JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
PAGES 3–38
2003
Alkali Picrites Formed by Melting of Old
Metasomatized Lithospheric Mantle:
Manı̂tdlat Member, Vaigat Formation,
Palaeocene of West Greenland
LOTTE M. LARSEN1,5∗, ASGER K. PEDERSEN2,5, BJØRN SUNDVOLL3
AND ROBERT FREI4,5
1
GEOLOGICAL SURVEY OF DENMARK AND GREENLAND, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
2
GEOLOGICAL MUSEUM, ØSTER VOLDGADE 5–7, DK-1350 COPENHAGEN K, DENMARK
3
MINERALOGISK–GEOLOGISK MUSEUM, SARS GATE 1, N-0562 OSLO, NORWAY
4
GEOLOGICAL INSTITUTE, UNIVERSITY OF COPENHAGEN, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K,
DENMARK
5
DANISH LITHOSPHERE CENTRE, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
RECEIVED DECEMBER 5, 2001; REVISED TYPESCRIPT ACCEPTED JULY 1, 2002
are required in the melting assemblage and dominate the Pb
isotope compositions. The various magma types and the partly
complementary relation between them suggest that the lithospheric
mantle had an ordered structure, possibly with old metasomatic
zones formed by successive trapping of elements in migrating fluids.
Alkaline picrites and basalts constitute 20–200 m of lava flows
and hyaloclastites in the middle part of an >2 km thick succession
of tholeiitic picrites and basalts formed during continental rifting of
West Greenland around 60 Ma. The alkaline rocks, found only
in northern Disko, have phenocrysts of olivine + chromite ±
clinopyroxene; lava flows contain abundant groundmass clinopyroxene and plagioclase, whereas pillow breccias contain abundant
fresh alkali basaltic glass. Six compositional types are present; all
are strongly but variably enriched in incompatible trace elements
[Ba, U, Nb, Ta, light rare earth elements (LREE)], yet their
major elements, with relatively high SiO2 and Al2O3 and low
Na2O, do not suggest an origin by small degrees of mantle melting.
The isotope compositions are unusual, with negative Nd and mostly
negative Sr (below the mantle array), high 206Pb/204Pb (below
the Northern Hemisphere Reference Line), and mostly negative Os.
The most likely source for the alkaline magmas is old metasomatized
lithospheric mantle in which melting was induced by the passing
hot, asthenosphere-derived, tholeiitic magmas. Simple mass-balance
calculations suggest that the melting assemblages consisted of >60%
pargasitic amphibole, 26–30% clinopyroxene, >9% olivine and
>1% apatite. Mica in the source is required for only the least
enriched magma type. For the most enriched magmas small amounts
of Ba–U–Nb–Sr–LREE-rich oxides (lindsleyite and hawthorneite)
Alkaline extrusive rocks are a volumetrically insignificant
component in most large igneous provinces (LIPs). LIPs
are overwhelmingly tholeiitic in character and are considered to be formed by relatively high degrees of melting,
mainly of the asthenosphere, and usually in the presence
of a mantle plume (e.g. Saunders et al., 1997). The alkaline
magmas are generally considered to be products of lower
degrees of melting; however, their mantle sources,
whether asthenospheric, lithospheric, or both, are con-
∗Corresponding author. Telephone: +45 38142252.
E-mail: [email protected]
Published by Oxford University Press
alkali picrite; amphibole melting; Greenland; lithosphere
melting; metasomatism
KEY WORDS:
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
andesites (Pedersen, 1985a; Pedersen et al., 1996;
Lightfoot et al., 1997). The volcanism of the Vaigat
Formation occurred in three main cycles which formed
the three main stratigraphic members; from older to
younger these are the Anaanaa, Naujánguit and Ordlingassoq Members. The interval with alkaline rocks is
the upper Naujánguit Member to lower Ordlingassoq
Member, and the main part of the alkaline rocks forms
one stratigraphic unit, which is formalized as the
Manı̂tdlat Member (Pedersen, 1985b). The geographical
distribution of these rocks is limited to northern Disko
(Fig. 1).
tentious. Alkaline extrusives have often been emplaced
late in the formation of the LIPs, e.g. the meimechites
capping the Siberian basalt plateau (Arndt et al., 1995,
1998), and the nephelinites and basanites capping the
East Greenland basalt plateau and nunataks (Brooks et
al., 1979; Brown et al., 1996; Bernstein et al., 2000). Both
the Siberian and the East Greenland alkaline extrusives
have been interpreted to be of asthenospheric origin
(Brooks et al., 1979; Arndt et al., 1995, 1998; Brown et
al., 1996; Bernstein et al., 2000) whereas in other LIPs,
such as the Deccan and Yemen, lithospheric components
are thought to be involved (e.g. Mahoney et al., 1985;
Baker et al., 1997). Probably, the alkaline rocks in LIPs
are polygenetic.
During Palaeogene rifting and continental break-up in
the North Atlantic, large volumes of flood basalts were
extruded on the continental margins of both West and
East Greenland. The magmas are thought to be generated
mainly by melting within the impacting Iceland mantle
plume (e.g. review by Saunders et al., 1997). In contrast
to East Greenland, the West Greenland flood basalts do
not include any nephelinites but terminate with transitional to mildly alkaline basalts (Clarke & Pedersen,
1976; Larsen, 1977). The voluminous main succession is
uniformly tholeiitic except within one limited interval in
the middle part, which, in addition to tholeiitic picrites,
contains three close-lying levels with alkali picrites and
alkali basalts with distinctive and highly unusual geochemical and isotopic characteristics. This paper explores
the petrogenesis of the alkaline melts and the nature of
their unusual mantle sources.
Alkaline volcanic rock units
Alkaline rocks occur at three close-lying stratigraphic
levels (Fig. 1). They are divided into a number of types
as shown in Table 1. The total estimated volume of the
alkaline rocks is around 30 km3; with possible extensions
to the west and north the original volume may have
been up to 50 km3, about 0·05% of the original volume
of onshore basalts.
The oldest alkaline unit is a volcanic neck with a few
associated lava flows of alkali basalt in the Stordal area
(Fig. 1). These rocks (Stordal type) are of very limited
extent and volume and are interbedded within picritic
lava flows of the uppermost Naujánguit Member.
The second alkaline unit is an up to 20 m thick series
of olivine-rich alkali picrite flows within the lowermost
part of the Ordlingassoq Member. This unit (Type 0)
can be followed over >20 km along the north coast of
Disko and was most probably produced from one volcanic
centre. It is overlain by 50–60 m of tholeiitic picrites,
which are in turn overlain by the rocks of the Manı̂tdlat
Member.
The third alkaline unit is the Manı̂tdlat Member, which
forms a purplish brown marker horizon within the grey
tholeiitic picrites of the Ordlingassoq Member (Fig. 2).
The Manı̂tdlat Member is found within a 30 km by 20
km area in northern Disko (Fig. 1). It is on average about
50 m thick and represents eruptions from several volcanic
centres, most probably fissure eruptions because all the
lava flows are of pahoehoe type and no traces of volcanic
edifices or explosive activity have been found. No eruption
sites are known but they must be local. The volcanic
rocks were produced from at least four alkaline centres
and two tholeiitic centres, which interfinger laterally. The
alkaline rocks are picrites and alkali basalts with four
element enrichment patterns (Types 1a, 1b, 2 and 3; see
Table 1). The youngest rocks are clinopyroxene-phyric
basalts (ankaramites, Type 3), probably erupted from a
volcanic centre in the southern part of the area. A dyke
cutting the whole Manı̂tdlat Member is tholeiitic but
somewhat enriched. Overlying the Manı̂tdlat Member
are tholeiitic picrites of the Ordlingassoq Member.
