Download Gabbroic Pegmatite Intrusions, Iberia Abyssal

JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 5
PAGES 885–905
2002
Gabbroic Pegmatite Intrusions, Iberia
Abyssal Plain, ODP Leg 173, Site 1070:
Magmatism during a Transition from
Non-volcanic Rifting to Sea-floor Spreading
JAMES S. BEARD1∗, PAUL D. FULLAGAR2 AND A. KRISHA SINHA3
1
VIRGINIA MUSEUM OF NATURAL HISTORY, 1001 DOUGLAS AVENUE, MARTINSVILLE, VA 24112, USA
2
DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF NORTH CAROLINA, CHAPEL HILL, NC 27599, USA
3
DEPARTMENT OF GEOLOGICAL SCIENCES, VIRGINIA TECH, BLACKSBURG, VA 24061, USA
RECEIVED FEBRUARY 21, 2001; REVISED TYPESCRIPT ACCEPTED NOVEMBER 7, 2001
On the Iberia Abyssal Plain (Ocean Drilling Program Site 1070),
gabbroic pegmatites and related rocks (127 ± 4 Ma, U–Pb zircon)
intrude upper mantle that was subsequently exposed and serpentinized
during Early Cretaceous non-volcanic rifting. The pegmatites include
a 3–4 m dike or sill (the ‘main’ pegmatite), numerous dikelets of
1–5 cm thickness, and clasts within the overlying ophicalcite breccia.
Exclusive of rodingitization, the main pegmatite contains 40–70%
calcic andesine, 25–35% kaersutitic amphibole (Mg# 60–70),
5–25% augite (Mg# 70–80) and 1–2% ilmenite. The dikelets
are more magnesian (Mg# up to 82 in kaersutite and 88 in
augite). Most indications are that the high Mg#s in the dikelets
reflect igneous compositions. Isotopic and elemental chemistry indicate
that the pegmatite-forming melt was enriched in incompatible
elements relative to normal mid-ocean ridge basalt, but not as
enriched as Azores basalts. The amphibole-bearing plagioclase
peridotites of the Iberia Abyssal Plain are an appropriate source for
the pegmatite melts. A combination of decompression accompanying
unroofing and heating from the upwelling asthenosphere beneath the
developing rift caused P–T conditions in the amphibole-bearing
lithosphere to exceed the dehydration-melting solidus (>1050°C),
producing small-volume, enriched, hydrous melts. Pegmatite intrusion
pre-dates unroofing at Site 1070 and post-dates syn-rift sedimentation and faulting of serpentinite seen inboard on the Iberia
Abyssal Plain. Thus, serpentinization and unroofing were timetransgressive and the age of the non-volcanic sea floor formed by
the unroofed mantle grows younger outboard, just as is the case for
normal, volcanic sea floor.
∗Corresponding author. E-mail: [email protected]
 Oxford University Press 2002
KEY WORDS:
abyssal plain; gabbro; Iberia; pegmatites; rifting
INTRODUCTION AND GEOLOGIC
SETTING
The west Iberian ocean–continent transition is a welldocumented example of a non-volcanic (a more precise
term might be magma-poor) rifted margin. Combined
seismic studies and drilling on Ocean Drilling Program
(ODP) Legs 103, 149, and 173 have revealed that the
oceanic basement beneath much of the Iberia Abyssal
Plain (IAP) consists of serpentinized peridotite, representing mantle rocks exhumed during continental rifting, and stranded fault blocks of continental crust (Boillot
et al., 1987; Krawczyk et al., 1996; Pickup et al., 1996;
Whitmarsh & Sawyer, 1996; Whitmarsh et al., 1998) (Figs
1 and 2).
Rifting at the margin began with the thinning of
the continental crust, probably in the latest Jurassic
(Whitmarsh et al., 2000; Whitmarsh & Wallace, 2001).
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VOLUME 43
NUMBER 5
MAY 2002
Fig. 1. Location map of the Iberia Abyssal Plain showing Site 1070 and other Ocean Drilling Program drill-hole sites from Legs 149 and 173.
Fig. 2. Schematic cross-section based on drill-hole data along Ocean Drilling Program Legs 149 and 173. Hole 1070 is the farthest outboard
site drilled to date on the Iberia Abyssal Plain.
886
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Studies of rocks from Sites 900, 1067, and 1068 on a
single basement high where Permian continental crust is
in fault contact with serpentinized peridotite (Figs 1 and
2) suggest a long period of rifting, dissection by several
generations of faults, and tectonic exhumation. At this
basement high (‘Hobby High’; Whitmarsh et al., 2000)
active tectonic dissection of the continental crust and
exhumation of mantle rocks was under way by 137 Ma
and continued for 10–20 my (Whitmarsh & Wallace,
2001; Manatschal et al., 2002). Site 1070 lies >100 km
outboard of Hobby High, with much of the intervening
oceanic basement being serpentinized peridotite (Fig. 2).
Linear magnetic anomalies in the vicinity of Site 1070
suggest the presence of ocean crust of 124–127 Ma age
(Whitmarsh & Miles, 1995), but no basalt or diabase was
found. Instead, drilling revealed pegmatoidal gabbro
intrusive into partially serpentinized peridotite. The basement rocks at Site 1070 were exhumed by late Aptian
(112–119 Ma) time (Whitmarsh & Wallace, 2001). Site
1070 lies 30 km east of the ‘J’ anomaly, a strong, linear
magnetic high found on both sides of the Atlantic (Rabinowitz et al., 1979; Tucholke & Ludwig, 1982; Masson
& Miles, 1984; Fig. 1). The ‘J’ anomaly is generally
considered to mark an episode of voluminous basaltic
magmatism dating from >120 Ma (i.e. between anomalies M0 and M1), just after the eruption of mid-ocean
ridge-type basalts (coincident with anomaly M2,
>122 Ma) off the Grand Banks (Tucholke & Ludwig,
1982; Vogt & Tucholke, 1989). Most workers (e.g. Tucholke & Ludwig, 1982; Whitmarsh et al., 1998) suggest
that ocean crust formation in the Grand Banks–Iberia
segment began at >125 Ma (anomaly M4 on the Southern Iberia Abyssal Plain), but, to date, no basalts of this
age have been recovered there.
Some syn-rift magmatism has been documented at
the west Iberian margin, especially at its northern and
southern ends. An undeformed gabbro–troctolite–diorite
lopolith on the Gorringe Bank, south of the IAP, has
yielded U–Pb zircon ages of 136–138 Ma, and younger
(122 Ma, U–Pb zircon) gabbros occur near the northern
end of the IAP on the Galicia Bank (Scharer et al., 2000;
Fig. 1). However, very few unequivocal syn-rift igneous
rocks have been recovered during drilling on the Iberia
Abyssal Plain proper. In particular, metamorphosed gabbroic rocks drilled at Sites 900, 1067, and 1068 on Legs
149 and 173 have now been shown to be Paleozoic in
age (Rubenach, 1999) and not syn-rift gabbro as originally
thought (e.g. Feraud et al., 1996). Leg 149 recovered
extremely weathered basalt and diabase clasts from serpentinite breccias in Hole 899B. These rocks have chemical characteristics similar to enriched mid-ocean ridge
basalt (EMORB) and probably reflect syn-rift sea-floor
magmatism (Seifert et al., 1997). However, the age and
source of these rocks remains unknown, and their exceptionally severe and pervasive alteration makes their
further study (as igneous rocks) problematic.
Basement at Site 1070 consists of a calcite-cemented
‘jigsaw’ breccia reminiscent of Alpine ophicalcites (e.g.
Bernoulli & Weissert, 1985) and apparently formed moreor-less in place during low-temperature alteration and
serpentinization (Skelton & Valley, 2000). This is immediately underlain by a thick (3–4 m recovered thickness
over a 10 m interval), relatively fresh, gabbro pegmatite
dike intruding the underlying serpentinite (Whitmarsh et
al., 1998). Beneath this is relatively massive, partially
(>70–95%) serpentinized peridotite. In addition to the
large dike, pegmatoidal gabbro also occurs as clasts in
the upper breccia and as strongly rodingitized dikelets
in the underlying serpentinite. Apart from serpentinite,
the only other rock type found in the Hole 1070A
basement is albitite, which, like the pegmatoidal gabbro,
occurs as small dikes in serpentinite and clasts in the
upper breccia.
The gabbro pegmatite yields Ar–Ar cooling ages of
119–121 Ma on igneous amphibole (kaersutite) (Turrin,
1999; Manatschal et al., 2002). We report here a new
U–Pb (ion probe on zircon from associated albitites) age
of 127 ± 4 Ma, which we interpret as an igneous age.
The Site 1070 gabbro pegmatites are the only igneous
rocks that are demonstrably coeval with rifting in the
southern IAP. As such, they help to document the transition from non-volcanic margin to normal ocean crust
formation in this part of the North Atlantic Basin. Because
the pegmatites represent magmas formed at a time and
place close to the onset of true mid-ocean ridge magmatism, they can also provide information on the processes involved in the inception of magmatism in oceanic
rift systems.
METHODS
887
The major element chemistry of all minerals was determined by electron microprobe in the Department of
Geological Sciences at Virginia Tech. Separates for trace
element and isotopic analysis were prepared by hand
selecting pegmatite fragments rich in a particular mineral,
coarsely crushing the fragment and hand picking a rough
mineral separate. Green and zoned amphibole fragments
were excluded from the rough separate. The separates
were crushed to 100 mesh in an agate mortar under
acetone, run through a magnetic separation and inspected
for obvious impurities. Feldspar separates were then
leached in dilute HF for 30 min and rinsed in deionized
water. One split of all separates was analyzed by Actlabs
using inductively coupled plasma (ICP) (for most transition metals) and ICP-mass spectrometry (ICP-MS) for
all other elements.
JOURNAL OF PETROLOGY
VOLUME 43
A second split was analyzed for Sr, Nd, and Pb isotopes.
