
0·00
0·02
0·00
0·07
0·11
0·02
0·04
0·07
MgO
0·02
0·02
0·01
0·01
0·04
0·05
0·02
0·07
0·00
0·03
Total
99·1
101·2
98·8
101·3
99·0
100·9
100·4
99·1
99·3
100·9
Cations normalized for 4 oxygens
Ti
1·91
1·92
1·95
2·02
1·97
1·99
1·93
1·91
1·79
1·96
Al
0·00
0·00
0·00
0·00
0·01
0·00
0·00
0·01
0·00
0·00
Cr
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·01
0·00
Fe3+
0·17
0·17
0·10
0·00
0·05
0·01
0·14
0·17
0·41
0·06
1·69
Fe2+
1·82
1·77
1·78
1·80
1·75
1·41
1·87
1·77
1·61
Mn
0·09
0·14
0·16
0·15
0·22
0·58
0·06
0·14
0·18
0·26
Mg
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·00
0·01
Mgnsf
0·1
0·1
0·1
0·1
0·2
0·3
0·1
0·4
0·0
0·2
Mgnet
96·0
99·4
94·4
99·4
98·4
99·3
93·0
98·5
96·6
97·6
0·0
Spinl
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
Hercy
1·6
0·2
4·3
0·4
1·0
0·1
0·1
0·3
0·3
0·2
Mgchr
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
Chrom
0·1
0·1
0·3
0·1
0·0
0·1
0·3
0·0
1·3
0·1
Mgulv
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
0·0
Feulv
2·3
0·3
1·0
0·1
0·4
0·3
6·6
0·8
1·9
1·9
Weight per cent Fe2O3 calculated from charge-balanced structural formula.
Elkins & Grove (1990) (Fig. 13). The highest An contents
indicate temperatures in excess of 1100°C, at the upper
limit of possible igneous crystallization conditions. Diffusion of Na and K from feldspar is common during cooling
and exsolution (Fuhrman & Lindsley, 1988), and this
probably accounts for low apparent temperatures and
internal disequilibrium indicated by the wide temperature
spread (Table 5). A concentration of values in the 450°C
range suggests that this temperature may represent the
effective closure of feldspars to re-equilibration. Late K
loss from plagioclase is suggested by consistently lower
temperatures indicated by plagioclase compositions (Fig.
13).
Hornblende is a late magmatic phase based on its
habit and could indicate conditions near the end of the
crystallization history. Although the feldspar thermometry indicates that Na and K equilibrated down to
relatively low temperatures (<450°C), it is likely that
plagioclase established its Na/Ca ratios at higher temperatures, possibly in equilibrium with hornblende.
Results from calibrations (Blundy & Holland, 1990;
Holland & Blundy, 1994) of the equilibria
edenite + 4 quartz = tremolite + albite
(A)
edenite + albite = richterite + anorthite
(B)
and
are reported for 27 rocks in Table 11 at an assumed
pressure of 5 kbar. Values range from 575 to 820°C for
the three calibrations, with narrower ranges for each. In
the calibrations taking account of amphibole non-ideality
(Holland & Blundy, 1994), most values fall between 680
and 760°C, with identical mean temperatures of 714°C
for Utsalik and Lake Minto domain rocks. The common
occurrence of pyroxene and amphibole in this temperature range has been demonstrated in many experimental studies (e.g. Naney, 1983; Conrad et al., 1988;
Patiño Douce, 1995). Thus these values may reflect
the range in which hornblende precipitated or reached
equilibrium with plagioclase in a subsolidus state.
1638
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
may be suitable as a thermometer at known pressure.
Using mineral compositions from Tables 3, 5 and 6, and
amphibole parameters from Mäder et al. (1994) in TWQ
(Berman, 1991), highly scattered values and temperatures
lower by 100–300°C than the hornblende–plagioclase
values were calculated. Either the thermometer is not
well calibrated for this bulk composition, or hornblende
may have grown directly from residual melt and did not
equilibrate with previously crystallized clinopyroxene.
Garnet–orthopyroxene and garnet–biotite assemblages
are present in two samples (U-6, U-10), along with
hornblende (Table 12). Relatively low temperatures (685,
730°C) are indicated by Fe–Mg exchange thermometers,
whereas the more refractory Al-in-orthopyroxene
thermometer (Aranovich & Berman, 1997) gives values
in the 800°C range.
Fe–Mg exchange between biotite and orthopyroxene
has been calibrated as a thermometer. Using values from
Tables 4 and 7 in TWQ yields a wide range of values,
from >515 to 950°C. The values less than >700°C do
not represent equilibrium conditions because they are
below the orthopyroxene stability field. Temperatures in
the 950°C range approach biotite’s thermal stability
limit, although Ti-saturated biotites may be stable at
temperatures as high as 1000°C (Patiño Douce, 1993;
Patiño Douce & Beard, 1995). The high TiO2 contents
of biotites (up to 6·27 wt %; Table 7) qualitatively support
the high crystallization temperatures.