GEOLOGICAL SETTING
The West Greenland basalts form part of the Nuussuaq
Basin, a fault-controlled extensional basin at the continental margin where Precambrian basement is covered
by Cretaceous–Palaeocene sediments overlain by volcanic rocks (Fig. 1; Chalmers et al., 1999). The major
part of the onshore volcanic succession was erupted
during a short time period around 60 Ma (Storey et al.,
1998). The lower part of the 2–4 km thick volcanic
succession is dominated by highly magnesian picritic
rocks (the Vaigat Formation), whereas the upper part
consists of more evolved, plagioclase-phyric basalts of
the Maligât Formation (Clarke & Pedersen, 1976). The
magnesian magmas of the Vaigat Formation were generated at high temperatures and very high production
rates in the asthenosphere and passed swiftly through
the lithosphere (Gill et al., 1992; Holm et al., 1993; Larsen
& Pedersen, 2000; Pedersen et al., 2002). Most magmas
escaped contamination, although a number of discrete
crustal contamination episodes led to the formation of
subordinate units of siliceous basalts and magnesian
4
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 1. Location and extent of the alkaline rocks of the Manı̂tdlat Member and associated units (Type 0 and Stordal centre) within the Vaigat
Formation, West Greenland. The lithological log shows only the relevant middle part (c. one-fourth) of the total volcanic stratigraphy. Lithologies
distinguished on the log are thin subaerial lava flows, subaqueous cross-bedded hyaloclastite breccias (hy), and flows from the Stordal volcanic
centre. The log was measured on the north coast of Disko >5 km west of the Maniillat (old spelling: Manı̂tdlat) gully. The lateral variations of
thicknesses and subaerial or subaqueous facies distributions are considerable (Fig. 2; Pedersen, 1985b). Arrow points to the location of Fig. 2.
2). Consequently, all the magma types discussed here
except the Stordal type exist in hyaloclastite facies with
extremely fresh glassy rocks; in contrast, many of
the subaerial lava flows are affected by zeolite-facies
metamorphism. Stordal type glass is found in a chilled
Most of the volcanic rocks were erupted subaerially.
However, the volcanic front was prograding laterally
towards the east and SE into a volcanic-dammed lake
(Pedersen et al., 1998), so that the thin lava flows pass
laterally into thick hyaloclastite breccia deposits (Fig.
5
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
Table 1: Divisions of the alkaline rocks of the Manı̂tdlat Member and associated units
Stratigraphy
Name
Vol. %
Dyke
Dyke
MMb
Type 3
15
—
MgO% range
Phenocrysts
17–18
ol + cr
Tholeiitic but increased Ba, Nb, REE, Sr, P
cpx + ol + cr
High Ba, Nb, REE, Sr, P
6–11
Distinction
MMb
Type 2
25
10–16
ol + cr ± cpx
High Rb, K, Zr, Ti
MMb
Type 1b
5
12–14
ol + cr ± cpx
As 1a but higher Ba, REE, Sr
ol + cr ± cpx
‘Main’ type; high Ba, U, Nb, REE, Sr, P
MMb
Type 1a
40
8–21
Below MMb
Type 0
10
21–23
Stordal centre
Stordal
5
8
ol + cr
Aphyric
MMb, Manı̂tdlat Member; ol, olivine; cr, chromite; cpx, clinopyroxene. The Stordal type and Type 0 are primarily defined
by their occurrence below the Manı̂tdlat Member, but they are also chemically distinctive, as apparent from the later Figs
6 and 9. The volume relations between the various types are crude estimates.
Fig. 2. Southwest-facing wall of the Kuugannguaq valley, Disko (see Fig. 1). Some boundaries are outlined in black. Vaigat Formation from below:
N, Naujánguit Member, subaerial lava flows, partly sediment contaminated (thick rusty flows); MM2 and MM3, Manı̂tdlat Member Types 2 and
3, brown subaerial lava flows transforming laterally towards the SE into brownish and bluish foreset-bedded hyaloclastite breccias; Oh, Ordlingassoq
Member, hyaloclastite breccias; Ol, Ordlingassoq Member, thin grey subaerial picrite lava flows. MF, Maligât Formation, thick brownish flows of
plagioclase-phyric basalts. Vertical north-trending dykes cut the wall obliquely. Height of section in photograph is 1000 m.
ANALYTICAL TECHNIQUES
neck contact. Mantle xenoliths have not been found
despite dedicated search.
In this paper, the alkaline units will be collectively
referred to as the Manı̂tdlat Member.
Microprobe analyses of phenocrysts, glass inclusions and
matrix glasses were made on a JEOL Superprobe at
the University of Copenhagen, using a combination
6
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
of wavelength-dispersive (WDS) and energy-dispersive
(EDS) detection systems. Normal operating conditions
were 15 kV acceleration voltage, 15 nA beam current,
20 s total counting time for WDS and 60 s live time for
EDS analyses. High-precision analyses of Ca, Cr, Ni and
Ti in olivine and Ni, Ti and V in chromite were made
using 100 nA beam current and 40 s total counting time
(WDS); the lower limits of detection for these conditions
are 15–30 ppm (Pedersen, 1985a). Glasses were analysed
with 15 nA beam current and extended total counting
times (WDS) for Ti (60 s), K (60 s) and P (120 s), yielding
reproducibilities of 0·03 wt % TiO2 and 0·02 wt % K2O
and P2O5 (2 on 11 repeat analyses on glass standards).
Sixty whole-rock samples from the alkaline units and
15 samples of the contemporaneous tholeiitic picrites
were analysed for major and trace elements by X-ray
fluorescence spectrometry (XRF). Major elements were
determined on fused glass discs at the Geological Survey
of Denmark and Greenland, as described by Kystol &
Larsen (1999). FeO was determined by titration. Trace
elements were determined on pressed powder pellets at
the Geological Institute, University of Copenhagen, using
a Philips PW 1400 spectrometer and standard analytical
methods with USGS reference materials for calibration.
A subset of 13 alkaline and four tholeiitic samples,
mainly fresh pillow breccias, was analysed for trace
elements by inductively coupled plasma mass spectrometry (ICP-MS) and for Sr, Nd and Pb isotopes. The
ICP-MS analyses were performed on a Perkin–Elmer
SCIEX Elan 6000 at the University of Durham, using
methods described by Turner et al. (1999). Reproducibility, based on replicate digestion of samples,
varied from 1·5% to 3% for most analyses.
Sr, Nd and Pb isotope ratios were determined on
unleached samples on a VG354 instrument at the
Mineralogisk–Geologisk Museum, Oslo, using methods
described by Griffin et al. (1988). Average values for
repeated standard analyses during the analytical period
were 87Sr/86Sr = 0·71023 ± 3 (2 SE) for NBS987 and
143
Nd/144Nd = 0·511112 ± 5 (2 SE) for the J&M
standard batch no. S819093A. The Pb standard NBS981
gave 206Pb/204Pb = 16·897 ± 0·005, 207Pb/204Pb =
15·434 ± 0·005 and 208Pb/204Pb = 36·540 ± 0·015 (2
SE).
Seven samples were analysed for Os isotopes at Geocentre Copenhagen. Samples were spiked with an 188Osand 187Re-enriched solution and digested in inversed
(14N HNO3:10N HCl = 3:1) aqua regia in carius tubes
at 230°C for 1 week (Shirey & Walker, 1995). Os was
distilled from aqua regia directly into 8N HBr (Nägler
& Frei, 1997) and purified following Roy-Barman &
Allègre (1994). Os isotope analyses were performed on
a VG Sector 54 solid-source negative thermal ionization
mass spectrometer, using a multi-collector static routine
and single multiplier peak jump mode for small Os
beams. Re was purified using the liquid extraction method
of Cohen & Waters (1996) and the concentrations were
measured by multiple-collector (MC)-ICP-MS on an
Axiom instrument, using Ir-doped sample solutions for
controlling mass fractionation of Re through monitoring
the 190Ir/194Ir ratio. Procedural blanks for Re were <30
pg and for Os <3 pg.