Isotopic tracer solutions (either 150Nd, 147Sm, 87Rb, and
84
Sr spikes, or a mixed 205Pb/233–236U spike) were added
to samples before dissolution. Samples were dissolved in
Savillex vials using HF and HNO3. Except for plagioclase
samples, AG50Wx8 cation exchange resin was used to
separate Rb, Sr, and bulk rare earth element (REE)
fractions from solutions using HCl. The bulk REE fraction was adsorbed onto AG50Wx4 resin equilibrated
with alpha-HIBA, and Nd and Sm were selectively eluted
with alpha-HIBA. The Sr fraction was purified using
EiChrom SrSpec resin. Separation of Pb and U from
plagioclase samples employed standard HBr and HCl
anion exchange chemistry.
The mass spectrometer used for the analyses is a VG
(Micromass) Sector 54 thermal ionization mass spectrometer in the Department of Geological Sciences, University of North Carolina at Chapel Hill (UNC-CH). Nd
was measured as the metal species using a dynamic
multicollector mode. Samarium was analyzed in static
multicollector mode. Mass fractionation was corrected
using 146Nd/144Nd = 0·7219 and 149Sm/152Sm =
0·51685. Recent results for standard isotopic reference
materials: La Jolla Nd standard, 143Nd/144Nd =
0·511855; Ames Nd metal, 143Nd/144Nd = 0·512140;
BCR-1, 143Nd/144Nd = 0·512631. Sr was analyzed in
dynamic multicollector mode. 87Sr/86Sr ratios were corrected for mass fractionation using an exponential mass
fractionation law. All 87Sr/86Sr ratios from the UNC-CH
laboratory are reported relative to a value of 0·710250
for this NBS 987 standard (typically, four standards were
analyzed with each turret of samples; if the 87Sr/86Sr
ratio for the four standards averaged 0·710225, a value
of 0·000025 was added to the ratio for each sample from
that turret). Internal precision (standard error) for Sr
analyses is typically 0·000006–0·000010. Isotopic compositions of Pb and U are measured in multicollector
static mode. Uncertainties (2) for measured Pb isotopic
ratios are typically below 0·1%. Recent replicate analyses
of SRM 981 Pb yield 206Pb/204Pb = 16·897 ± 9, 207Pb/
206
Pb = 0·91352 ± 15, and 208Pb/206Pb = 2·1617 ±
9. Typical total procedural blanks are insignificant.
U and Pb isotopic abundances in zircon (Table 1) were
measured using the Cameca IMS 1270 ion microprobe in
the Department of Earth and Space Sciences at UCLA
following the method of Schumacher et al. (1994).
Because of reasonable counting statistics associated with
206
Pb (as opposed to 207Pb), we base our age estimates on
206
Pb/238U.
ZIRCON U–Pb AGES
A large (1 mm) zircon was hand picked from a partially
altered albitite clast sampled from the autochthonous
NUMBER 5
MAY 2002
breccia above the main pegmatite at Site 1070 (Whitmarsh et al., 1998). The zircon was hosted by a large
(>1 cm) single grain of largely unaltered sodic plagioclase.
Similar zircons have been found in several other albitite
clasts and dikes.
Six spot analyses were attempted on the zircon; four
of these yielded statistically significant results. Three of
these spots yielded 206Pb/238U ages of 124·1, 126·8, and
130·2 Ma (all ±7 Ma, Table 1); the fourth yielded an
age of 112 Ma. The three older dates (average 127 Ma,
range consistent with all analyses, 123–131 Ma) we interpret as reflecting the igneous age of the gabbroic
pegmatite. This age is similar to the ocean crust age
predicted from magnetic anomalies in the vicinity of Site
1070 (Whitmarsh & Wallace, 2001) and also intermediate
between the Gorringe Bank (137 Ma) and Galicia Bank
(122 Ma) gabbro ages. This is consistent with the northto-south progression that has been proposed for the onset
of magmatism at the west Iberian margin (Scharer et al.,
2000). We interpret the 112 Ma age as reflecting lead
loss, possibly incurred during serpentinization close to
the time of exhumation (e.g. Whitmarsh & Wallace,
2001).
PETROGRAPHY AND MINERAL
CHEMISTRY
Overall
The Site 1070 gabbro pegmatites include a large (recovered thickness 3–4 m) dike or sill (the ‘main’ pegmatite), numerous small (1–5 cm thick) dikelets intrusive
into the serpentinite basement, and fragments within the
overlying parautochthonous serpentinite breccia. Exclusive of alteration products, the main pegmatite consists
of 40–70% plagioclase, 25–35% red–brown amphibole,
5–25% augite, and 1–2% ilmenite (Whitmarsh et al.,
1998). Overall, the main pegmatite has igneous textures
resembling an extremely coarse-grained (grain sizes up
to 7 cm) amphibole gabbro. Conspicuous red–brown
amphibole poikolitically encloses equant to tabular
plagioclase and replaces augite. Rodingite-like (e.g. zoisite, prehnite, analcime) alteration is widespread. In the
main pegmatite, rodingite zones and patches are evident
in all thin sections and hand samples, with plagioclase
being particularly strongly affected. In the clasts, and
especially the thin dikelets, rodingitization may be pervasive or spare only remnant pyroxene, kaersutite, and
ilmenite. Fresh plagioclase was not found in any of the
small gabbroic dikelets. In addition to low-temperature
alteration, the main pegmatite is locally deformed in
discrete, high-temperature, shear zones. Within the shear
zones, igneous composition plagioclase and amphibole
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BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Table 1: U–Pb ion probe data from ODP Site 1070 zircon [ODP sample 1070A-8R4, 30–33 cm, piece
3 (albitite)]
Spot no.
207
Pb/235U
Abs. SE
206
Pb/238U
206
Abs. SE
Pb/238U
Error
207
Pb/235U
Error
age (Ma)
(Ma)
age (Ma)
(Ma)
119·9
1
0·1124
0·1314
0·0204
0·001202
130·2
7·591
108·2
2
0·1393
0·07255
0·01987
0·001104
126·8
6·979
132·4
3
–0·04555
0·1274
0·01765
0·001199
112·8
7·596
–47·34
4
0·1732
0·106
0·01944
0·001059
124·1
6·695
162·2
64·66
135·6
91·71
SE, standard error.
(i.e. andesine–labradorite and kaersutite) were recrystallized, suggesting deformation occurred at or near
granulite-facies conditions (Whitmarsh et al., 1998).
In addition to the gabbroic rocks, several albitite breccia
clasts and dike fragments were recovered. The albitites
consist of 95% coarsely crystalline, mildly to pervasively
altered, weakly zoned sodic plagioclase with minor altered
biotite, altered amphibole (tremolite), and accessory zircon. They contain no quartz.
within large amphibole crystals. Within the remnants,
patchy replacement of augite by amphibole along fracture
or cleavage planes is common. Crystallographically oriented ilmenite plates, presumably formed during subsolidus exsolution, occur in many augite grains. The
augite is moderately aluminous (2–3% Al2O3) and moderately titanian (0·5–1% TiO2) (Table 2). In the main
pegmatite, augite has Mg# ranging from 72 to 77,
whereas that in breccia clasts and small pegmatite dikelets
is substantially more magnesian (Mg# 82–88) (Table 2,
Fig. 3).
Feldspar
Within the main pegmatite body, plagioclase occurs as
large (up to 5 cm), weakly zoned, euhedral to subhedral
crystals. Alteration ranges from minor recrystallization
along cleavage planes to complete replacement. Alteration minerals include prehnite, zoisite, zeolites(?), and
analcime, but not albite. Potassium feldspar (Or98, rare)
is probably an alteration phase, but late crystallization
as an igneous phase cannot be ruled out. Plagioclase in
breccia clasts and in the smaller gabbroic dikelets is
usually pervasively altered. In the main pegmatite, plagioclase compositions (average of all analyses; An46·5Ab52·2Or1·3) range from An28 to An53, with most rims
An35–47 and most cores An43–51. The orthoclase component
of the plagioclase typically ranges from 1 to 2 mol %
(Table 2).
Plagioclase in the albitites is coarse-grained (commonly
>1 cm), albite–sodic oligoclase (An6–12; Table 2). The sodic
plagioclase in the albitites appears to be somewhat more
resistant to low-temperature alteration than the calcic
plagioclase in the gabbros. The only fresh plagioclase
preserved in the breccia clasts and dikelets is in the albitites.
Amphibole
Bright red–brown kaersutite (pargasitic hornblende with
up to 5·3 wt % TiO2; Fig. 4, Table 2) occurs as thick
mantles about augite and as large (up to 7 cm) anhedral,
interstitial to poikoblastic crystals enclosing all other
igneous phases. The kaersutites are F-bearing hydroxylamphiboles (Table 2). Many amphibole grains are weakly
zoned with Al- and Ti-poor, magnesian hornblende at
grain boundaries and along fractures. This type of zoning
probably represents subsolidus re-equilibration. A few
grains, however, are spectacularly zoned from reddish
kaersutite cores to deep green (sometimes nearly black),
Na-, Fe-, Al-rich, Ti-poor hastingsite or ferropargasite
rims (Fig. 5, Table 2). This zoning, also characterized
by a drop in Mg# from 65 in the kaersutite core to 29
at the rim, appears to be an igneous feature. Amphibole
in the smaller dikes and clasts is more magnesian than
that in the main pegmatite (Mg# up to 82). Unlike
secondary amphiboles that are clearly related to alteration, however, the amphibole in the clasts and dikes
retains its Ti-rich and pargasitic composition (Fig. 4).
Clinopyroxene
Ilmenite
Pinkish brown clinopyroxene (augite) is an early crystallizing phase that is partially to wholly replaced by late
igneous amphibole. It occurs as irregular, remnant cores
Rounded grains of ilmenite [il/(il + hm) = 88–92] up
to 1 cm in diameter occur in all samples, commonly as
inclusions in amphibole, augite, and plagioclase. The
889
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MAY 2002
Table 2: Major-element analyses of minerals
Amphibole
Clinopyroxene
Main pegmatite
dikelet
albitite
main pegmatite
dikelet
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
9R-2
9R-2
9R-2
9R-1
9R-1
14R-3
8R-4
9R-2
9R-2
9R-1
14R-3
22 cm
22 cm
64 cm
75 cm
82 cm
85 cm
30 cm
22 cm
64 cm
45 cm
85 cm
Fe-parg
meta.
meta.