Table 10: Electron microprobe analyses of
garnet
SiO2
TiO2
Al2O3
M-10
U-6
U-10
36·6
37·6
37·5
0·03
21·6
0·02
20·3
Cr2O3
0·07
0·00
Fe2O3
3·33
0·00
FeO
26·7
32·7
0·07
20·4
0·00
0·02
33·0
MnO
2·12
1·88
1·79
MgO
7·19
1·16
1·57
CaO
1·64
5·49
Total
99·2
99·0
5·73
100·1
Cations normalized for 24 oxygens
Si
5·78
6·11
6·05
Ti
0·00
0·00
0·01
Al
4·02
3·89
3·88
Cr
0·01
0·00
0·00
Fe3+
0·40
4·44
4·45
Fe2+
3·53
0·00
0·00
Mn
0·28
0·28
0·38
Mg
1·70
0·26
0·25
Ca
0·28
0·96
0·99
mg-no.
0·30
0·06
0·05
Pyrope
27·6
4·7
6·2
Almand
57·6
74·8
73·4
Spessa
4·6
4·4
4·0
Uvarov
0·2
0·0
0·0
Grossu
0·5
16·1
16·3
Andrad
9·5
0·0
0·1
Water fugacity
Weight per cent Fe2O3 calculated from charge-balanced structural formula; mg-number = Mg/(Mg + Fe2+ + Fe3+).
The spatial distribution of temperature values from
equilibrium (B) is illustrated in Fig. 12. The Lake Minto
and Utsalik domains exhibit similar ranges (670–790°C)
and identical mean temperatures (718°C), with a slightly
higher standard deviation in the Utsalik domain. There
is no indication of higher values in the western Lake Minto
domain, where high garnet–orthopyroxene temperatures
are recorded in paragneiss enclaves (Bégin & Pattison,
1994).
If hornblende–plagioclase assemblages reached equilibrium with igneous clinopyroxene and quartz, then the
equilibrium
2 diopside + 2 quartz + tschermakite = tremolite +
2 anorthite
Minerals in several assemblages provide estimates of
water fugacity, assuming equilibrium. The assemblage
biotite–quartz–orthopyroxene–K-feldspar is a common
hygrometer in igneous and metamorphic rocks. It returns
TWQ temperatures (5 kbar) in the 760–840°C range at
unit water activity (Table 12). Assemblages of orthopyroxene–clinopyroxene–plagioclase–hornblende–quartz
yield TWQ values of aH2O, at 5 kbar and temperatures
defined by the hornblende–plagioclase thermometer, in
the range 0·06–0·18. Using Burnham’s (1979) equations,
these values correspond to magmatic water contents in
the range 2–3 wt %. The significance of these calculations
is unclear in light of the wide range of crystallization
temperatures inferred for pyroxene, biotite and amphibole; however, the results appear consistent with experimentally derived phase equilibria (see below).
Oxygen fugacity
Estimates of oxygen fugacity based on the assemblage
orthopyroxene–clinopyroxene–ilmenite–magnetite–
quartz were made with QUILF (Anderson et al., 1993).
Values range from FMQ +0·6 to +2, consistent with
1639
T-1
T-2
M-1
5·6
4·8
4·9
661
610
746
4
5
6
−22·2
801
575
740
748
760
815
701
744
738
710
769
744
5·5
M-3
734
749
752
4·4
M-4
705
773
821
5·2
M-5
717
737
738
4·6
M-7
−23·3
699
719
791
638
5·7
M-8
671
684
711
M-9
−26·3
578
782
817
777
5·4
724
735
742
4·8
678
719
715
3·9
653
693
714
4·1
M-11 M-12 M-13 G-2
791
622
765
4·8
U-1
714
728
775
5·3
U-2
−26·3
670
773
708
766
4·9
U-3
698
704
707
4·9
U-4
672
711
686
4·1
U-5
750
722
604
5·4
U-6
666
683
702
4·1
U-7
495
720
677
705
4·1
U-9
−36·6 −47·5
457
748
714
769
5·0
U-8
759
712
763
5·5
804
700
716
757
5·0
679
694
725
695
3·9
−21·7−17·9−28·6
817
646
654
743
6·5
U-10 U-11 U-13 U-14
1640
T-1
723
1010
870
703
M-8
760
895
M-10
5.0
796
630
M-12
670
770
U-3
938
840
920
U-4
625
U-5
5.5
790
725
555
U-6
744
920
U-8
5.6
650
680
705
920
U-10
972
765
895
U-11
950
U-12
959
820
U-13
References: 1, orthopyroxene–biotite Fe–Mg exchange thermometer (TWQ ; Berman, 1991); 2, biotite–quartz = orthopyroxene + K-feldspar + H2O equilibrium
(TWQ ; Berman, 1991, pers. comm., 1996); 3, two-pyroxene QUILF thermometer (Anderson et al., 1993); 4, almandine+enstatite = pyrope + ferrosilite
equilibrium (TWQ ; Aranovich & Berman, 1997); 5, pyrope = enstatite + Al in orthopyroxene (TWQ ; Aranovich & Berman, 1997); 6, anorthite+enstatite =
grossular+ferrosilite+quartz equilibrium (TWQ ; Aranovich & Berman, 1997).