Mössbauer analyses were performed on handpicked
glass chips at the Royal Veterinary and Agricultural
University, Copenhagen. The spectra were obtained at
295 and 80 kV, using a constant acceleration spectrometer, and were fitted using three Fe2+ doublets and
one Fe3+ doublet.
Analytical results are presented in Tables 2–6. The
complete dataset is available for downloading from the
Journal of Petrology web site at http://www.petrology.oupjournals.org (Electronic Appendix A).
PETROGRAPHY AND MINERALOGY
The alkaline rocks have simple mineralogies with olivine
and chromite as the earliest phenocryst phases, just as
in the contemporaneous tholeiites (Larsen & Pedersen,
2000). In contrast to the tholeiites, clinopyroxene was
the next phase to crystallize, followed by plagioclase. The
glassy rocks are 1–3 cm thick pillow rims with phenocrysts
(up to 1–2 mm) and microphenocrysts embedded in
70–80 vol. % clear, pristine, pale yellow glass, which has
yielded well-defined 39Ar–40Ar ages of 60·3 ± 1·0 Ma
and 60·7 ± 0·5 Ma (Storey et al., 1998). In the glassy
rocks the amount of clinopyroxene increases gradually
with decreasing Mg content of the rocks, from none or
just a few tiny microphenocrysts in the glasses of the most
Mg-rich rocks (Type 0), through increasingly frequent
microphenocrysts, often in clusters, in Type 1 and 2
glasses, to up to 1 mm phenocrysts in the glasses of Type
3 where clinopyroxene is the dominant phenocryst phase.
Plagioclase is absent from the glassy rocks of Type 0,
and it forms microlites in the Type 1 and 2 glasses, and
sparse microphenocrysts in the Type 3 glasses. The
Stordal rocks are aphyric to olivine–clinopyroxene microphyric.
In the crystalline lava samples the groundmass consists
of clinopyroxene, plagioclase, olivine, Fe–Ti oxides and
apatite in intersertal, intergranular or subophitic textures.
No primary mica or amphibole has been found. The
alkaline and calcic character is reflected in very high
modal proportions of clinopyroxene (17–40% normative
di), with crystals often showing hourglass zoning and
purplish colours. Late-stage segregation veins contain
purple, zoned, prismatic clinopyroxene crystals, zoned
plagioclases often heavily zeolitized, semi-skeletal magnetite and ilmenite and frequent apatite crystals, all
embedded in a matrix of fine-grained zeolite–smectite
7
7·51
0·156
FeO
MnO
8
101·29
21·24
29·43
38·92
0·11
3·58
8·95
0·12
18·17
0·32
100·29
12·17
Al2O3
Cr2O3
V 2O 3
Fe2O3
FeO
MnO
MgO
NiO
Sum
47·00
3·93
cr-no.
fe-no
92·11
78·35
mg-no.
Ol mg-no.
0·73
Fe2+/Fe
(88)
10·86
44·54
69·39
0·59
0·27
15·92
0·19
12·51
9·70
0·15
0
(87)
10·10
47·22
63·40
0·65
22·40
100·10
0·21
14·14
0·25
14·55
8·73
0·12
35·00
26·23
0·66
0·21
327099
3
88·77
99·96
0·489
0·300
47·76
0·179
10·77
0·066
0·011
40·39
0
4
0·544
0·221
45·71
0·241
13·19
0·056
0·018
40·14
4
0·74
0·15
88·74
3·95
47·10
71·76
0·79
14·81
100·90
0·26
16·61
0·21
11·64
3·52
0·17
38·56
29·04
1a
326787
86·06
100·12
0
327099
5
0·574
0·217
44·40
0·238
14·79
0·050
0·018
39·84
5
0·84
0·09
87·03
4·11
53·35
62·68
0·83
18·18
101·09
0·07
14·15
0·23
15·02
3·52
0·65
41·93
24·59
1a
326787
84·25
100·13
0
327099
6
0·88
0·10
83·48
3·98
53·43
57·63
0·85
19·98
101·14
0·08
12·90
0·27
16·90
3·42
0·61
41·64
24·34
1a
326787
6
87·21
99·76
0·444
0·283
46·59
0·191
12·18
0·060
0·010
40·00
1b
264120
7
7
87·21
11·19
38·49
67·70
0·60
22·55
101·70
0·24
15·79
0·21
13·43
10·13
0·14
29·43
31·54
0·69
0·10
1b
264120
86·08
100·06
0·442
0·222
45·74
0·231
13·18
0·046
0·005
40·19
1b
264120
8
85·91
13·38
47·06
61·58
0·60
25·73
100·91
0·25
13·88
0·23
15·44
11·44
0·21
33·17
25·02
1·16
0·11
1b
264120
8
80·56
99·44
0·579
0·161
41·41
0·294
17·81
0·041
0·026
39·12
1b
264120
9
0·10
0·24
14·94
0·16
13·74
8·78
0·17
38·72
22·64
1·38
88·91
10·34
53·42
65·96
0·63
21·64
100·87
2
264113
9
85·35
99·66
0·259
0·400
45·13
0·279
13·80
0·082
0·011
39·70
1b
264120
10
0·17
0·15
13·87
0·21
15·44
5·37
0·20
39·49
24·03
1·33
85·73
6·33
52·43
61·54
0·76
20·28
100·26
2
264113
10
84·37
99·99
0·273
0·342
44·65
0·294
14·74
0·079
0·013
39·60
1b
264120
11
0·10
1·18
84·84
11·54
49·23
58·86
0·65
25·04
99·89
0·12
13·09
0·27
16·30
9·71
0·25
34·80
24·07
3
264093
11
85·97
99·79
0·423
0·213
45·46
0·235
13·22
0·060
0·017
40·16
1b
264120
12
0·49
1·38
82·14
11·74
51·94
54·45
0·67
26·39
99·83
0·20
11·92
0·26
17·77
9·58
0·35
35·72
22·16
3
264093
12
83·48
99·96
0·513
0·183
43·74
0·277
15·42
0·037
0·024
39·77
1b
264120
13
87·93
9·78
57·35
64·72
0·65
21·18
99·95
0·27
14·19
0·17
13·78
8·22
0·18
41·51
20·70
0·84
0·09
dyke
326785
13
86·88
99·25
0·264
0·354
45·98
0·165
12·37
0·099
0·007
40·01
1b
176718
14
84·01
9·35
59·11
55·85
0·71
24·02
100·64
0·17
12·05
0·21
16·98
7·83
0·16
42·69
19·80
0·61
0·14
1b
176718
14
82·93
100·15
0·262
0·273
43·71
0·241
16·03
0·069
0·024
39·54
1b
176718
15
83·02
32·69
64·13
50·81
0·46
40·89
100·40
0·28
10·84
0·25
18·70
24·66
0·31
31·00
11·63
2·47
0·26
1b
176718
15
85·37
99·42
0·353
0·237
45·01
0·232
13·74
0·039
0·022
39·79
1b
176718
16
88·93
0·80
37·16
0·28
79·51
101·24
0·34
7·39
0·21
22·27
63·62
0·17
0·06
5·02
2·06
0·10
dyke
326785
16
84·98
99·77
0·511
0·144
44·88
0·248
14·14
0·034
0·007
39·81
3
264093
17
57·05
1·39
50·86
0·36
57·98
101·40
0·54
12·14
0·23
20·90
41·20
0·35
0·41
19·55
5·90
0·18
dyke
326785
17
79·65
99·59
0·628
0·114
40·85
0·333
18·60
0·012
0·021
39·03
3
264093
NUMBER 1
FeO∗
28·10
0·53
33·66
0·66
0·16
TiO2
3
327099
VOLUME 44
SiO2
0·13
0
0
Type:
2
92·11
327099
1
Oxides
92·26
99·66
0·276
0·422
50·13
0·113
7·65
0·127
0·014
40·93
Sample no.: 327099
Analysis:
mg-no.