SiO2
42·72
39·80
42·81
41·65
49·54
44·87
53·29
53·52
52·24
53·05
53·31
TiO2
5·19
0·27
4·64
4·28
0·32
4·30
1·26
0·37
0·63
0·63
0·87
Al2O3
11·87
15·59
11·24
11·78
6·82
11·40
3·85
1·21
2·25
2·35
2·70
Cr2O3
0·04
0·03
0·06
0·05
0·07
0·20
0·08
0·01
0·04
0·04
0·17
MgO
12·72
5·78
12·41
13·47
15·42
16·18
21·37
14·86
14·10
14·36
16·50
CaO
11·17
11·68
11·38
10·95
12·63
11·78
11·39
21·13
21·62
21·36
21·74
MnO
0·24
0·11
0·23
0·29
0·17
0·16
0·04
0·28
0·33
0·32
0·19
FeO
11·62
21·99
12·89
12·11
11·14
6·77
3·67
8·93
8·31
8·71
4·69
Na2O
2·94
3·13
2·77
2·71
1·88
2·70
1·77
0·48
0·57
0·55
0·66
K2O
0·58
0·25
0·62
0·63
0·10
0·64
0·09
0·00
0·02
0·00
0·04
F
0·43
0·00
0·32
0·61
0·20
0·09
0·00
Total
99·53
98·63
99·37
98·51
98·29
99·08
96·81
100·78
100·10
101·36
100·85
Mg#
66·1
31·9
63·2
66·5
71·2
81·0
91·2 En
42·3
41·0
41·4
47·3
Fe2O3
A-site
0·67
0·89
0·68
0·63
0·49
0·65
0·21 Fs
14·3
13·6
14·1
7·6
AlIV
1·80
1·91
1·74
1·85
0·86
1·64
0·64 Wo
43·3
45·3
44·4
45·0
Ilmenite
Plagioclase
main pegmatite
dikelet
main pegmatite
albitite
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
9R-2
9R-2
9R-1
14R-3
9R-2
9R-2
9R-2
9R-2
9R-2
9R-1
8R-4
22 cm
64 cm
45 cm
85 cm
64 cm
SiO2
0·05
0·02
0·02
0·06
TiO2
49·01
47·82
48·75
53·69
Al2O3
0·09
0·11
0·14
0·04
Cr2O3
0·05
0·05
0·06
0·14
MgO
2·30
3·64
3·54
6·88
22 cm
22 cm
64 cm
22 cm
rim
core
rim
core
75 cm
30 cm
calcic
sodic
58·52
56·59
57·53
56·04
55·25
58·60
66·87
26·75
27·95
27·48
28·33
28·50
25·90
21·42
0·03
0·04
0·02
0·02
0·02
0·04
0·03
8·61
10·14
9·46
10·44
11·06
7·87
2·00
0·26
0·32
0·18
0·27
0·24
0·13
0·19
Na2O
6·57
5·75
6·21
5·63
5·22
6·95
9·81
K2O
0·31
0·29
0·27
0·23
0·19
0·13
0·26
CaO
MnO
0·67
0·63
0·59
0·76
FeO
39·35
35·95
36·95
35·32
8·54
10·92
10·64
3·01
Fe2O3
F
Total
100·06
99·14
100·70
101·04
101·07
101·13
100·95
100·48
99·62
100·58
81·9
74·8
75·8
70·9 An
41·3
48·6
45·0
49·9
53·3
38·2
10·0
Hem
8·0
10·2
9·8
2·7 Ab
57·0
49·8
53·5
48·8
45·6
61·1
88·5
Gk
8·4
13·4
12·9
24·5 Or
1·8
1·6
1·5
1·3
1·1
0·8
1·5
Ilm
99·90
Fe-parg, Fe-pargasite; meta., metamorphic.
890
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Fig. 3. Mg# vs Ti (+) and Al (Ε) in clinopyroxene. Plotted points are average analyses for individual samples. The fields outline the range of
spot analyses. It should be noted that the main pegmatite is substantially less magnesian than the clast and dike samples.
Fig. 4. Mg# vs Ti (+), Na (stars), and AlIV (Ε) in igneous amphibole (kaersutite). Plotted points are average analyses for individual samples.
The fields outline the range of spot analyses. As with pyroxene, the clasts and dike are more magnesian than the main pegmatite. It should be
noted that even the most magnesian amphiboles are still kaersutite or titanian pargasite.
ilmenite is magnesian (geikelite component up to 14%
in the main pegmatite and 25% in one small dikelet) and
is usually zoned in Mg. There are two sets of very fine
(width usually <5 m) exsolution lamellae. One set is
hematite-rich (approximately hm25) ilmenite. The other,
much more sparsely developed, is baddeleyite (ZrO2).
891
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MAY 2002
Fig. 5. Mg#, Ti, AlIV and A-site occupancy variation with distance across the rim of a zoned amphibole (kaersutite core, hastingsite and ferropargasite rim) from the main pegmatite. This type of zoning, although optically and compositionally spectacular, is uncommon and has been
seen only in the main pegmatite.
Other minerals
Zircon occurs as rather large (up to 1 mm) grains in the
albite rocks. A strongly chloritized phyllosilicate that may
have once been biotite occurs in these rocks and in a
single sample from the main pegmatite. Sphene occurs
sporadically as epitaxial growths on ilmenite and is a late
magmatic or subsolidus phase.
Trace element chemistry of mineral
separates
Plagioclase contains 200–300 ppm Ba and 800–1300
ppm Sr, with Sr/Rb averaging >250. Despite acid
leaching, K in two out of three bulk analyses is higher
than the averaged microprobe data (Tables 2 and 3),
whereas Rb in a small (10 mg) split analyzed by isotope
dilution is five times (31 vs 6 ppm) that of a larger
(100 mg) sample analyzed for Rb by ICP. This suggests
that, despite leaching, a K- and Rb-rich component,
probably K-feldspar, was present in at least one and
possibly two of the separates. Concentrations of most
other trace elements are low, especially compared with
amphibole. Plagioclase is light REE (LREE) enriched
with a large positive Eu anomaly and very low abundances
of heavy REE (HREE) (Fig. 6).
Amphibole contains substantial Nb, transition metals
and REE, and is probably the major reservoir for most
of these elements in the pegmatite. Both LREE and (to
a lesser extent) HREE are depleted with respect to middle
REE (MREE). There is a small negative Eu anomaly (Fig.
6). The three amphibole separates yield exceptionally
consistent results for all trace elements for both ICP and
isotope dilution, although Rb from isotope dilution (2–3
ppm) is slightly lower than the value (4 ppm) from ICP.
Both K and Ti from ICP are consistent with microprobe
results (Tables 2 and 3). The average of the three analyses
is taken as a realistic estimate of the average kaersutite
composition in the main pegmatite.
The single augite separate has element distribution
patterns (e.g. chondrite-normalized REE) similar to the
amphibole, but much lower abundances for most elements. An exception to this is compatible transition
metals, which are as abundant or (in the case of Cr)
more abundant in augite than in amphibole. A problem
with the augite separate is its likely contamination by
very fine included septa and stringers of replacement
kaersutite. Both K and Ti measured in the separate are
higher than the averaged microprobe analyses, enough
so to suggest 10–15% amphibole contamination. High-K
and high-Ti augite analyses were screened and eliminated
during compilation of the microprobe average as a check
on amphibole contamination. However, the high Ti in
the analysis of the mineral separate may reflect, in part,
ilmenite exsolution rather than amphibole contamination.
The single ilmenite separate has low abundances of
most trace elements with several notable exceptions.
The high concentrations of large ion lithophile elements
892
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Table 3: Trace element analyses of mineral separates (values in ppm)
Sample:
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
9R-3
9R-2
9R-1
9R-2
9R-2
9R-1
9R-2
9R-1
20 cm
28 cm
75 cm
64 cm
28 cm
75 cm
64 cm
45 cm,
Mineral:
plag
plag
plag
amph
amph
amph
cpx
ilm
Sr
1286
928
802
166
148
142
35
12
K
8050
7139
4068
3984
4233
4233
664
913
Rb
6
4
3
4
4
4
2
3
Ba
212
287
230
193
161
170
116
92
Cs
0·2
0·1
0·1
0·1
Th
0·1
0·0
0·1
0·39
0·35
0·34
0·3
U
0·03
0·02
0·04
0·14
0·09
0·11
0·18
Nb
27·8
Ta
P
Zr
26·9
1·94
1·97
27·3
2·01
2·2
0·1
130
130
130
218
262
699
130
9
10
10
90
109
108
57
Hf
0·2
Ti
700
840
Y
2
2
Ge
2·8
2·2
Ga
0·1
16
17
4·1
846
1·1
18
4·1
3·9
2·2
0·34
0·11
54
3·48
197
2·3
27671
29981
28665
6127
297843
83
90
86
47
5
4·4
4·5
16
17
47·6
3·9
16
53·2
5·3
2·6
7
4
Co
53
Ni
160
146
117
114
V
910
996
977
557
2085
Cr
142
154
106
233
230
109
108
108
140
Sc
0·7
1·1
0·7
La
1·96
2·08
2·70
Ce
3·61
3·83
4·54
Pr
0·40
0·41
0·47
Nd
1·65
1·73
1·75
34·5
35·4
34·4
Sm
0·25
0·26
0·25
12·1
12·3
11·9
Eu
0·97
1·15
1·32
Gd
0·23
0·23
0·22
8·79
7·67
28·6
28·5
5·45
5·53
4·01
3·97
15·7
15·7
Tb
0·03
0·03
0·03
Dy
0·13
0·13
0·11
Ho
0·03
0·02
0·02
3·34
3·39
Er
0·07
0·06
0·04
8·9
9·01
1·15
Tm
2·88
2·88
16·5
7·61
27·4
5·31
3·78
14·9
4·99
108
56·3
4·14
12·7
4·6
2·3
0·45
14·5
5·36
1·49
0·3
1·74
0·13
7
0·34
1·32
0·06
7·88
0·37
3·17
1·59
0·08
8·45
4·45
0·25
1·19
1·12
0·60
0·05
16·7
2·73
45·7
15·8
Yb
0·06
0·05
0·03
6·75
7·06
6·58
3·67
0·46
Lu
0·008
0·007
0·004
0·893
0·923
0·862
0·508
0·101
(LILE), especially K and Ba, undoubtedly reflect inclusions of a K-rich phase, probably orthoclase replacing
plagioclase. Such inclusions are exceptionally difficult to
detect in an opaque mineral. A large (1 g) sample
processed for isotopic analysis yielded (by isotope dilution)
much lower Rb (0·05 vs 3 ppm) and Sr (3·5 vs 12ppm)
than the smaller (100 mg) separate used for ICP, consistent with feldspar contamination of the latter split. The
high Zr content of the ilmenite is consistent with observed
baddeleyite exsolution lamellae. High concentrations of
Nb and transition metals, especially V, are not unexpected. REE concentrations are low, and the
893
JOURNAL OF PETROLOGY
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NUMBER 5
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Fig. 6. Chondrite-normalized (Sun & McDonough, 1989) rare earth elements in mineral separates from the main pegmatite.