6
Pressure (kbar)
805
759
760
M-7
5
678
885
M-6
700
515
M-5
4
3
2
M-1
NUMBER 9
1
Temperature (°C)
Ref.
VOLUME 43
Table 12: Equilibration conditions based on pyroxene and garnet compositions
References: 1, hornblende barometer (Schmidt, 1992); 2, garnet–orthopyroxene–plagioclase–quartz barometer ( TWQ ; Berman, 1991); 3, edenite–
quartz–tremolite–albite thermometer (Blundy & Holland, 1990); 4, edenite–quartz–tremolite–albite thermometer (Holland & Blundy, 1994); 5, edenite–
albite–richterite–anorthite thermometer (Holland & Blundy, 1994); 6, QUILF oxide thermometer (Anderson et al., 1993); 7, fugacity O2 from QUILF (Anderson et
al., 1993).
7
Fugacity O2
711
3
3·8
M-2
Temperature estimates (°C)
2
1
Pressure estimates (kbar)
Ref.
Table 11: Equilibration conditions based on hornblende and oxide compositions
JOURNAL OF PETROLOGY
SEPTEMBER 2002
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 12. Map showing pressure–temperature estimates. Most pressures are derived from Schmidt’s (1992) calibration of the Al-in-igneous
hornblende barometer and most temperatures from hornblende–plagioclase assemblages [Holland & Blundy (1994) calibration]. Also plotted
are pressures from TWQ for garnet-, orthopyroxene- and garnet-, sillimanite-bearing rocks, as well as two-pyroxene temperatures.
414–815°C (Table 11), reflecting post-crystallization reequilibration.
The presence of the assemblage clinopyroxene–
ilmenite in the more calcic compositions of the suite can
be used to derive relatively reducing conditions relatively
early in the crystallization history (Czamanske & Wones,
1973). Occurrences of amphibole–titanite–magnetite, developed in relatively potassic compositions, indicate that
oxygen fugacity may have increased with hydration during advanced crystallization (Wones, 1989; Frost et al.,
2000a).
Summary
Fig. 13. Ternary and plagioclase feldspar compositions plotted on
5 kbar isotherms (Elkins & Grove, 1990).
an arc environment (Anderson & Lindsley, 1988; Frost
& Lindsley, 1992). Magnetite–ilmenite pairs (Tables 8
and 9) yield QUILF temperatures in the range
Although reliable pressure-sensitive assemblages are
scarce in the calc-alkaline rocks, available geobarometers
yield relatively consistent pressure values of 5 ± 1 kbar.
Results of the Al-in-hornblende barometer are supported
by garnet–orthopyroxene–plagioclase–quartz data for igneous rocks and garnet–sillimanite–plagioclase–quartz
equilibria in enclosed metamorphic rocks.
Conversely, a wide range of temperature conditions is
indicated by various mineral geothermometers. Feldspars
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JOURNAL OF PETROLOGY
VOLUME 43
record an >700°C range, from >1100°C for cryptoperthitic ternary compositions to the 400°C range for
discrete pairs. The high temperatures for crystallization
of pyroxene-bearing assemblages are supported by several
high two-pyroxene and biotite–orthopyroxene temperatures. The range is narrower (>400°C) for twopyroxene assemblages (1010–625°C), and oxide pairs
(815–415°C). The wide observed temperature range
probably reflects variable resetting from crystallization
conditions. The variable degree of retrogression could
have been produced in a cooling magmatic system at
depth, through evolution of fluids during advanced crystallization. Even relatively anhydrous, high-level (3 kbar)
intrusions have been shown to be susceptible to diffusive
retrograde re-equilibration (e.g. Eggins & Hensen, 1987;
Weiss & Troll, 1989). Alternatively, later metamorphism,
crustal magmatism and hydrothermal activity
(2700–2630 Ma) could also have contributed to resetting.