100·29
0·269
CaO
Sum
0·523
NiO
50·24
0·135
Cr2O3
MgO
0·010
41·45
TiO2
SiO2
0
0
Type:
2
327099
1
Sample no.: 327097
Analysis:
Olivines
Table 2: Microprobe analyses of minerals from the Manı̂tdlat Member volcanic rocks, West Greenland
JOURNAL OF PETROLOGY
JANUARY 2003
1b
MgO
9
7·85
fe-no
En
0·01
15·57
0·26
2·85
2·46
1·04
85·03
78·39
1·42
41·28
(85·5)
40·36
5·50
61·93
0·35
45·77
(86)
30·49
7·28
65·21
0·41
37·86
85·88
10·34
12·40
74·60
0·56
20·27
(84)
13·15
46·36
63·45
0·59
24·62
Wo
Fs
5·06
47·85
5·26
46·89
0·56
2
48·00
5·55
46·45
0·75
4·49
99·33
0·24
22·85
0·04
15·90
0·00
3·39
1·23
0·70
2·78
0·95
51·26
1a
326787
3
48·65
6·84
44·51
0·79
5·02
99·92
0·24
22·71
0·06
14·94
0·13
3·96
1·18
0·63
4·81
1·06
50·20
1b
264120
4
48·43
8·24
43·33
0·72
6·71
100·47
0·22
22·09
0·01
14·21
0·00
4·81
2·11
0·29
6·61
1·56
48·56
1b
264120
5
48·40
6·09
45·51
0·82
4·52
98·73
0·18
22·91
0·01
15·49
0·00
3·69
0·92
0·50
3·40
0·82
50·81
3
264091
6
50·90
10·79
38·30
0·80
7·53
98·89
0·31
22·31
0·03
12·07
0·00
6·06
1·64
0·03
7·73
2·15
46·57
3
264091
7
50·61
5·02
44·36
0·63
4·44
99·30
0·26
23·28
0·00
14·67
0·14
2·82
1·80
0·66
5·16
0·88
49·63
3
264091
8
47·99
9·28
42·73
0·88
6·07
99·22
0·26
22·17
0·00
14·19
0·12
5·37
0·78
0·03
5·48
1·27
49·55
3
264091
9
54·18
0·77
45·05
0·05
8·17
99·83
0·71
22·43
0·00
13·41
0·00
0·41
8·63
0·00
6·58
2·25
45·42
3
264091
10
50·23
5·19
44·58
0·57
5·30
99·27
0·32
22·94
0·03
14·64
0·00
3·04
2·51
0·45
4·87
1·07
49·40
3
264091
Elements in wt % oxides. FeO∗, total iron as FeO; Fe2O3 and FeO in oxides and clinopyroxenes distributed according to mineral stoichiometry. mg-number is
atomic 100Mg/(Mg + Fe2+), values in parentheses are estimated; cr-number is atomic 100Cr/(Cr + Al); fe-number is atomic 100Fe3+/(Cr + Al + Fe3+). All sample
numbers are Geological Survey of Greenland (GGU) numbers.
Petrographic notes. Olivines: 1 and 2, large low-Ca phenocrysts; 3 and 4, core and rim of a typical high-Ca phenocryst; 5, Mg-poor rim of phenocryst; 6 and 7,
high-Ca phenocrysts; 8, microphenocryst; 9–12, from core to rim of a zoned phenocryst shown in Fig. 4; 13–15, from core to rim of a xenocryst; 16 and 17,
phenocryst and microphenocryst. Oxides: 1–13, primary magmatic chromite crystals; 14 and 15, chromites in two olivine xenocrysts; 16 and 17, ferroan
magnesioferrite xenocryst, core and rim; 18, spinel sensu stricto, tiny blebs in olivine core; 19–23, magnesioferrite, spinel and chromite grains within a single
olivine crystal, from core to rim of the olivine. Clinopyroxenes: 1 and 2, core and rim of microphenocryst; 3 and 4, core and rim of microphenocryst; 5–8, from
core to rim of large phenocryst; 9 and 10, core and margin of microphenocryst.
86·21
1·58
cr-no.
Ol mg-no.
78·95
mg-no.
0·29
73·12
17·02
0·58
FeO∗
Fe2+/Fe
99·65
0·16
14·19
0·23
14·57
11·17
0·21
32·63
4·80
1·18
49·50
1a
0·23
99·94
0·25
18·66
0·19
11·32
9·94
0·20
10·16
1·03
0·14
25·32
1b
1
326787
100·00
99·96
100·31
99·50
0·33
16·21
0·22
15·41
24·95
0·24
3·94
48·13
0·95
0·14
1b
23
264124
Na2O
99·64
4·41
0·16
33·63
1b
22
264124
Sum
0·34
14·54
0·24
15·93
33·16
0·19
2·56
29·54
2·98
0·16
1b
21
264124
22·10
0·28
8·31
0·24
21·05
57·87
0·07
0·22
20
264124
Clinopyroxenes
CaO
0·31
0·12
20·66
MnO
NiO
8·01
0·14
V2 O 3
9·81
1·41
Cr2O3
FeO
59·06
Al2O3
Fe2O3
1·74
0·60
TiO2
10·04
0·14
0·19
SiO2
1b
Type:
19
264124
18
Sample no.: 176718
Analysis:
Oxides
Table 2: continued
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
JOURNAL OF PETROLOGY
VOLUME 44
aggregates. In contrast to the fresh glassy pillow breccias,
the crystalline samples are affected by zeolite-facies metamorphism, and interstices are filled with green and brown
smectites and colourless fine-grained aggregates of zeolites
and Ca-hydrosilicates. These interstices probably include
the breakdown products of nepheline. Vesicles are filled
with massive zeolites. Sulphides form secondary pyrite
in alteration zones, and primary sulphide liquid drops
infrequently preserved as tiny 1–5 m spherules in glass
inclusions in olivine phenocrysts in the Mg-rich Types 0
and 1a.
NUMBER 1
JANUARY 2003
enclosed in olivine and, in Type 0 picrites only, also
within the matrix glass. Some olivine cores are speckled
with tiny oxide inclusions ranging from dust-size particles
to greenish brown bleb-like grains and opaque vermicular
grains. Rare opaque oxide xenocrysts, fringed with clinopyroxene crystals, occur free in the matrix glass.
The compositional variation of the oxides is illustrated
in Fig. 4 and Table 2. The primary magmatic chromites
have cr-numbers [atomic 100Cr/(Cr + Al)] between 40
and 60, generally lower than in the tholeiitic rocks (Fig.
4a). The dyke chromites have high cr-numbers. There is a
good correlation between the mg-number of the chromites
and the enclosing olivines (not shown); chromites with
mg-number 75–80 are enclosed within olivines with mgnumber 91–92·5, and the iron-rich olivine xenocrysts
contain correspondingly iron-rich chromites.
The primary magmatic chromites show differences in
the iron oxidation states between the magma types (Fig.
4b; Table 2). The Type 1a chromites have the highest
Fe2+/Fe whereas those in Type 1b have the lowest.
With progressive crystallization (decreasing mg-number)
the chromites become more reduced (higher Fe2+/Fe),
in contrast to the tholeiitic chromites, which become
more oxidized (lower Fe2+/Fe).