chondrite-normalized pattern probably reflects mixing of
an ilmenite signal (possibly HREE-rich) with feldspar
and/or other contaminants (Fig. 6).
MODELS AND ESTIMATES OF MELT
COMPOSITION
We have estimated the whole-rock chemistry of the 1070
gabbro pegmatites by two methods: modal mass balance
and partition coefficient inversion. Modal models are
end-member calculations that tacitly assume that the
minerals in the rock are not cumulate, but reflect a
bulk liquid composition. Partition coefficient modeling
provides another end-member by assuming that the
minerals are entirely cumulate. Taken together, the two
end-member models provide some constraints on likely
liquid compositions.
Isotope geochemistry
Sr, Nd, and Pb isotopic compositions were determined
on splits of the mineral separates that were analyzed for
trace elements. Nd and Sr were measured in amphibole
and pyroxene; Sr and Pb in plagioclase. Results are given
in Table 4. Nd values for pyroxene and amphibole cluster
in a tight group with initial (125 Ma) Nd of 4·7–4·9
(modern values 5·0–5·3). For amphibole, pyroxene and
plagioclase, initial 87Sr/86Sr ranges from 0·70362 to
0·70463. Two of the amphiboles have initial 87Sr/86Sr
<0·70366; all other analyses yielded initial 87Sr/86Sr
>0·704. The feldspar has average 206Pb/204Pb, 207Pb/
204
Pb, and 208Pb/204Pb, respectively, of 18·65; 15·48, and
38·3 (Table 4).
Nd and Sr were also analyzed in the ilmenite separate,
which has the lowest concentrations of Sr, Rb, Nd and
Sm, and the highest initial 87Sr/86Sr (0·70623) and lowest
Nd (4·4) of any sample analyzed for isotopes. Given the
low concentrations of REE in the ilmenite (Table 3), Nd
probably lies within measurement error of the other
separates. The anomalously radiogenic ilmenite Sr presumably reflects contamination, probably with feldspar
alteration products (Fig. 7).
Modal estimates of major and trace
element composition
The coarse-grained nature of the pegmatite and the
necessarily small size of the samples precludes meaningful
determination of a thin-section-based mode. The modal
estimates used here to calculate bulk-rock chemistry rely
heavily upon shipboard modal data. These modes are
based on visual inspection of the whole core and on a
semiquantitative scanned image analysis (Whitmarsh et
al., 1998). The volume modes are then converted to mass
(wt %) modes for the whole-rock calculation.
The general relationship used for calculating wholerock chemistry from modal data in these rocks is
894
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Table 4: Isotopic compositions of mineral separates
Sample:
Mineral:
87
Sr/86Sr
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
1070a-
9R-2
9R-2
9R-1
9R-2
9R-3
9R-1
9R-2
9R-1
28 cm
64 cm
75 cm
28 cm
20 cm
75 cm
64 cm
45 cm
amph
amph
amph
plag
plag
plag
cpx
ilm
0·703656
0·704273
0·703618
0·512909
0·512911
0·512893
0·704322
0·704625
0·704005
143
Nd/144Nd
206
Pb/204Pb
18·63
18·73
18·57
207
Pb/204Pb
15·50
15·51
15·45
208
Pb/204Pb
38·30
38·39
38·19
0·704461
0·706227
0·512912
0·512867
Sr and Nd are initial ratios calculated at 127 Ma. Pb ratios are uncorrected.
Fig. 7. (a)
Ciwr = Xpl × Cipl + Xcpx × Cicpx + Xamp × Ciamp
+ Xilm × Ciilm
mode. The third mode is also mafic, but more amphibole
rich.
where X is the mass fraction of the phase (plagioclase,
clinopyroxene, amphibole, ilmenite) and Ci is the concentration of a given element i. The accuracy of the
calculation is limited by uncertainties in the mode, by
analytical error, and, in the case of trace elements, by
the purity of the separates.
Three calculated whole-rock compositions based on
modal data are presented (Table 5). One is based on a
plagioclase-rich mode, the second, on a more mafic
Trace element estimates based on partition
coefficients
895
Any attempt to reconstruct melt compositions by inverting partition coefficients will involve large errors. We
have attempted to manage the error in our calculations
in several ways. First, we chose a wide range of possible
partition coefficients, based on experimental studies of
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 5
MAY 2002
Fig. 7. (a) Covariation of Sr and Nd isotopes. MAR, Mid-Atlantic Ridge; OFZ, Oceanographer Fracture Zone. (b) Covariation of Pb isotopes.
MORB fields from Staudigel et al. (1984) and Wilson (1989); OFZ data from Shirey et al. (1987); Azores data from Turner et al. (1997).
Results of modeling
basalts or ion probe phenocryst–matrix studies for all
four major mineral phases (Nielsen, 2000, GERM database). We complemented this with data from one widely
accepted compilation (Rollinson, 1993). Second, we used
the partition coefficients to frame a range of potential
melt compositions, rather than try to select a single ‘best’
partition coefficient value for any given element. Third,
we interpret this range largely on the basis of relative,
rather than absolute, element abundances. Model melt
compositions are given in Table 6. The partition coefficients and primary data sources are given in the
Appendix.
896
All modal models yield broadly basaltic major element
compositions. Absolute and relative quantities of most
elements vary largely as a function of chosen mineral
mode. TiO2 ranges from 2·2 to 3·4% and exceeds 1·5%
in most models even when ilmenite is (unjustifiably)
excluded from the mode. Na2O is nearly 3·2% even in
the most plagioclase-poor model.
In the modal models, most trace element abundances
are controlled by amphibole. In model 3, for example,
70–85% of Nb, Ta, Y and REE (except La and Eu)
reside in amphibole. Partly in consequence of this, the
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Table 5: Modal estimates of bulk
composition
Table 6: Models of trace-element composition
for bulk rock
%Plag:∗
70
40
45
Low Kd
High Kd
Modal
%Cpx:
4
30
18
model
model
model 3
%Amph:
25
28
35
%Ilm:
1
2
2
Sr
1875
111
406
K
12694
5101
2713
SiO2
51·85
49·71
49·05
Rb
80
TiO2
2·17
3·11
3·42
Ba
1519
4·3
162
2·74
188
Al2O3
21·50
13·89
15·84
Cs
MgO
4·32
8.l5
7·69
Th
18
0·35
0·23
CaO
10·42
13·48
12·12
Ta
25
1·29
0·86
MnO
0·09
0·18
0·16
Nb
342
23
FeO
4·76
7·93
7·82
Ce
317
34
4·7
Na2O
4·80
3·17
3·61
Zr
814
66
59·2
K2O
0·34
0·27
0·32
Hf
18
2·6
F
0·09
0·09
0·12
Sm
33
6·7
Total
100·32
100·33
100·16
Mg#
61·77
65·66
63·66
Modes are given in volume percent, but composition calculated based on weight percent.
calculated whole-rock REE have a concave-down pattern
that parallels amphibole (Figs 6 and 8). In this model
plagioclase accounts for 86% of the Sr, 40–60% of most
other LILE, and >20% of the La and Eu.
The model based on the highest partition coefficients
(i.e. those yielding the lowest abundances of a given
element in the melt) resembles the modal model, except
that it predicts slight LREE enrichment (Fig. 8) and yields
lower values for compatible elements. This similarity is
not surprising, as both models would be expected to yield
near minimum values. The model based on the lowest
partition coefficients predicts very high concentrations,
especially for incompatible elements.
DISCUSSION
Comparisons with Iberian margin and
other oceanic rocks
The isotopes and modeled bulk compositions of the Site
1070 gabbro pegmatites are similar to EMORB in most
respects. The pegmatites are generally less enriched than
Azores basalts and, even allowing for fractionation, more
enriched than most normal MORB (NMORB).
The variation in Sr isotopes in the Site 1070 pegmatite
mineral separates is consistent with contamination with
Cretaceous seawater ( 87Sr/86Sr = 0·70725; Veizer et al.,
1999; Fig. 7a). There is little variation in Nd isotopic
1·90
1·90
0·11
12·3
2·1
5·6
Ti
33072
9281
21629
Y
216
52
42
Yb
14
4·14
Sc
173
26
69
Cr
137
7
102
Ni
76
8
74
La
166
15
Pr
Nd
3·3
4·98
2·65
121
26
16·47
Eu
12·7
2·5
Gd
25
7·7
7·19
2·02
1·32
8·1
7·66
1·1
1·55
Tb
Dy
Ho
Er
3·14
26
7·5
16
5·1
Tm
Lu
2·26
4·16
0·55
2·8
0·57
0·44
Kd models based on partition coefficients given in Table A1
(Appendix). Modal model is model 3 from Table 5.
composition. In an Sr–Nd isotopic diagram, the samples
with the least radiogenic Sr plot at the juncture of overlap
amongst (non-plume?) EMORB from the Oceanographer
Fracture Zone (Shirey et al., 1987), basalts from the
Azores (Turner et al., 1997), and the least depleted Atlantic
NMORB (Staudigel et al., 1984) (Fig. 7a). Feldspar Pb
isotopes are similar both to NMORB and to EMORB
from the Oceanographer Fracture Zone. They are less
radiogenic than basalts from the Azores (Fig. 7b).