CRYSTALLIZATION HISTORY AND
CONSTRAINTS ON WATER
CONTENTS
The crystallization history of the Leaf River suite can
be interpreted from mineral assemblages, reaction and
exsolution textures, temperatures derived from various
thermometers, and bulk compositional changes. Diorites
represent the most primitive compositions of the suite
(Stern et al., 1994) and may be parental; comagmatic
gabbro and pyroxenite enclaves may be cumulates. All
of these rock types, as well as quartz diorites and some
granodiorites, have assemblages of ortho- and clinopyroxene, plagioclase, biotite, hornblende and Fe–Ti
oxides, with or without quartz. Normally zoned plagioclase, and clinopyroxene evolving from subcalcic to diopside compositions in this compositional range reflect
declining temperature. Alkali feldspar is present sporadically in the highest-temperature assemblages, apparently following pyroxenes, biotite and plagioclase in
the crystallization order. Subsequent cooling led to exsolution in pyroxenes (001 pigeonite; 100 orthopyroxene)
and feldspar (cryptoperthite).
In the granodiorite to granite compositional range,
hornblende appears as an additional phase and alkali
feldspar is abundant. Based on textural observations of
amphibole as initial overgrowths on pyroxene, and in
more felsic compositions as the main ferromagnesian
mineral, it is possible that hornblende and alkali feldspar
are products of the subsolidus decomposition of clinopyroxene–biotite–quartz. However, based on TWEQU
calculations using natural compositions, the equilibrium
diopside + biotite + quartz = hornblende +
K-feldspar
NUMBER 9
SEPTEMBER 2002
does not occur in the 600–800°C temperature range. It
is more likely that hornblende crystallized directly from
the magma, in response to changes in bulk composition,
including increasing water content.
In summary, orthopyroxene, clinopyroxene, plagioclase, biotite and some alkali feldspar formed early in
the crystallization sequence. Fractionation of water into
the residual magma promoted some resorption of pyroxene and crystallization of hornblende in more evolved
compositions.
The crystallization order of early biotite and late hornblende differs from that of hydrous I-type granitoid rocks
(Eggler, 1972; Burnham, 1979; Wones & Gilbert, 1982;
Sisson & Grove, 1993). However, it matches closely that
of H2O-undersaturated, K-bearing magmas (Wones &
Gilbert, 1982; Naney, 1983; Weiss & Troll, 1989; Kilpatrick & Ellis, 1992). Experimental results for synthetic
granite and granodiorite (Naney, 1983) demonstrate qualitative crystallization orders at 2 and 8 kbar for a variety
of water contents. In only the driest magmas (<2 wt %
H2O) does biotite precede alkali feldspar in the crystallization order. Hornblende appears in the crystallization sequence only in granodioritic bulk
compositions containing >4 wt % H2O (Sisson & Grove,
1993; Beard, 1995). Additional indirect evidence for low
magmatic water contents in the calc-alkaline pyroxenebearing rocks includes low intrinsic water contents
(<1 wt % H2O; Table 1), lack of hydrothermal sericitization of plagioclase and lack of pyroxene alteration
products (Figs 3 and 4).
The crystallization history inferred from assemblages,
textures and thermometry is compared with a schematic
isobaric (5 kbar) T–XH2O diagram derived from Naney’s
(1983) experimental results in Fig. 14. Qualitative relations have been preserved in the nonrigorous construction of the phase diagram, with a few exceptions. A
modification was necessary to allow clinopyroxene and
hornblende to coexist in the same bulk compositional
space: Naney’s synthetic compositions did not produce
clinopyroxene in granodiorite or hornblende in granite.
The phase diagram qualitatively predicts the observed
assemblages and crystallization order; however, several
discrepancies are also apparent. The common hornblende–perthite (alkali feldspar) assemblage in granite is
not stable and must form in bulk compositions not
represented experimentally. Biotite begins to crystallize
in the synthetic compositions at <900°C, in contrast to
indications of higher temperatures in the natural biotites
which were probably stabilized as a result of high TiO2
contents (Patiño Douce, 1993). Full occupancy of the
M2 site by Ti and Al in high-temperature biotite (Table
7) may account for the abundance of magnetite in these
rocks, as all Fe3+ at given fH2O conditions would be
sequestered in magnetite (Pilkington & Percival, 2001).
1642
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
Fig. 14. Schematic T–water content diagram showing stability fields of
minerals (abbreviations as in Table 1) in calc-alkaline felsic compositions
[modified after Naney’s (1983) experimental results on synthetic granodiorite by linear interpolation between 2 and 8 kbar data to produce
an approximate 5 kbar isobaric section]. Epidote is not present in the
rocks under study and was omitted from consideration. Plotted on the
diagram are observed equilibrium assemblages (Β) from a range of
bulk compositions (Table 1; Fig. 5). Assemblages of hornblende–alkali
feldspar, common in Leaf River granitic compositions, are not stable
in the experimental range investigated by Naney. At the left margin is
the range of temperature estimates from various thermometers.