The opaque oxide xenocrysts are ferroan magnesioferrites, which have thin rims that are more reduced
and more magnesian than the centres. They are extremely
low in both Cr and Ti (Table 2, numbers 16 and 17).
The vermicular and bleb-like oxides in some olivine
crystals range from ferroan magnesioferrite in olivine
cores through greenish bleb-like Al-rich spinel sensu stricto
to chromian spinel approaching normal magmatic compositions in the olivine rims (Fig. 4c). These oxides
are not magmatic but are solid-state high-temperature
oxidation and re-equilibration products, as discussed
below.
Olivine
Olivine comprises several textural types similar to those
from the tholeiitic rocks described by Larsen & Pedersen
(2000). Most olivines are clear euhedral to subhedral to
skeletal phenocrysts; some have inclusion-filled zones
and healed cracks, and others have cores speckled with
numerous small inclusions of oxides and sometimes glass.
Some olivines are obviously xenocrystic, with anhedral
and serrated outlines.
The olivines span the compositional range mg-number
92·3–77·4 (Table 2) with a compositional gap around
mg-number 90. The most magnesian olivines (mg-number
>90) are found in the most magnesian rocks and there
is a crude correlation between the olivine compositional
range within a sample and the bulk-rock MgO contents,
as also found in the tholeiitic rocks (Larsen & Pedersen,
2000). All the olivines, including those with mg-number
>90, have glass inclusions and high contents of CaO and
Cr2O3, indicating a magmatic origin. Possible mantle
xenocrysts would have very low contents of CaO and
Cr2O3 (Larsen & Pedersen, 2000) and have not been
found.
The minor elements MnO, CaO, Cr2O3 and NiO,
measured with high precision, show distinct differences
between olivines in alkaline and tholeiitic rocks (Fig. 3).
First, olivines in the alkaline rocks show a far greater
scatter than those in the tholeiites. Second, the main
olivine populations in the alkaline rocks have distinctly
higher contents of CaO (and MnO, not shown), similar
or higher Cr2O3, and lower NiO than the tholeiitic
olivines. Within a single sample, olivine crystals with
widely different minor-element contents and zoning patterns may exist side by side. A few olivine crystals, often
the larger ones, have minor-element contents similar to
those of the tholeiitic olivines. This is particularly evident
for some olivines with low CaO. The zoning patterns in
some individual crystals are also shown in Fig. 3, and
the significance of the data is discussed below.
Clinopyroxene and plagioclase
The clinopyroxenes show complicated oscillatory zoning,
which is a feature very commonly found in alkaline
rocks. The microphenocrysts generally correspond to the
outermost 2–3 zones of the larger phenocrysts. Compositionally, however, all clinopyroxene phenocrysts span
a relatively narrow range, En42–48 Fs4–10 Wo42–52, and they
are thus classical diopsides (Table 2). There is little or no
difference between the clinopyroxenes from the various
chemical rock types, except that those of Type 2 tend to
have slightly higher Ti and slightly lower Wo.
Plagioclase microphenocrysts with slight normal zoning
(in Type 3) and microlites (in Types 1–3) span the
compositional range An87–71 Or0·7–2·3. There are no differences between the various rock types.
Oxides
The primary magmatic oxides are brown, semi-transparent chromites, which occur as small euhedral crystals
10
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 3. Minor elements (wt % oxides) in olivines in the alkaline rocks (high-precision analyses). The compositional fields for the contemporaneous
tholeiitic picrites are based on data of Larsen & Pedersen (2000). Left panel: all analyses, showing the far greater scatter for the alkaline than
for the tholeiitic rocks. Right panel: compositional variation from core (c) to rim (r) of four phenocrysts from two samples of Type 1b (GGU
264120 and 264124). The compositional scatter and the zoning patterns can be explained by mixing and re-equilibration of low-CaO olivines
from tholeiitic into alkaline magmas, as discussed in the text.
Primary sulphides
Matrix glasses and glass inclusions
Tiny globules of sulphide preserved within glass inclusions
in olivine have chemical compositions close to Fe–Ni
monosulphide with small amounts of Cu. Most globules
are too small for ‘clean’ microprobe analyses, but energy
spectra show about equal amounts of Fe and Ni in
sulphide in olivine with mg-number 92, and successively
decreasing Ni in sulphides in olivines with mg-number
88–85. Similar globules are also present in the tholeiitic
rocks. The globules seem to be too large to be exsolved
from the trapped liquid and are considered to be trapped
as liquid sulphide together with the silicate melts.
Matrix glasses were analysed in all compositional types.
Glass inclusions in olivine phenocrysts, representing melts
at earlier stages of crystallization, were analysed in Mgrich samples of Types 0, 1 and 2. The glass inclusions
were not homogenized and have lost olivine as a result
of post-entrapment crystallization, but as long as there
are no other daughter minerals the incompatible element
ratios of the trapped melts should be unaffected.
The alkaline matrix glasses (Table 3) are homogeneous
and relatively fractionated, with 6·4–7·9 wt % MgO and
elevated contents of the incompatible elements Ti, Na,
11
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
the main trend goes from the matrix glasses towards the
TiO2 apex. A plot of CaO/Al2O3 versus the composition
of the olivine (Fig. 5c) shows that low-CaO olivines have
glass inclusions with low CaO/Al2O3; these are the same
as those with high TiO2/P2O5 in Fig. 5b.
The glass inclusions in the high-CaO olivines have
CaO/Al2O3 ratios of 0·9–1·2. Except for Type 0 there
is a general tendency of declining CaO/Al2O3 with
increasing crystallization (decreasing olivine mg-numbers)
to low values in the matrix glasses, distinctly lower than
in the bulk rocks. This is ascribed to the formation of
the abundant clinopyroxene microphenocrysts in the
matrix glasses in these rocks.
Crystallization temperatures and oxidation
states of alkaline vs tholeiitic melts
During quenching the last olivine rims that formed in
contact with the alkaline matrix glasses have mg-numbers
varying from 83·5 in the Type 0 picrites to 79·6 in the
Type 3 ankaramites. When the oxidation state of iron
in the glass is derived by assuming an olivine–melt Fe–Mg
distribution coefficient at 1 atm of 0·30 (Roeder & Emslie,
1970), quench temperatures can be calculated after Ford
et al. (1983) and range from 1190°C in the Type 0
picrites to 1150°C for the ankaramites. In most cases the
measured and calculated olivine compositions are very
similar, indicating equilibrium between olivine rims and
glass. In comparison, the quench temperatures for the
tholeiites are in the range 1210–1180°C for rocks without
plagioclase phenocrysts (Larsen & Pedersen, 2000).
The complicated oxides indicate early intratelluric
events of oxidation of the olivine phenocrysts, leading to
oxidation-exsolution of magnesioferrite (Khisina et al.,
1995), and subsequent solid-state re-equilibration at low
oxygen fugacities, leading to formation of Fe3+-poor, Alrich spinel and then to Fe3+-poor Cr-spinel (Fig. 4c).
During both primary chromite crystallization and reequilibration, the oxides from the alkaline rocks show
progressively decreasing oxidation states, in contrast to
Fig. 4. Compositional variation of the oxides in the alkaline rocks.
Primary magmatic chromites are shown with different symbols for
the various alkaline types. Chromites in iron-rich olivine xenocrysts,
magnesioferrite xenocrysts, and solid-state oxidation and re-equilibration products of bleb-like oxides in olivine are shown with one
symbol each, irrespective of the type they occur in. A fine line connects
individual oxide blebs within a single olivine crystal, with the arrow
pointing from core to rim of the olivine (GGU 264124, Type 1b). For
the magnesioferrite xenocrysts, arrows in (b) connect cores and rims of
grains. The compositional fields for primary chromites from the tholeiitic
picrites are based on data of Larsen & Pedersen (2000). (a) cr-number
[100Cr/(Cr + Al)] vs mg-number [100Mg/(Mg + Fe2+)]. (b) Fe2+/
Fe vs mg-number. Oxides from Types 0, 1a and 3 are highlighted to
clarify the different levels of oxidation state. (c) Oxides projected onto
the end of the spinel prism Al–Fe3+–Cr.