Analyses of metagabbro from the Iberia Abyssal Plain
yield Nd ranging from +6 to +10 (Seifert et al., 1997;
Cornen et al., 1999). However, these rocks are now
897
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MAY 2002
Fig. 8. Chondrite-normalized (Sun & McDonough, 1989) REE for the model melt compositions.
recognized as Paleozoic continental basement (Rubenach, 1999). On the Galicia Bank, basaltic and gabbroic rocks have Nd of +2·2 to +8·8, with older rocks
(e.g. 122 Ma) having less depleted values than younger
(e.g. 100 Ma) rocks (Charpentier et al., 1998). This trend
was interpreted by Charpentier et al. (1998) as reflecting
a decreasing lithospheric and increasing asthenospheric
component with time. Young (60–66 Ma) alkaline basalts
from the Gorringe Bank have initial 87Sr/86Sr of 0·7031
and Nd of +6·6 (Bernard-Griffith et al., 1997). A dolerite
from the west end of Gorringe Bank has Nd of +5·1
and initial 87Sr/86Sr of 0·70484 (Cornen et al., 1999).
Scharer et al. (2000) reported strongly depleted Hf of
+19·5 to +20·5 for 135 Ma gabbroic rocks from Gorringe Bank and slightly less depleted values (+14·0 to
+14·6) for 122 Ma gabbro on the Galicia Bank.
Despite the uncertainties in the bulk compositional
models, several general characteristics of the magma
from which the pegmatites formed can be recognized
(Figs 8 and 9):
(1) the model based on the minimum partition coefficients (i.e. the model yielding the highest concentrations of the modeled elements) appears to be largely
unrealistic (predicting, for example, 340 ppm Nb). However, abundance patterns in all three models are roughly
parallel for most elements.
(2) None of the models exhibit the strong LREE depletion typical of NMORB. As noted above, however,
chondrite-normalized La/Yb is higher in the partition
coefficient models (3–10) than in the modal models (about
unity) (Fig. 8).
(3) All models are enriched in the most incompatible
elements with respect to the NMORB of Sun & McDonough (1989). For many strongly incompatible elements,
this enrichment is greater than can be accommodated
by any realistic (e.g. 75%) degree of fractionation (Fig.
9).
(4) The modal and high-Kd models resemble the
EMORB of Sun & McDonough (1989).
(5) The least radiogenic (i.e. excluding lavas from São
Miguel) Azores samples of Turner et al. (1997) are mostly
bracketed by the models.
(6) Nb and Ta are not depleted in the models.
(7) Yb and Y are unique in that they are higher in all
models than in NMORB, EMORB, or Azores lavas (Fig.
9).
Highly altered basalts, of unknown age, but interpreted
as syn-rift by Seifert et al. (1997), were recovered from
ODP Site 899, >40 km east of Site 1070. Some of these
basalts are Ti rich (>3% TiO2) and contain Ti-augite,
kaersutite and Ti-rich biotite (Cornen et al., 1999). Most
of the basalts are LREE enriched and have been interpreted as coming from a somewhat enriched mantle
source (Seifert et al., 1997; Cornen et al., 1999). Gorringe
Bank includes a younger series of 60–77 Ma alkaline
basalts, diorites, and gabbros, and an older series (137 Ma)
of cumulate gabbros and dioritic, doleritic and basaltic
differentiates. These latter rocks have flat to slightly
LREE-enriched REE patterns, interpreted by Cornen et
al. (1999) as transitional MORB ranging to a mixed
MORB–OIB (ocean-island basalt) source.
898
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Fig. 9. Modal model number 3 (Table 5; Ε), average Azores basalt (Φ; Turner et al. (1997), excluding São Miguel) and partition coefficient
models (shaded field) normalized to (a) NMORB and (b) EMORB of Sun & McDonough (1989).
Many mafic igneous rocks, including gabbro, diorite,
and basalt, have been recovered from the Galicia Bank.
These include sheared kaersutite–plagioclase diorite dikes
with coarse (greater than centimeter-sized) relict kaersutite porphyroclasts and accessory biotite and ilmenite
(Beslier et al., 1990). Many Galicia Bank basaltic rocks
also contain accessory brown amphibole and biotite and
show mild enrichment in LREE (Charpentier et al., 1998;
Cornen et al., 1999).
In summary, both elemental and isotopic chemistry
and mineralogy of the Site 1070 gabbros are consistent
with derivation from an enriched mantle source, possibly
subcontinental lithosphere, and mitigate against a depleted, NMORB-type asthenospheric source. The
pegmatites have mineralogical and chemical similarities
to many other syn-rift intrusive and extrusive rocks found
elsewhere in the region. In the continuum of lithospheric
vs asthenospheric magmas recognized at the west Iberian
margin (Charpentier et al., 1998; Cornen et al., 1999),
the Site 1070 rocks lie near the lithospheric end-member.
Origin of high Mg in gabbroic clasts and
dikelets
899
The interpretation of the origin of the high-Mg pegmatitic
dikes and clasts bears on the overall interpretation of the
pegmatitic rocks at Site 1070. If the high Mg (and Cr)
JOURNAL OF PETROLOGY
VOLUME 43
content of their mafic phases (ilmenite, augite, kaersutite)
simply reflects chemical exchange with the peridotite
host, they provide little information on pegmatite petrogenesis. If the high Mg is a primary feature, however,
these dikelets and fragments arguably represent melts
parental to the main Site 1070 gabbro. The most plausible
explanations for the high Mg#s are that they reflect
either original igneous values or high-temperature exchange with peridotite.
Because the minerals in question are essentially unzoned, any high-T diffusive exchange would have to have
gone to completion. Fe/Mg diffusion coefficients for
clinopyroxene at 900–1200°C range from 10−18 to 10−14
cm2/s (Dimanov & Sautter, 2000). Complete exchange
with centimeter-sized crystals hence requires times of the
order of 1 my at 1200°C, increasing to >10 my at 1150°C.
Evidence suggests that the pegmatites are slowly cooled:
the argon closure temperature for amphibole (>600°C)
was attained several million years after the intrusion age
inferred from zircon (Turrin, 1999). If diffusion were
aided by grain-boundary diffusion along cracks and cleavage planes, it is just possible that the pegmatites were
hot enough for long enough to effect the required Fe/
Mg exchange.
However, we argue that several other lines of reasoning
support an igneous origin for the high-Mg minerals. First,
the augite and kaersutite, although relict in rodingite,
still resemble their counterparts in the main pegmatite.
The amphibole in the dikelets, for example, retains its
characteristic red–brown color. With the exception of
elevated Mg# (and Cr content), the relict mineral chemistry of the dikelets and clasts resembles that of igneous
minerals in the main pegmatite (Figs 3 and 4; Table 2).
In our view, this unrealistically restricts high-T exchange
to only Fe, Mg, and Cr.
Second, and perhaps more telling, there is substantial
variation in Mg# and Cr amongst the various dike and
clast samples, although all are more magnesian than
the main pegmatite. If subsolidus processes (e.g. high-T
diffusive exchange with the surrounding peridotite) were
controlling Mg–Fe–Cr exchange, then Mg# and Cr
content should be more or less the same in all of the
exchanged samples. This is clearly not the case (Figs 3
and 4). Locally incomplete exchange cannot explain this
lack of uniformity, as the variation reflects the compositions of relatively homogeneous minerals and is not
an artifact of strong zoning or other intra-sample inhomogeneities. Finally, the Mg#s of coexisting amphibole
and augite are consistent: high-Mg augite coexists with
high-Mg amphibole (Figs 3 and 4).
Origin of the albitites
The coarse feldspar in the albitites appears to be of
igneous origin. Indications for this include zoning, low,
NUMBER 5
MAY 2002
but finite calcium content, and the observation that some
of the albitites have granoblastic textures (similar to those
locally developed in the gabbroid rocks) suggesting highT deformation of the albitic feldspar. This last observation
is consistent with the albitites (and the sodic feldspar)
being present before the onset of serpentinization. The
lack of albite as a plagioclase alteration product in the
pegmatites and rodingites (analcime is the sodic alteration
phase) is circumstantial evidence for an igneous origin
of the albite. The altered amphibole in one albitite,
although it is basically tremolitic hornblende, still contains
>1 wt % TiO2, probably as submicroscopic exsolved
plates of ilmenite (Table 2). We interpret the albitites
as modified differentiation products of the Site 1070
gabbroids.
Timing of gabbro pegmatite intrusion vis-àvis rift development
The Site 1070 pegmatites were emplaced at 127 Ma,
underwent non-pervasive shearing at granulite facies
conditions, and cooled through amphibole closure
(>600°C) by 119 Ma and through plagioclase closure
(>250°C) at 110 Ma (Manatschal et al., 2002). This long
deformation and cooling history shows that melting predates complete unroofing and serpentinization.
Furthermore, pegmatite intrusion post-dates syn-rift sediments that overlie serpentinite elsewhere on the Iberia
Abyssal Plain (Whitmarsh et al., 1998). Pegmatite intrusion
also post-dates the initiation of faulting at Hobby High,
100 km inboard of Site 1070. At the same time, the
gabbro pegmatites appear to be slightly younger than
the ocean crust in the western IAP (Whitmarsh & Wallace,
2001). All of this indicates that serpentinization, unroofing
and magmatism were time-transgressive during nonvolcanic (more accurately, magma-poor) rifting and that
the age of the sea floor formed by the unroofed mantle
grows younger outboard, just as is the case for normal,
volcanic ocean crust.