The maximum stability of alkali feldspar in the experiments is >1000°C, compared with higher values
recorded in natural ternary compositions. Hornblende is
stable at higher temperatures than those recorded by
hornblende–plagioclase thermometers. This discrepancy
may be explained if the water content of the natural
magmas only rose to values above >3 wt % below
temperatures of >800°C. The general T–H2O crystallization path is indicated by the array of assemblages
in Fig. 14.
Evidence supporting a primary igneous origin for pyroxene in calc-alkaline rocks of the Lake Minto and Utsalik
domains includes field-based, petrographic, thermometric
and geochemical observations. Field constraints include
the occurrence of large bodies of homogeneous, massive,
coarse-grained rocks with dehydrating contact relationships to enclaves and country rocks (Fig. 2d and
e). All bulk compositions, from pyroxenite to granite,
contain both hydrous and anhydrous mafic minerals;
however, the more felsic rocks generally contain more
hornblende and biotite. Where gradations between units
with pyroxene- and hornblende-dominant units were
observed, pyroxenes are invariably rimmed by hornblende (Figs 2b and c, and 4).
Mineralogical thermometers record a wide range of
temperatures from >1100 to 400°C with considerable
scatter both within and between samples. The highest
temperatures, sparsely recorded by ternary feldspar and
pyroxene compositions, approach the limits of crustal
metamorphic conditions. Had the region been metamorphosed to such extreme conditions, one would expect
to see evidence of tectonic fabrics, migmatization,
anhydrous mineralogy independent of bulk composition,
and better preservation of high-temperature features, the
effect of dehydration being to isolate the metamorphic
rocks from potentially rehydrating fluids. Instead, the
geochemically more evolved rocks exhibit relatively
hydrous assemblages, and the extreme spread in temperature results suggests re-equilibration in the presence
of an evolving fluid.
Geochemical features of the Utsalik domain rocks
support an igneous origin. Typical igneous calc-alkaline
fractionation trends (Figs 6–10) characterize the mineralogical change from pyroxene- to hornblende-, biotitebearing units. REE and LILE characteristics (Figs 7–9)
are consistent with igneous fractionation rather than
metamorphic depletion. Identical U–Pb zircon crystallization ages are recorded by hornblende (2724,
2725 Ma) and pyroxene-bearing granodiorites (2729,
2724 Ma). Zircons in pyroxene-bearing rocks rarely have
metamorphic rims (Percival et al., 2001).
PETROGENESIS OF IGNEOUS
PYROXENE-BEARING GRANITOID
ROCKS
In comparison with geochemically similar calc-alkaline
batholiths of younger continental margins, the Minto
rocks are notably less hydrous, although a continuum of
water undersaturation exists, as in calc-alkaline magmas
in general (Eggler, 1972; Hildreth, 1981; Clemens, 1984;
Beard, 1995). Models of subduction-zone magmatism
hold that slab-derived fluids initiate melting through
mantle metasomatism and are retained through the various stages of magma genesis (e.g. Wyllie et al., 1976;
Burnham, 1979; Gill, 1981; Merzbacher & Eggler, 1984;
Sisson & Grove, 1993; Sobolev & Chaussidon, 1996;
Peacock & Hyndman, 1999). However, granitoid magmas
are commonly water undersaturated (Eggler, 1972; Clemens, 1984; Merzbacher & Eggler, 1984) and pyroxenes
are common liquidus phases of calc-alkaline magmas
(Frost & Lindsley, 1992), as indicated by phenocrysts in
dacites of continental margin settings (e.g. Ewart, 1979).
Water contents of these explosive, shallow-level
1643
JOURNAL OF PETROLOGY
VOLUME 43
magmas are of the order of 1–6 wt % (Merzbacher &
Eggler, 1984; Beard, 1995; Kawamoto, 1996).
Arc plutonic rocks are presumed to have evolved
through more hydrous magmatic conditions in which
pyroxenes were resorbed and amphibole–biotite crystallized (Sisson & Grove, 1993). However, magmatic
systems are commonly zoned with respect to water content (Hildreth, 1981) and some plutons have waterundersaturated root zones [e.g. orthopyroxene-bearing
tonalites of the southern Sierra Nevada batholith (Ross,
1985, 1989; Sams & Saleeby, 1988; Barth & May, 1992)].