K and P relative to the bulk rocks. In a K2O–TiO2–P2O5
triangular diagram (Fig. 5), the matrix glasses of the
various alkaline types have well-defined K2O–TiO2–P2O5
ratios (Fig. 5a), with Types 0, 1a and 1b being closely
similar. The tholeiitic glasses with their very low K2O
and P2O5 contents plot close to the TiO2 apex. The
inclusion glasses (Fig. 5b) show considerable scatter, but
12
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Table 3: Analyses of alkaline and tholeiitic matrix glasses, West Greenland
Alkaline rocks
Sample:
264166
327099
327097
326786
326788
326787
264120
176718
264122
Type:
Stordal
0
0
1a
1a
1a
1b
1b
1b
45·24
SiO2
48·38
45·69
46·26
46·45
46·96
47·05
45·30
45·40
TiO2
2·00
1·64
1·67
2·07
2·11
2·07
1·97
2·11
1·94
Al2O3
14·62
13·94
14·01
14·89
14·54
14·59
14·71
14·69
14·78
Cr2O3
0·070
0·052
0·052
0·040
10·98
0·035
11·12
0·038
FeO∗
10·46
11·06
11·12
MnO
0·16
0·22
0·23
MgO
6·97
7·88
7·91
7·23
6·95
6·98
NiO
0·003
0·014
0·015
0·004
0·008
0·005
n.a.
0·26
11·04
0·017
0·018
0·019
12·16
11·96
11·90
0·20
0·19
0·21
6·52
6·53
6·48
0·009
0·010
n.a.
0·012
CaO
12·48
15·59
15·64
13·71
13·84
13·91
13·76
13·80
13·76
Na2O
2·44
2·09
2·07
2·20
2·22
2·18
2·75
2·70
2·62
K2O
0·77
0·83
0·86
1·17
1·17
1·13
1·10
1·02
1·02
P2O5
0·31
0·61
0·62
0·61
0·64
0·63
0·66
0·59
0·63
Sum
98·67
99·61
100·46
99·35
99·85
99·62
99·15
99·02
98·61
H2O
2·21
0·32
0·34
0·62
0·31
CO2
0·07
0·06
0·09
0·07
0·13
Fe2+/FeTotal
0·925
Alkaline rocks
Tholeiitic rocks
Sample:
264113
264114
264093
176701
326785
136943
327100
264109
264137
Type:
2
2
3
3
Dyke
Below1
Between1
In MMb1
Above1
48·48
SiO2
46·93
46·70
44·15
44·49
48·77
47·72
47·55
48·38
TiO2
2·69
2·74
1·76
1·76
1·67
1·60
1·78
2·72
1·61
Al2O3
14·37
14·40
14·75
14·75
14·87
15·11
14·88
13·51
15·05
Cr2O3
0·043
0·040
FeO∗
11·04
11·31
MnO
0·18
MgO
6·66
6·80
NiO
0·009
0·006
n.a.
0·007
0·008
0·053
0·049
0·035
0·042
12·21
11·34
10·62
10·48
11·02
0·21
0·21
0·19
0·17
0·15
6·46
6·66
8·10
8·35
8·28
7·19
0·007
0·007
0·016
0·017
0·018
0·016
0·040
11·27
10·49
n.a.
n.a.
8·48
0·020
CaO
13·04
13·33
14·62
14·71
12·83
12·97
13·15
12·18
12·72
Na2O
2·04
1·96
2·63
2·60
2·20
2·00
2·17
2·40
2·07
K2O
1·17
1·13
1·32
1·37
0·40
0·16
0·25
0·46
0·16
P2O5
0·41
0·38
1·06
1·06
0·23
0·12
0·17
0·29
0·16
Sum
98·58
98·80
99·18
98·97
99·95
98·74
99·45
98·46
99·28
H2O
1·55
0·28
CO2
0·20
0·07
Fe2+/FeTotal
0·14
0·02
0·955
0·821
Major elements by microprobe, H2O and CO2 by CHN-analysis of separated glass chips. Analyses in wt % oxides. Fe2+/FeTotal
by Mössbauer analysis of separated glass chips; analyst C. Bender-Koch. The analyses are arranged in the same order as
the bulk-rock analyses in Table 4. The glass sample 264166 corresponds to bulk sample 264167 in Table 4. Glass inclusions
are not tabulated. n.a., not analysed.
1
Stratigraphic position: Below indicates below Type 0; Between indicates between Type 0 and the Manı̂tdlat Member; Above
indicates above Manı̂tdlat Member.
13
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
than in the tholeiitic glasses. Thus, there is circumstantial
evidence that the alkaline melts were significantly more
reduced than the tholeiitic melts.
WHOLE-ROCK GEOCHEMISTRY
Major elements
Representative XRF analyses are shown in Table 4, and
plots of the major elements vs MgO are shown in Fig.
6. The alkaline rocks contain 7–23 wt % MgO, and all
have much higher contents of P2O5 and K2O than the
contemporaneous tholeiites. They form a number of
compositional groups (types, presented in Table 1), which
are particularly evident in the P2O5 and TiO2 diagrams.
Most of the alkaline rocks have low SiO2 and high CaO
compared with the tholeiites. Types 0 and 3 show the
greatest relative enrichment in CaO and P2O5 but no
enrichment in TiO2. In contrast, Type 2 rocks have
relatively low CaO and P2O5 but the greatest enrichment
in TiO2 and K2O. The alkaline rocks have only slightly
lower Al2O3 than the tholeiites, whereas the levels of
FeO∗ are the same or slightly lower. Perhaps surprisingly,
the alkaline rocks are not enriched in Na2O; indeed,
many have lower Na2O contents than the tholeiites.
Although Na2O is somewhat scattered as a result of
secondary alteration, low Na2O is a primary feature of
the magmas and is also seen in the analyses of the fresh
matrix glasses (Table 3 and Fig. 6). The tholeiitic glasses
have the same Na2O contents as the glasses of Types 0
and 2; the highest Na2O contents are seen in the glasses
of Types 1 and 3 that were quenched after significant
clinopyroxene crystallization. The K2O diagram shows
considerable scatter, much of which is due to secondary
redistribution of K in the lava samples, which, as described above, are often altered whereas the pillow breccias are fresh. The data from the pillow breccias alone
strongly suggest that the Type 2 melts were generated
with higher K2O than the other types. The dyke has
tholeiitic abundances of most of the major elements but
has slightly elevated contents of P2O5 and K2O.
Three Stordal samples and four Type 1a samples have
increased SiO2 and decreased CaO relative to other
rocks of the same type; together with other elemental
fingerprints this suggests these samples are crustally contaminated, as discussed below.