The Penninic nappes of eastern
Switzerland: an on-land analog
The mafic and ultramafic rocks of the Penninic nappes
in eastern Switzerland probably constitute the best documented on-land analog of the Iberian ocean–continent
transition (Froitzheim & Manatschal, 1996; Manatschal
& Bernoulli, 1999; Desmurs et al., 2002). Structural and
petrological features (e.g. ophicalcites) all suggest that
ultramafic rocks in these nappes were exhumed during
continental rifting, serpentinized, and exposed at the sea
floor. The serpentinites are intruded by a variety of
igneous rocks including magnesian gabbros, Fe–Ti-enriched gabbros, dioritic and gabbroic pegmatites, and
900
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
albitites (Desmurs et al., 2002). The gabbroic rocks constitute >10% of the outcrop of the serpentinites. Overlying the mantle section are sedimentary rocks that are
sporadically intercalated with pillowed basalt flows. Postrift, deep-water sediments overlie the entire package. The
sequence of emplacement of igneous rocks has been
interpreted as reflecting the transition from a non-volcanic margin to a slow spreading ridge.
Petrogenesis
The elemental and isotopic chemistry of the pegmatites
is consistent with a moderately enriched mantle source
(Figs 7 and 9). We argue that the composition and
mineralogical character of the local peridotite (i.e. peridotites recovered on the IAP) permits it to be the source
for the magmas that produced the pegmatites. Although
the precise nature (supra-subduction zone or subcontinental lithosphere) of the IAP peridotite is a matter
of debate, there is general agreement that it is moderately
enriched (in LREE and LILE, for example) with respect
to abyssal peridotite (Seifert & Brunotte, 1996; Abe,
2002; Hebert et al., 2002). In addition, Nd isotopic data
(Nd = 4·0; Charpentier et al., 1998) on west Iberian
mantle rocks both confirm their enriched character and
are consistent with the isotopic composition of the Site
1070 gabbro pegmatites. The presence of modal kaersutite in several mantle samples from the IAP and elsewhere at the West Iberian margin (Agrinier et al., 1988;
Cornen et al., 1996, 1999; Abe, 2002) is consistent with
its enriched character, demonstrates that the mantle is
not only enriched, but hydrous, and links the mantle
with hydrous, kaersutite-bearing igneous rocks at Site
1070 and elsewhere at the West Iberian margin.
More difficult to constrain than the nature of the
mantle source is the issue of how and where the melting
of the source occurred. We postulate that the gabbroic
pegmatites represent melts that were generated in the
upper mantle (probably in the plagioclase stability field)
then collected and fractionated close to the site where
they were generated. Our reasoning is as follows:
(1) plagioclase peridotites (some kaersutite bearing) are
widespread at the West Iberian margin (Evans & Girardeau, 1988; Cornen et al., 1996; Whitmarsh et al., 1998;
Abe, 2002). These peridotites locally preserve evidence for
partial melting (Cornen et al., 1996; Charpentier et al.,
1998), although the age of the melting event is unknown.
(2) Minerals in some Site 1070 pegmatite dikes are
nearly magnesian enough to be in equilibrium with the
mantle. These dikes may be parental to a suite of more
differentiated rocks, particularly the main pegmatite. If
so, this suggests that the local magmatic framework
involved collection and storage of melt, a necessary (albeit
not sufficient) characteristic of magma source regions.
(3) The pegmatites, including the very magnesian examples, were probably near or at vapor saturation at the
time of emplacement and, thus, not capable of sustained
upward movement because of pressure-quench constraints.
(4) The enrichment of the pegmatitic rocks in Y and
HREE suggests melting at pressures less than garnet
stability, and the sodic and aluminous nature of the
pegmatites is consistent with hydrous melting in a plagioclase-bearing, low-pressure system (e.g. Rapp, 1995).
Magmatism at Site 1070 was driven by a combination
of decompression that accompanied unroofing and passive upwelling of hot, asthenospheric mantle. The presence
of amphibole in the lithospheric mantle at the west
Iberian margin is, we argue, key to magma genesis in
the region, as amphibole peridotite in the plagioclase
stability field begins to melt (dehydration-melting) at
temperatures below 1050°C (Niida & Green, 1999). Our
model for how this melting initiated is as follows:
(1) a metasomatic event of unknown age, but probably
related to Paleozoic rifting (Cornen et al., 1996) or subduction (Abe, 2002) results in the enrichment and, most
importantly, partial hydration of the Iberian lithosphere
with concomitant formation of plagioclase and kaersutitic
amphibole.
(2) A combination of unroofing owing to crustal thinning (Whitmarsh & Wallace, 2001) and northward migration of an upwelling asthenospheric bulge beneath
the northward-opening rift (Scharer et al., 2000) causes
the mantle to approach solidus conditions.
(3) Where kaersutite is present in the lithosphere, the
solidus temperature is relatively low and dehydrationmelting occurs at temperatures as low as 1050°C. This is
>50°C below the dry solidus of depleted asthenospheric
mantle at the same pressure (Niida & Green, 1999).
Dehydration-melting generates small volumes of enriched, hydrous melts. Plagioclase is likely to be at least
partially consumed during the dehydration-melting reactions [Niida & Green (1999); also see Vielzeuf &
Schmidt (2001) for a review of dehydration-melting reactions and P–T geometry].
(4) As rifting progresses, asthenospheric melting begins
and, eventually, progresses to the point where true
oceanic crust is generated. Where the lithosphere is
anhydrous, it is likely that melting of the much hotter
asthenosphere will dominate magmatism throughout the
early and, indeed, entire history of the rift. The observation that a hydrous lithosphere, where present, is
the source for early melts in an evolving oceanic rift
probably holds as a general case.
ACKNOWLEDGEMENTS
901
I would like to thank the scientific and ship’s crew of
ODP Leg 173 for a productive and stimulating cruise.
JOURNAL OF PETROLOGY
VOLUME 43
Thanks go to Jay Thomas for preparing the mineral
separates. The manuscript has benefited considerably
from the reviews of D. Bernoulli, O. Muntener, K.
Seifert, and J. Shervais. This work was supported by
USSAC/JOI grant F00719 to J.S.B.
REFERENCES
Abe, N. (2002). Petrochemistry of serpentinized peridotite from the
Iberia Abyssal Plain (ODP Leg 173): its character intermediate
between sub-oceanic and sub-continental upper mantle. In: Wilson,
R. C. L., Whitmarsh, R. B., Taylor, B. & Frotzheim, N. (eds) NonVolcanic Rifting of Continental Margins: A Comparison of Evidence from Land
and Sea. Geological Society, London, Special Publications 187 (in press).
Adam, J. & Green, T. H. (1994). The effects of pressure and temperature
on the partitioning of Ti, Sr and REE between amphibole, clinopyroxene and basanitic melts. Chemical Geology 117, 219–233.
Adam, J., Green, T. H. & Sie, S. H. (1993). Proton microprobe
determined partitioning of Rb, Sr, Ba, Y, Zr, Nb and Ta between
experimentally produced amphiboles and silicate melts with variable
F content. Chemical Geology 109, 29–49.
Agrinier, P., Mevel, C. & Girardeau, J. (1988). Hydrothermal alteration
of the peridotites cores at the ocean–continent boundary of the
Iberian margin: petrologic sand stable isotope evidence. In: Boillot,
G. & Winterer, E. L., et al. (eds) Proceedings of the Ocean Drilling Program,
Scientific Results, 103. College Station, TX: Ocean Drilling Program.
Bernard-Griffith, G., Gruau, G., Cornen, G., Azambre, B. & Mace, J.
(1997). Continental lithospheric contribution to alkaline magmatism:
isotopic (Nd, Sr, Pb) and geochemical (REE) evidence from Serra
de Monchique and Mount Ormonde complexes. Journal of Petrology
38, 115–132.
Bernoulli, D. & Weisseret, H. (1985). Sedimentary fabrics in Alpine
ophicalcites, south Pennine Arosa zone, Switzerland. Geology 13,
755–758.
Bindeman, I. N., Davis, A. M. & Drake, M. J. (1998). Ion microprobe
study of plagioclase–basalt partition experiments at natural concentration levels of trace elements. Geochimica et Cosmochimica Acta
62(7), 1175–1193.
Boillot, G., Winterer, E. L., Meyer, A. W., et al. (eds) (1987). Proceedings
of the Ocean Drilling Program, Initial Reports, 103. College Station, TX:
Ocean Drilling Program, 663 pp.
Brenan, J. M., Shaw, H. F., Ryerson, F. J. & Phinney, D. L. (1995).
Experimental determination of trace-element partitioning between
pargasite and a synthetic hydrous andesitic melt. Earth and Planetary
Science Letters 135, 1–11.
Charpentier, S., Kornprobst, J., Chazot, G., Cornen, G. & Boillot, G.
(1998). Interaction entre lithosphere et asthenosphere au cours de
l’ouverture océanique: données isotopiques preliminaires sur the
marge passive de Galice (Atlantique-Nord). Comptes Rendus de l’Académie des Sciences 326, 757–762.
Cornen, G., Beslier, M.-O. & Girardeau, J. (1996). Petrologic characteristics of the ultramafic rocks from the ocean/continent transition
in the Iberia Abyssal Plain. In: Whitmarsh, R. B., Sawyer, D. S.,
Klaus, A. & Masson, D. G. (eds) Proceedings of the Ocean Drilling Program,
Scientific Results, 149. College Station, TX: Ocean Drilling Program,
pp. 377–395.
Cornen, G., Girardeau, J. & Monnier, C. (1999). Basalts, underplated
gabbros, and pyroxenites record the rifting process of the West
Iberian margin. Mineralogy and Petrology 67, 111–142.