In the central Andes, Lucassen & Franz (1996) described a
magmatic history involving H2O-undersaturated igneous
crystallization to form orthopyroxene–clinopyroxene–
biotite ‘granulites’, followed by magmatic autometamorphism yielding amphibole-bearing ‘amphibolitefacies’ assemblages.
I-type felsic volcanic rocks commonly contain pyroxenes and may be linked to water-undersaturated plutonic bodies. For example, the charnockitic, I-type
Ballachulish igneous complex has rhyodacitic equivalents
of the most fractionated granitic compositions (Weiss &
Troll, 1989; Troll & Weiss, 1991). M-type (mantlederived) magmas may also fractionate to anhydrous granitoid compositions. For example, the Barrington Tops
granodioritic batholith of eastern Australia, emplaced at
>3 kbar levels, has a high-temperature crystallization
history involving ortho- and clinopyroxene (Eggins &
Hensen, 1987).
It is probable that water undersaturation is the norm for
granitoid magmas of various compositions (e.g. Clemens,
1984). The most compelling evidence for low water
contents throughout the magma solidification process is
crystallization orders without early amphibole (Wones &
Gilbert, 1982), implying that water undersaturation is an
intrinsic characteristic of the magma.
Models for production of virtually anhydrous (charnockitic) magmas include high-temperature (>1000°C)
melting of anhydrous crustal source rocks (Kilpatrick &
Ellis, 1992; Langdenberger & Collins, 1996; Young et
al., 1997; Zhao et al., 1997), producing distinctive ‘Ctype’ geochemical features. A second model involves
fractionation of primitive, mantle-derived, water-undersaturated magmas followed by high-temperature crystallization (e.g. Eggins & Hensen, 1987). A third model
considers the role of a CO2 component in felsic magmas
(e.g. Wendlandt, 1981; Frost et al., 1989; Peterson &
Newton, 1990), possibly derived from mixing of crustal
melts and CO2-bearing mafic magmas (Frost et al., 2000b).
Intermediate models involve mixing of variably hydrous
mantle-derived basalt and crustal components (MASH:
melting, assimilation, storage, homogenization; Hildreth
& Moorbath, 1988; Emslie & Hunt, 1990; Barker et al.,
1992; Stern et al., 1994; Patiño Douce, 1995) and probably
NUMBER 9
SEPTEMBER 2002
apply to the genesis of most hydrous and charnockitic
granitoid rocks.
Whereas pyroxene-bearing granitic plutons are not
uncommon in Phanerozoic terranes, batholiths of charnockitic rock are dominantly a Precambrian feature. In
particular, large pyroxene-bearing plutons of Neoarchaean (>2·7 Ga) and Palaeoproterozic (>1·85 Ga)
age are integral parts of orogens formed during major
episodes of crustal growth (see Condie, 1998). These
broad continental magmatic arcs such as the 2·725 Ga
Lake Minto and Utsalik domains of the Minto block,
2·63 Ga Louis Lake batholith of the Wyoming craton
(Frost et al., 2000b), 1·86–1·85 Ga Wathaman (Meyer et
al., 1992), Cumberland (St-Onge et al., 1998) and de Pas
(Dunphy & Skulski, 1996) batholiths of the Trans-Hudson
and New Quebec orogens, and 1·85 Ga central Finnish
batholith (Lahti, 1995), may mark a thermal regime in
Earth’s history conducive to the production of voluminous, water-undersaturated, arc magmas. Potential
mantle temperatures at 2·7 Ga may have been higher
than those at present by >100°C, based on theoretical
(Davies, 1998) and metamorphic (Galer & Mezger, 1998)
constraints. Given a constant flux of fluid emanating
from subducting lithosphere (see Peacock et al., 1994;
Schmidt & Poli, 1998) and comparable subduction rates
(Davies & Bickle, 1991; Kincaid & Sacks, 1997; Davies,
1999), the higher mantle temperatures (Davies, 1995)
would have evoked higher degrees of mantle wedge fusion
(Kawamoto, 1996; Gaetani & Grove, 1998), yielding
voluminous, relatively water-deficient magma. In the
Minto block example, the arc magmas rose into older
(2·77–3·0 Ga) continental crust, where they assimilated
small proportions (10–20%) of older material, fractionated, and rose to upper-crustal (>15 km) levels, in a
manner similar to that envisaged by Johnston & Wyllie
(1988), Ague & Brimhall (1988) and Patiño Douce
(1995). Early crystallization of pyroxene–biotite assemblages led to fractionation of more hydrous granitic
compositions, some of which may have risen to higher
crustal levels (see Holland & Lambert, 1975; Hubbard
& Whitley, 1979; Eggins & Hensen, 1987; Frost et al.,
1989).