For each rock type, the compositional variation seen
in Fig. 6 is dominantly caused by olivine fractionation
and accumulation; the clinopyroxene-phyric Type 3 rocks
also show evidence of clinopyroxene fractionation, or
perhaps accumulation, in the changed slope of the CaO
trend. CaO/Al2O3 is not changed by olivine fractionation
or accumulation, and the different levels of this ratio in
the various magma types (Fig. 7) may be features of the
primary magmas. Types 0, 1b and 3 have the highest
Fig. 5. Microprobe analyses of glasses. (a) and (b) the incompatible
elements K2O, TiO2 and P2O5 in matrix glasses and in glass inclusions
in olivine. The glass inclusions with relatively high TiO2, and also those
with low P2O5, are hosted in low-Ca olivines. (c) CaO/Al2O3 in glass
inclusions in olivine plotted against the composition of the enclosing
olivine; average matrix glasses are plotted against the composition of
the outermost olivine rims. Some low-Ca olivines are zoned with CaOrich rims, and lines connect inclusions within core and rim of two such
olivine crystals. Bars at the left side of the diagram show CaO/Al2O3
ratios for the bulk rocks.
the evolution trend in the chromites from the tholeiitic
rocks (Fig. 4b). Mössbauer analysis of matrix glasses
(Table 3) also shows higher Fe2+/Fe in the alkaline
14
15
50·10
1·44
13·86
3·10
7·24
0·17
7·76
10·05
2·24
0·84
0·31
2·46
99·57
61·01
46·95
1·80
13·96
4·75
6·06
0·16
7·61
12·77
1·80
0·71
0·41
2·82
99·80
59·83
Trace elements by XRF
Rb
37
17
Ba
545
555
Pb
5
2
Sr
400
580
La
49
72
Ce
82
128
Nd
37
59
Y
27
27
Th
5
5
Zr
150
153
Nb
26
65
Zn
88
76
Cu
19
64
Co
46
48
Ni
39
65
Sc
25
33
V
260
282
Cr
770
326
Ga
22
20
SiO2
TiO2
Al2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
Volat.
Sum
mg-no.
264167
Stordal
Neck
18
418
1
455
56
100
46
27
5
144
59
95
47
55
70
32
277
498
23
47·51
1·89
13·58
5·66
6·03
0·17
7·97
11·28
2·25
0·73
0·30
2·01
99·38
59·17
264168
Stordal
Neck
Stordal volcanic centre
Sample no.: 2641651
Type:
Stordal
Lithology:
Lava
Strat. pos.:
Alkaline rocks
3·1
577
3
453
106
152
54
17
0
123
67
91
94
91
896
28
219
1580
12
42·64
0·99
8·82
2·52
8·53
0·20
22·34
9·80
1·19
0·21
0·39
2·63
100·26
80·71
327099
0
Pillow
4·7
865
8
609
107
154
57
18
3
128
70
86
99
90
876
30
221
1630
14
42·37
1·03
9·02
2·18
8·87
0·19
21·33
11·31
0·75
0·23
0·40
2·54
100·22
79·93
327097
0
Pillow
6·1
1380
8
750
136
191
71
18
2
131
74
90
103
88
851
27
223
1480
12
42·59
1·03
9·26
2·04
8·86
0·20
20·92
11·21
0·71
0·37
0·43
2·31
99·93
79·82
327096
0
Pillow
Low in Ordlingassoq Mb
18
572
4
587
101
143
52
18
1
139
86
131
149
76
569
26
228
1310
12
42·74
1·30
9·36
2·33
8·76
0·19
20·25
10·29
0·70
0·57
0·39
2·68
99·56
79·05
135989
1a
Lava
6·2
527
6
551
86
127
48
19
0
142
87
91
54
76
411
32
251
1250
18
43·82
1·31
10·41
2·08
8·58
0·19
17·31
11·80
0·84
0·37
0·39
2·90
100·00
77·01
326786
1a
Pillow
Manı̂tdlat Member
7·1
557
6
614
94
142
53
18
2
145
92
88
54
72
387
31
255
1380
13
43·04
1·35
10·68
2·04
8·72
0·19
17·00
11·61
1·03
0·33
0·40
3·66
100·15
76·51
326788
1a
Pillow
5·7
651
3
660
93
135
52
20
2
151
92
87
57
72
360
31
255
1240
16
44·27
1·38
10·96
2·10
8·56
0·19
16·42
12·12
1·00
0·46
0·41
2·27
100·14
76·07
326787
1a
Pillow
34
457
7
364
58
93
36
21
0
148
61
88
53
61
246
30
256
953
21
45·60
1·61
11·90
3·05
7·96
0·19
14·80
9·98
1·42
0·88
0·33
2·37
100·09
73·66
1369411
1a
Lava
28
719
8
680
128
188
68
22
3
177
110
88
62
59
256
39
316
820
19
44·60
1·69
12·70
3·10
7·50
0·19
12·30
13·40
1·20
0·79
0·55
1·97
99·99
70·74
113254
1a
Lava
17
652
7
1030
81
137
62
27
7
182
93
133
169
47
94
30
286
383
15
45·47
1·94
13·28
3·01
8·34
0·20
8·47
13·14
1·72
0·78
0·56
2·75
99·66
60·79
135931
1a
Lava
3·6
2290
8
816
194
270
98
22
2
141
99
95
122
68
309
30
292
828
19
43·92
1·54
11·10
4·01
7·84
0·21
12·72
13·72
0·80
0·43
0·48
2·62
99·39
69·21
264120
1b
Pillow
1·8
1060
11
805
174
240
89
23
3
149
87
93
121
68
291
32
293
752
24
43·68
1·59
11·74
3·68
7·99
0·21
12·34
13·34
1·65
0·18
0·44
3·35
100·19
68·83
264116
1b
Pillow
3·8
2280
12
960
185
260
98
22
7
147
97
94
133
64
258
36
305
734
13
43·79
1·56
11·96
2·59
8·79
0·21
12·28
13·49
1·22
0·39
0·46
2·66
99·40
69·07
264124
1b
Pillow
Table 4: Chemical analyses of alkaline rocks of the Manı̂tdlat Member and contemporaneous tholeiitic volcanic rocks, West Greenland
1·7
1250
10
827
148
217
81
22
5
142
83
92
125
62
296
35
304
801
19
44·26
1·66
12·08
3·24
8·45
0·21
12·04
13·32
1·55
0·33
0·47
2·37
99·98
68·18
176718
1b
Pillow
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
45·20
1·66
12·46
2·33
9·29
0·20
10·70
14·04
0·95
0·60
0·52
1·06
99·01
65·53
44·85
1·88
11·01
2·49
8·92
0·18
15·66
10·48
1·36
0·77
0·28
1·59
99·47
73·95
30
232
4
238
23
37
20
19
0
222
77
85
103
71
263
33
307
735
17
44·10
2·14
12·24
3·55
8·03
0·18
13·09
10·53
1·03
0·88
0·31
4·32
100·40
70·23
136916
2
Lava
48
312
0
260
25
43
26
20
1
229
79
90
116
68
224
34
314
651
20
44·34
2·20
12·50
3·93
7·77
0·18
11·97
11·10
1·16
1·11
0·32
3·79
100·37
68·17
136915
2
Lava
34
275
5
362
28
42
21
21
3
193
70
91
94
73
183
35
342
733
21
45·92
2·30
12·46
2·79
8·79
0·18
11·26
12·22
1·27
0·91
0·32
1·32
99·74
66·84
264114
2
Pillow
33
270
5
269
27
52
27
21
0
236
83
84
101
60
189
38
356
672
19
45·20
2·40
12·90
5·40
6·30
0·17
10·70
11·60
1·24
0·96
0·39
3·58
100·84
65·98
113253
2
Lava
16
1130
7
1020
152
232
92
24
5
193
121
104
84
55
132
39
324
520
18
44·97
1·47
11·81
1·04
10·06
0·21
10·47
15·73
1·43
0·69
0·80
1·53
100·21
65·82
176705
3
Pillow
14
1570
7
1150
153
233
93
26
5
202
126
107
97
60
114
35
312
418
23
43·92
1·47
12·50
2·05
9·16
0·21
9·74
15·45
1·22
0·60
0·78
2·75
99·85
64·16
264093
3
Pillow
13
1060
6
1060
161
244
98
25
3
200
125
106
91
53
114
38
320
444
22
44·65
1·51
12·48
0·24
10·76
0·22
9·61
15·50
1·82
0·51
0·85
1·96
100·11
63·91
176701
3
Pillow
26
1230
6
1170
169
266
103
30
5
230
142
112
112
54
81
26
305
245
25
43·48
1·73
13·06
3·81
8·04
0·22
8·33
14·56
1·63
1·05
0·96
1·59
98·46
59·50
264145
3
Lava
34
1910
7
1400
207
307
117
32
4
251
171
112
142
51
56
20
261
123
20
44·03
1·86
13·72
4·93
8·00
0·23
6·66
14·44
1·73
1·39
1·15
0·89
99·03
52·00
264106
3
Lava
2
82
4
234
12
24
11
18
2
84
16
91
120
84
688
30
280
1340
15
45·71
1·23
11·25