Desmurs, L., Manatschal, G. & Bernoulli, D. (2002). The Steinmann
Trinity revisited: mantle exhumation and magmatism along an
NUMBER 5
MAY 2002
ocean–continent transition: the Platta nappe, eastern Switzerland.
Journal of the Geological Society, London (in press).
Dimanov, A. & Sautter, V. (2000). ‘Average’ interdiffusion of (Fe,Mn)Mg in natural diopside. European Journal of Mineralogy 12, 749–760.
Drake, M. J. & Weill, D. F. (1975). The partition of Sr, Ba, Ca, Y,
Eu2+, Eu3+ and other REE between plagioclase feldspar and magmatic silicate liquid. Geochimica et Cosmochimica Acta 39, 689–712.
Dunn, T. (1987). Partitioning of Hf, Lu, Ti, and Mn between olivine,
clinopyroxene and basaltic liquid. Contributions to Mineralogy and Petrology 96, 476–484.
Dunn, T. & Sen, C. (1994). Mineral/matrix partition coefficients
for orthopyroxene, plagioclase, and olivine in basaltic to andesitic
systems: a combined analytical and experimental study. Geochimica
et Cosmochimica Acta 58(2), 717–733.
Evans, C. A. & Girardeau, J. (1988). Galicia margin peridotites:
undepleted abyssal peridotites from the North Atlantic. In: Boillot,
G., Winterer, E. L., et al. (eds) Proceedings of the Ocean Drilling Program,
Scientific Results, 103. College Station, TX: Ocean Drilling Program,
pp. 195–207.
Feraud, G., Beslier, M.-O. & Cornen, G. (1996). 40Ar/39Ar dating of
gabbros from the ocean/continent transition of the West Iberia
Margin: preliminary results. In: Whitmarsh, R. B., Sawyer, D. S.,
Klaus, A. & Masson, D. G. (eds) Proceedings of the Ocean Drilling Program,
Scientific Results, 149. College Station, TX: Ocean Drilling Program,
pp. 519–527.
Forsythe, L. M., Nielsen, R. L. & Fisk, M. R. (1994). High fieldstrength element partitioning between pyroxene and basaltic to
dacitic magmas. Chemical Geology 117, 107–125.
Froitzheim, N. & Manatschal, G. (1996). Kinematics of Jurassic rifting,
mantle exhumation, and passive-margin formation in the Austroalpine and Penninic nappes (eastern Switzerland). Geological Society
of America Bulletin 108, 1120–1133.
Fujimaki, H. & Tatsumoto, M. (1984). Partition coefficients of Hf, Zr
and REE between phenocrysts and groundmass. Proceedings of the
14th Lunar and Planetary Science Conference, Part 2. Journal of Geophysical
Research 89(Supplement), B662–B672.
Green, T. H. & Pearson, N. J. (1985). Experimental determination of
REE partition coefficients between amphibole and basaltic to andesitic liquids at high pressure. Geochimica et Cosmochimica Acta 49,
1465–1468.
Green, T. H. & Pearson, N. J. (1987). An experimental study of Nb
and Ta partitioning between Ti-rich minerals and silicate liquids at
high pressure and temperature. Geochimica et Cosmochimica Acta 51,
55–62.
Green, T. H., Adam, J. & Sie, S. H. (1993). Proton microprobe
determined trace element partition coefficients between pargasite,
augite and silicate or carbonatitic melts (abstract). EOS Transactions,
American Geophysical Union 74, 340.
Green, T. H., Sie, S. H., Ryan, C. G. & Cousens, D. R. (1989). Proton
microprobe-determined partitioning of Nb, Ta, Zr, Sr and Y between
garnet, clinopyroxene and basaltic magma at high pressure and
temperature. Chemical Geology 74, 201–216.
Hart, S. R. & Dunn, T. (1993). Experimental cpx/melt partitioning of
24 trace elements. Contributions to Mineralogy and Petrology 113, 1–8.
Hauri, E. H., Wagner, T. P. & Grove, T. L. (1994). Experimental and
natural partitioning of Th, U, Pb and other trace elements between
garnet, clinopyroxene and basaltic melts. Chemical Geology 117, 149–
166.
Hebert, R., Gueddari, K., Lafleche, M. R., Beslier, M.-O. & Gardien,
V. (2002). Petrology and geochemistry of exhumed peridotites and
gabbros at non-volcanic margins: ODP Leg 173, West Iberian
ocean–continent zone. In: Wilson, R. C. L., Whitmarsh, R. B.,
Taylor, B. & Frotzheim, N. (eds) Non-Volcanic Rifting of Continental
902
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Margins: A Comparison of Evidence from Land and Sea. Geological Society,
London, Special Publications 187 (in press).
Jenner, G. A., Foley, S. F., Jackson, S. E., Green, T. H., Fryer, B. J.
& Longerich, H. P. (1994). Determination of partition coefficients
for trace elements in high pressure–temperature experimental run
products by laser ablation microprobe-inductively coupled plasmamass spectrometry (LAM-ICP-MS). Geochimica et Cosmochimica Acta
58, 5099–5103.
Johnson, K. T. M. (1994). Experimental cpx/ and garnet/melt partitioning of REE and other trace elements at high pressures: petrogenetic implications. Mineralogical Magazine 58A, 454–455.
Johnson, K. T. M. & Kinzler, R. J. (1989). Partitioning of REE, Ti,
Zr, Hf, and Nb between clinopyroxene and basaltic liquid: an ion
microprobe study (abstract). EOS Transactions, American Geophysical
Union 70, 1388.
Krawczyk, C. M., Reston, T. J., Beslier, M.-O. & Boillot, G. (1996).
Evidence for detachment tectonics on the Iberia Abyssal Plain rifted
margin. In: Whitmarsh, R. B., Sawyer, D. S., Klaus, A. & Masson,
D. G. (eds) Proceedings of the Ocean Drilling Program, Scientific Results, 149.
College Station, TX: Ocean Drilling Program, pp. 603–615.
LaTourrette, T. Z. & Burnett, D. S. (1992). Experimental determination
of U and Th partitioning between clinopyroxene and natural and
synthetic basaltic liquid. Earth and Planetary Science Letters 110, 227–244.
Lundstrom, C. C., Shaw, H. F., Ryerson, F. J., Phinney, D. L.,
Gill, J. B. & Williams, Q . (1994). Compositional controls on the
partitioning of U, Th, Ba, Pb, Sr, and Zr between clinopyroxene
and haplobasaltic melts: implications for uranium series disequilibria
in basalts. Earth and Planetary Science Letters 128, 407–423.
Masson, D. G. & Miles, P. R. (1984). Mesozoic seafloor spreading
between Iberia, Europe, and North America. Marine Geology 56,
279–287.
Manatschal, G. & Bernoulli, D. (1999). Architecture and tectonic
evolution of non-volcanic margins: present-day Galicia and ancient
Adria. Tectonics 18, 1099–1119.
Manatschal, G., Frotzheim, N., Rubenach, M. J. & Turrin, B. (2002).
The role of detachment faulting in the formation of an ocean–
continent transition: insights from the Iberia Abyssal Plain. In:
Wilson, R. C. L., Whitmarsh, R. B., Taylor, B. & Frotzheim, N.
(eds) Non-Volcanic Rifting of Continental Margins: A Comparison of Evidence
from Land and Sea. Geological Society, London, Special Publications 187 (in
press).
McCallum, I. S. & Charette, M. P. (1978). Zr and Nb partition
coefficients: implications for the genesis of mare basalts, KREEP,
and sea floor basalts. Geochimica et Cosmochimica Acta 42, 859–869.
McKay, G., Le, L., Wagstaff, J. & Crozaz, G. (1994). Experimental
partitioning of rare earth elements and strontium: constraints on
petrogenesis and redox conditions during crystallization of Antarctic
angrite Lewis Cliff 86010. Geochimica et Cosmochimica Acta 58, 2911–
2919.
McKay, G. A., Wagstaff, J. & Yang, S.-R. (1986). Zirconium, hafnium
and rare earth element partition coefficients for ilmenite and other
minerals in high-Ti lunar mare basalts. Proceedings of the 16th Lunar and
Planetary Science Conference. Journal of Geophysical Research 91, 293–302.
Nielsen, R. L. (2000). Partition coefficient database. On line at http://
earthref.org/GERM/data/partitioning.ref.
Nielsen, R. L., Gallahan, W. E. & Newberger, F. (1992). Experimentally
determined mineral–melt partition coefficients for Sc, Y and REE
for olivine, orthopyroxene, pigeonite, magnetite and ilmenite. Contributions to Mineralogy and Petrology 110, 488–499.
Niida, K. & Green, D. H. (1999). Stability of pargasitic amphibole in
MORB pyrolite under upper mantle conditions. Contributions to
Mineralogy and Petrology 135, 18–40.
Paster, T. P., Schauwecher, D. S. & Haskin, L. A. (1974). The behavior
of some trace elements during solidification of the Skaergaard layered
series. Geochimica et Cosmochimica Acta 38, 1549–1577.
Pickup, S. L. B., Whitmarsh, R. B., Fowler, C. M. R. & Reston, T. J.
(1996). Insight into the nature of the ocean–continent transition off
West Iberia from a deep multichannel seismic reflection profile.
Geology 24, 1079–1082.
Rabinowitz, P. D., Cande, S. C. & Hayes, D. E. (1979). The J-anomaly
in the central North Atlantic Ocean. In: Initial Reports of the Deep Sea
Drilling Project, 43. Washington, DC: US Government Printing Office,
pp. 879–885.
Rapp, R. P. (1995). Amphibole-out phase boundary in partially melted
metabasalt, its control over liquid fraction, composition, and source
permeability. Journal of Geophysical Research 100, 15601–15510.
Ringwood, A. E. (1970). Petrogenesis of Apollo 11 basalts and implications for lunar origin. Journal of Geophysical Research 75, 6453–
6479.
Rollinson, H. (1993). Using Geochemical Data: Evaluation, Presentation,
Interpretation. Harlow, UK: Longman, 352 pp.