The thermal regime preceding 2·7 Ga, when mantle
potential temperatures were still higher, produced mainly
tonalite–trondhjemite–granodiorite compositions, which
may have involved slab melting (e.g. Drummond &
Defant, 1990; Martin, 1994; Peacock et al., 1994; de Wit,
1998; Smithies & Champion, 2000), rather than fluiddriven suprasubduction zone processes. The scale, character and duration of calc-alkaline magmatism in the
Minto block suggest that arc processes resembling those
at modern long-standing consuming margins operated
effectively by 2725 Ma.
1644
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
SUMMARY AND CONCLUSIONS
Plutonic rocks of the Archaean Minto block formed
during a period of continental arc magmatism
(2775–2690 Ma). The pyroxene-bearing granitoid rocks
include a calc-alkaline quartz diorite to granite suite
(2725 Ma), as well as younger (2696–2688 Ma) crustderived granodiorite and granite suites. Water-undersaturated magmatic conditions are indicated by unaltered
pyroxene and feldspar in granitic rocks, and the presence
of orthopyroxene-rich reaction selvages where granites
are in contact with hornblende-bearing mafic bodies.
In the dominantly calc-alkaline suite, textures and
mineral compositions define a crystallization history
through bulk compositions ranging from pyroxenite to
granite. In the quartz diorite–granodiorite range, assemblages of ortho- and clinopyroxene, biotite, Fe–Ti
oxides, plagioclase, alkali feldspar and quartz predominate, whereas in transitional granodiorites and granites, hornblende occurs as a late magmatic phase. The
2729–2724 Ma pyroxene- and hornblende-bearing units
form a geochemical continuum (48–71 wt % SiO2) with
a typical continental arc signature, including LREE enrichment and negative Nb and Ti anomalies. The crystallization history is tracked by mineralogical
thermometers, which record early alkali feldspar growth
at >1100°C, pyroxene ± biotite crystallization at
>1000–800°C, and later hornblende growth and equilibration with plagioclase at 800–600°C, 5 ± 1 kbar.
Subsequent metamorphism is recorded in resetting of
feldspar and oxide temperatures into the 400°C range.
Field, geochemical and petrological evidence supports a
model of 5 kbar crystallization of hot (>1100°C), waterundersaturated I-type magma that evolved to hydrous
compositions. Bulk magma compositions produced in a
continental arc setting were water deficient owing to the
elevated temperature of the convecting mantle wedge at
2725 Ma. Abundant igneous charnockite in 2·7 and 1·85
Ga orogens suggests that such thermal conditions may
have existed in mantle wedges over >900 Myr of geological history.
ACKNOWLEDGEMENTS
Gratitude is due to the dedicated members of the Minto
Transect field team, particularly K. D. Card, R. A. Stern,
N. J. Bégin, N. Aleksejev, G. T. Shore, A. Ross and S.
H. Schwarz, whose research flourished through trying
field conditions. Discussions with R. A. Stern and T.
Skulski on petrology, geochemistry and geochronology
have been enlightening. Olga Ijewliw and John Stirling
provided reliable microprobe analyses and data reduction. Early versions of the manuscript benefited from
critical comments by Jean Bedard, Ken Currie and
particularly Tom Skulski. Journal reviewers Ron Frost
and John Tarney provided additional helpful comments
and insight. This paper is Geological Survey of Canada
contribution no. 2002007.