Rubenach, M. J. (1999). Microstructures, metamorphism, and tectonic
significance of metabasic rocks, ocean–continent transition, Iberia
Abyssal Plain. Australian Geological Society, Abstract Series 53, 231–232.
Scharer, U., Girardeau, J., Cornen, G. & Boillet, G. (2000). 138–121 Ma
asthenospheric magmatism prior to continental break-up in the
North Atlantic and geodynamic implications. Earth and Planetary
Science Letters 181, 555–572.
Schumacher, M., Chambost, E., McKeegan, K. D., Harrison, T. M.
& Migeon, H. (1994). In situ dating of zircon with the Cameca ims
1270. In: Benninghoven, A., et al. (eds) Secondary Ion Mass Spectrometry
SIMS IX. New York: John Wiley, pp. 919–922.
Seifert, K. E. & Brunotte, D. (1996). Geochemistry of serpentinized
mantle peridotite from site 897 in the Iberia Abyssal Plain. In:
Whitmarsh, R. B., Sawyer, D. S., Klaus, A. & Masson, D. G. (eds)
Proceedings of the Ocean Drilling Program, Scientific Results, 149. College
Station, TX: Ocean Drilling Program, pp. 413–424.
Seifert, K. E., Chang, C.-W. & Brunotte, D. A. (1997). Evidence from
Ocean Drilling Program Leg 149 mafic igneous rocks for ocean crust
in the Iberia Abyssal Plain ocean–continent transition. Journal of
Geophysical Research 102, 7915–7928.
Shirey, S. B., Bender, J. F. & Langmuir, C. H. (1987). Three-component
isotopic heterogeneity near the Oceanographer transform, MidAtlantic Ridge. Nature 325, 217–223.
Sisson, T. W. (1994). Hornblende–melt trace-element partitioning
measured by ion microprobe. Chemical Geology 117, 331–344.
Skelton, A. D. L. & Valley, J. W. (2000). The relative timing of
serpentinisation and mantle exhumation at the ocean–continent
transition, Iberia: constraints from oxygen isotopes. Earth and Planetary
Science Letters 178, 327–338.
Skulski, T., Minarik, W. & Watson, E. B. (1994). High-pressure
experimental trace-element partitioning between clinopyroxene and
basaltic melts. Chemical Geology 117, 127–147.
Staudigel, H., Zindler, A., Hart, S. R., Leslie, T., Chen, C.-Y. &
Clague, D. (1984). The isotope systematics of a juvenile intraplate
volcano: Pb, Nd, and Sr isotope ratios of basalts from Loihi Seamount, Hawaii. Earth and Planetary Science Letters 69, 13–29.
Sun, S.-s. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and
processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the
Ocean Basins. Geological Society, London, Special Publications 42, 313–345.
Tucholke, B. E. & Ludwig, W. J. (1982). Structure and origin of the
J-anomaly ridge, western North Atlantic. Journal of Geophysical Research
87, 9389–9401.
903
JOURNAL OF PETROLOGY
VOLUME 43
Turner, S., Hawkesworth, C., Rogers, N. & King, P. (1997). U–Th
disequilibria and ocean island basalt generation in the Azores.
Chemical Geology 139, 145–164.
Turrin, B. (1999). 40Ar/39Ar ages from the ocean–continent transition
zone, Iberia Abyssal Plain, ODP leg 173: chronologic and thermochronologic constraints on the kinematics of lithospheric extension
and continental breakup (abstract). In: Wilson, R. C. L., Whitmarsh,
R. B., Taylor, B. & Frotzheim, N. (eds) Non-Volcanic Rifting of Continental
Margins: A Comparison of Evidence from Land and Sea. Geological Society,
London, Special Publications, 187, (in press).
Veizer, J., Ala, D., Azmy, K., Bruckschen, P., Buhl, D., Bruhn, F.,
Carden, G. A. F., Diener, A., Ebneth, S., Godderis, Y., Jasper, T.,
Korte, C., Pawellek, F., Podlaha, O. G. & Strauss, H. (1999). 87Sr/
86
Sr and 18O evolution of Phanerozoic seawater. Chemical Geology
161, 59–88.
Vielzeuf, D. & Schmidt, M. W. (2001). Melting relations in hydrous
systems revisited: application to metapelites, metagreywackes and
metabasalts. Contributions to Mineralogy and Petrology 141, 251–267.
Vogt, P. R. & Tucholke, B. E. (1989). North Atlantic Ocean basin;
aspects of geologic structure and evolution. In: Bally, A. W. &
Palmer, A. R. (eds) The Geology of North America—An Overview. The
Geology of North America, Volume A. Boulder, CO: Geological Society
of America, pp. 53–80.
Whitmarsh, R. B. & Miles, P. R. (1995). Models of the development
of the West Iberia rifted continental margin at 40°30′N deduced
from surface and deep-tow magnetic anomalies. Journal of Geophysical
Research 100, 3789–3806.
Whitmarsh, R. B. & Sawyer, D. S. (1996). The ocean–continent
transition beneath the Iberia Abyssal Plain and continental rifting
to seafloor-spreading processes. In: Whitmarsh, R. B., Sawyer, D.
S., Klaus, A. & Masson, D. G. (eds) Proceedings of the Ocean Drilling
Program, Scientific Results, 149. College Station, TX: Ocean Drilling
Program, pp. 713–733.
Whitmarsh, R. B. & Wallace, P. J. (2001). The rift-to-drift development
of the west Iberia non-volcanic continental margin: a summary and
review of the contribution of Ocean Drilling Program Leg 173. In:
Beslier, M.-O., et al. (eds) Proceedings of the Ocean Drilling Program,
Scientific Results, 173. College Station, TX: Ocean Drilling Program,
pp. 1–36. Available online at //www-odp.tamu.edu/publications/
173SR.
Whitmarsh, R. B., Beslier, M.-O., Wallace, P.J., et al. (eds) (1998).
Proceedings of the Ocean Drilling Program, Initial Reports, 173. College
Station, TX: Ocean Drilling Program, 493 pp.
Whitmarsh, R. B., Dean, S. M. & Minshull, T. A. (2000). Tectonic
implications of exposure of lower continental crust beneath the
Iberia Abyssal Plain, northeast Atlantic Ocean: geophysical evidence.
Tectonics 19, 919–942.
Wilson, M. (1989). Igneous Petrogenesis. New York: Chapman and Hall,
466 pp.
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APPENDIX: K d VALUES AND DATA
SOURCES USED TO FORMULATE
INVERSE COMPOSITIONAL MODELS
The Kd values used to bracket the inverse compositional
models (Table 6; Figs 8 and 9) are given in Table A1.
Kd values are taken from the compilation of Rollinson
(1993) (reference 1 in Table 1A) and from experimental
and ion probe phenocryst–matrix studies on mafic rocks
as compiled by Nielsen (2000) in the GERM database.
Because of inherent large errors, very low Kd values (e.g.
0·01) for phases with very low concentrations of a given
element (e.g. Zr in plagioclase) were excluded from the
models. Only amphibole and clinopyroxene Kd values
were considered for REE because (1) there were virtually
no data for ilmenite and (2) Kd values and concentrations
for HREE in plagioclase were very low. In every case
(except Cs), Kd values for at least two mineral species
were considered, although if the high and low model
were based on the same mineral, this may not be evident
from Table A1. Sources for Table A1 are given in the
footnote to the table. Other studies that were considered
(and yielded intermediate values) are: Ringwood (1970);
McCallum & Charette (1978); McKay et al. (1986); Dunn
(1987); Green et al. (1989); Johnson & Kinzler (1989);
LaTourrette & Burnett (1992); Nielsen et al. (1992); Adam
et al. (1993); Hart & Dunn (1993); Forsythe et al. (1994);
Jenner et al. (1994); Johnson (1994); Lundstrom et al.
(1994); McKay et al. (1994); Bindeman et al. (1998).
BEARD et al.
GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN
Table A1: Partition coefficients and data sources for inverse compositional models
High value partition coefficient
Element
Mineral
Partition
Low value partition coefficient
Source
Mineral
coefficient
Partition
Source
coefficient
Sr
cpx
0·05
8
amph
0·08
K
amph
0·96
1
plag
0·17
3
1
Rb
amph
0·58
3
plag
0·05
12
Ba
plag
1·5
12
plag
0·16
13
Cs
plag
0·07
12
plag
0·07
12
Th
plag
0·19
12
amph
0·02
7
Ta
ilm
2·7
2
amph
0·08
3
Nb
ilm
2·3
2
amph
0·08
3
Ce
amph
0·84
1
cpx
0·04
8,9
Zr
amph
1·56
1
cpx
0·06
8
Hf
amph
1·53
1
cpx
0·12
9
Sm
amph
1·8
1
amph
0·37
4
Ti
amph
3·1
5
amph
0·87
4
Y
cpx
0·9
1
amph
0·4
3
Yb
amph
1·64
1
cpx
0·32
9
Sc
amph
4·2
1
cpx
0·81
10
Cr
cpx
34
1
cpx
1·7
10
Ni
cpx
14
1
cpx
1·5
1
La
amph
0·54
1
cpx
0·03
8
Nd
amph
1·34
1
cpx
0·12
8,9
Eu
amph
1·56
1
amph
0·31
6
Gd
amph
2·01
1
amph
0·63
1
Tb
amph
1·4
1
cpx
0·42
11
Dy
amph
2·02
1
cpx
0·30
8,9
Ho
amph
3
6
amph
0·44
4
Er
amph
1·74
1
amph
0·55
1
Lu
amph
1·57
1
amph
0·32
4
Sources: 1, Rollinson (1993); 2, Green & Pearson (1987); 3, Green et al. (1993); 4, Adam & Green (1994); 5, Sisson (1994); 6,
Green & Pearson (1985); 7, Brenan et al. (1995); 8, Skulski et al. (1994); 9, Fujimaki & Tatsumoto (1984); 10, Hauri et al.
(1994); 11, Paster et al. (1974); 12, Dunn & Sen (1994); 13, Drake & Weill (1975).
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