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1648
PERCIVAL AND MORTENSEN
WATER-DEFICIENT CALC-ALKALINE MAGMAS
APPENDIX: DETAILS OF ROCKS ANALYSED
Sample
Archival
Mafic
Rock
number
number
mineral
type
UTMX
UTMY
Grid
zone
assemblage
Tikkerutuk domain
T-1
PBAC89-77
OHB
gdi
454250
6303950
18
T-2
PBAS89-11
HB
gdi
486884
6327366
18
T-3
PBA89-111
B
gdi
495971
6334374
18
M-1
PBA89-28c
OCHB
gdi
492297
6336263
18
M-2
PBAS89-76
CHB
gdi
488900
6345800
18
M-3
PBAS89-53
CHB
gdi
514788
6370795
18
M-4
PBAS89-31
OCHB
gdi
477818
6375211
18
M-5
PBA89-137
OCHB
gdi
487652
6388424
18
M-6
PBAS89-47
OCB
gdi
522738
6356086
18
M-7
PBAS89-45
OCHB
gdi
513094
6349235
18
M-8
PBA89-144
OCHB
gdi
517698
6409039
18
M-9
PBAS89-81
CH
gdi
556676
6376769
18
M-10
PBA90-28
OBG
gdi
583671
6391210
18
M-11
PBA90-50
HB
gdi
579695
6409362
18
M-12
PBA89-90
OCHB
gdi
570629
6437919
18
M-13
PBA90-137a
CHB
gdi
479220
6378922
18
M-14
PBAS89-29
OB
grnt
478146
6375399
18
Lake Minto domain
Goudalie domain
G-1
PBA91-130
B
gdi
631837
6439468
18
G-2
PBA90-76
HB
gdi
625161
6412056
18
U-1
GS90-3
CHB
gdi
547461
6414553
18
U-2
GS90-11
CHB
gdi
547461
6414553
18
U-3
PBA90-94
OCHB
gdi
648708
6467660
18
U-4
PBA91-182
OCHB
gdi
641803
6481984
18
U-5
PBA91-181
OCH
gdi
643130
6484636
18
U-6
PBA91-179
OHBG
gdi
647114
6489676
18
U-7
PBA91-28
CHB
gdi
712181
6452275
18
U-8
PBA91-134
OCHB
gdi
725194
6456254
18
U-9
PBA90-101
HB
gdi
684029
6506757
18
U-10
PBA91-158
OCHGB
gdi
686154
6517101
18
U-11
PBA90-117
OCHB
gdi
716165
6510470
18
U-12
PBA90-113
OCB
dte
724397
6513653
18
U-13
PBA90-168
OCHB
gdi
745405
6497208
18
U-14
PBA90-124
HB
gdi
766412
6481823
18
U-15
PBA91-26
H
gbr
714751
6445389
18
U-16
GS1-2
CHB
gbr
547461
6414553
18
U-17
PBAS90-107
CH
gbr
716458
6488243
18
U-18
PBA91-161
B
gbr
676365
6516914
18
U-19
PBA91-184
OCH
gbr
695621
6475456
18
U-20
GS1-15
OCB
mzt
547461
6414553
18
U-21
GS1-16
CHB
mdt
547461
6414553
18
U-22
GS1-17
CH
dte
547461
6414553
18
Utsalik domain
1649
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 9
SEPTEMBER 2002
APPENDIX: continued
Sample
Archival
Mafic
Rock
number
number
mineral
type
UTMX
UTMY
Grid
zone
assemblage
Utsalik domain
U-23
PBA91-147
OCH
gbr
694263
6522266
U-24
GS1-5
CB
mdt
547461
6414553
18
18
U-25
GS1-6
CB
dte
547461
6414553
18
U-26
PBA91-29
CB
gbr
711749
6455842
18
U-27
PBA91-7
OCB
dte
703788
6466735
18
U-28
PBA91-9
OCB
gbr
703953
6468756
18
U-29
PBA91-19
CB
gbr
692218
6474672
18
U-30
PBA91-37A
H
dte
725271
6484428
18
U-31
PBA91-6
OCB
gdi
703354
6465619
18
U-32
PBA91-22
OCHB
gdi
698545
6476856
18
U-33
PBAS90-93
OCHB
gdi
655129
6444516
18
U-34
PBAS90-99
OCHB
gdi
664430
6450723
18
U-35
PBA91-30
OCHB
gdi
735474
6490192
18
U-36
PBA91-109
OCHB
gdi
727615
6517567
18
U-37
PBA91-159
OCHB
gdi
678369
6516739
18
U-38
PBA91-8
CB
gdi
703784
6467140
18
U-39
GS1-1
CB
gbr
547461
6414553
18
U-40
GS1-4
CB
gbr
547461
6414553
18
U-41
PBAS90-113
CHB
gdi
707392
6460073
18
U-42
PBAS 90-111
CHB
gdi
728692
6498024
18
U-43
PBA91-3
CHB
gdi
682749
6458805
18
U-44
PBA91-11
CHB
gdi
700267
6470539
18
U-45
PBA91-147b
CHB
gdi
694263
6522266
18
U-46
PBAS90-89
HB
gdi
681560
6466030
18
U-47
PBAS90-112
HB
gdi
666643
6451296
18
U-48
GS3-12
HB
gdi
547461
6414553
18
U-49
PBAS90-82
HB
gdi
671822
6457151
18
U-50
PBAS90-100
HB
gdi
665775
6452573
18
U-51
PBA91-155
HB
gdi
686692
6516998
18
U-52
PBA91-27
B
grnt
713641
6446660
18
U-53
PBAS90-101
HB
gdi
669716
6454498
18
U-54
PBA91-183
OCHB
gdi
693332
6474842
18
Mineral abbreviations: B, biotite; C, clinopyroxene; H, hornblende; O, orthopyroxene; G, garnet. Rock type abbreviations:
dte, diorite; gbr, gabbro; gdi, granodiorite; grnt, granite; mdt, monzodiorite; mzt, monzonite.
1650