1·96
8·85
0·18
17·04
10·49
1·28
0·22
0·19
1·78
100·18
76·46
326785
Dyke
Chill
Dyke
Below
Between In M.Mb Above
0·7
31
5
111
3
6
5
16
1
58
4·4
86
113
88
802
32
258
1260
13
44·70
0·96
11·39
2·84
8·05
0·18
18·81
9·93
1·32
0·05
0·10
2·21
100·55
78·20
0·9
31
1·5
138
4
9
8
19
0
67
3·7
81
141
71
802
33
299
1327
15
45·08
1·17
11·23
3·63
7·62
0·17
18·50
9·59
1·30
0·18
0·09
1·44
100·00
77·46
0·6
35
0
163
4
10
8
16
0
73
6·9
86
85
81
710
30
295
1240
19
44·18
1·25
10·65
4·46
7·20
0·18
18·70
9·93
1·26
0·08
0·14
2·42
100·45
77·13
2·5
56
6
190
6
12
10
17
2
73
5·7
97
116
83
761
30
301
1270
18
45·00
1·40
10·70
3·78
7·72
0·17
18·11
8·90
1·78
0·12
0·17
2·07
99·92
76·71
2·3
23
4
132
1
12
7
17
0
64
4
95
124
84
848
33
281
1260
17
45·42
1·14
10·40
2·36
8·98
0·17
19·84
8·66
1·38
0·12
0·12
1·18
99·77
78·33
326783
136943
327100
264109
264137
Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite
Pillow
Pillow
Pillow
Pillow
Pillow
Below
Tholeiitic rocks of the Ordlingassoq Member
VOLUME 44
16
NUMBER 1
Major elements in wt %, trace elements in ppm. Trace elements by John Bailey, Geological Institute, University of Copenhagen. Volat., loss on ignition corrected
for oxygen uptake during ignition (Kystol & Larsen, 1999). mg-number is atomic 100Mg/(Mg + Fe2+), calculated with Fe2O3/FeO (wt %) adjusted to 0·15. For
stratigraphic position of tholeiites, see Table 3.
1
Sample 264165 (Stordal) with high SiO2 is considered to be crustally contaminated. Sample 136941 (Type 1a) represents a group of samples that have increased
levels of SiO2, K2O and Rb, and decreased levels of CaO, P2O5, Nb and several other trace elements, also considered to be due to crustal contamination.
Trace elements by XRF
Rb
9·6
26
Ba
2390
271
Pb
12
2
Sr
1210
315
La
207
27
Ce
295
47
Nd
107
24
Y
24
18
Th
6
8
Zr
159
153
Nb
111
58
Zn
94
87
Cu
146
60
Co
69
71
Ni
210
368
Sc
35
33
V
307
293
Cr
596
1270
Ga
18
15
SiO2
TiO2
Al2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
Volat.
Sum
mg-no.
264113
2
Pillow
Manı̂tdlat Member
Sample no.: 264122
Type:
1b
Lithology:
Pillow
Strat. pos.:
Alkaline rocks
Table 4: continued
JOURNAL OF PETROLOGY
JANUARY 2003
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 6. Bulk-rock major elements vs MgO for the alkaline rocks of the Manı̂tdlat Member. All data recalculated to 100%, volatile-free. The
contemporaneous tholeiites of the Ordlingassoq Member are shown as fields labelled ‘tholeiites’. Some alkaline magma types in some diagrams
have been outlined for clarity. Grey symbols in all diagrams denote samples with increased SiO2 and K2O and decreased CaO and P2O5,
considered to be due to crustal contamination.
CaO/Al2O3 ratios, up to 1·35, caused by both high CaO
and low Al2O3 (Fig. 6). The sloping CaO/Al2O3 trend
in the Type 3 rocks is produced by clinopyroxene fractionation or accumulation. Type 2 and Stordal have
CaO/Al2O3 ratios corresponding to the tholeiites.
All samples except for three evolved ones contain
<3 wt % total alkalis, and according to the IUGS
classification (Le Bas, 2000) the rocks are simply picrites
and basalts. The alkaline character is better reflected in
the CIPW norms, particularly of the matrix glasses (Fig.
8). None of the Type 2 and Stordal bulk rocks are ne
normative; Type 1 is mixed, and all rocks of Type 0 and
Type 3 are ne normative although the maximum ne is
only 5·3%. However, all matrix glasses except Stordal
and the dyke are ne normative, with maximum ne =
9·8–10·2% in Type 3. The three high-Si Stordal lavas
17
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
contrast, Type 2 is relatively more enriched in Zr and
has lower Nb/Zr ratios of 0·3–0·4, but still higher than
the tholeiitic values of <0·14. The dyke is close to the
tholeiitic values for many elements but is clearly enriched
in all the incompatible elements (see also Fig. 11). It has
a Nb/Zr ratio of 0·19.
The compatible elements Ni and Cr show the effects
of olivine and chromite control; however, the alkaline
rocks have distinctly lower Ni contents than tholeiites
with similar MgO. The diagrams for V and Sc show a
maximum at 10–11 wt % MgO, indicating clinopyroxene
fractionation in magmas with <10 wt % MgO, in accordance with the major-element variations.
A wider range of trace elements was obtained by ICPMS analysis of two samples of each of the alkaline types
and one of the dyke (Table 5). The two samples from
each alkaline type gave closely similar results, except for
the Stordal samples, of which one (264165) is a high-Si
variety. These data are presented in Figs 10 and 11 as
rare earth element (REE) and multi-element diagrams
with comparisons.
Two element enrichment patterns can be clearly distinguished from the multi-element diagrams. One pattern
is common to the rocks of Types 0, 1a, 1b and 3. It
shows extreme enrichment in Ba, U, Nb–Ta and light
to middle REE (LREE to MREE), deep troughs for Rb,
Th, K, Pb and Ti, and lesser troughs for Sr, P and Zr–Hf.
In comparison, melilitites, nephelinites and meimechites
show similar levels of enrichment for many elements but
have much smoother spectra, with moderate K troughs
being the most distinctive (Rogers et al., 1992; Arndt et
al., 1995, 1998; Hegner et al., 1995; Wilson et al., 1995;
Bernstein et al., 2000; Späth et al., 2001). Enriched picrites
from Nuanetsi have incompatible element concentrations
up to 200 times primitive mantle; however, their spectra
show no K anomalies and have distinct Nb–Ta troughs
Fig. 7. CaO/Al2O3 vs MgO for the alkaline rocks of the Manı̂tdlat
Member. Grey symbols denote crustally contaminated samples. Samples
with MgO <10 wt % have lost CaO by clinopyroxene fractionation.
The contemporaneous tholeiites are shown as an outlined area.
are slightly Q normative (not plotted). Figure 8 also shows
the high contents of normative diopside, the Type 3
ankaramites attaining a maximum of 40% di. The Type
3 glasses have somewhat decreased di because of clinopyroxene fractionation, whereas the Type 0 glasses with
39% di were quenched just before clinopyroxene saturation was reached.
Trace elements
Figure 9 shows a range of trace elements plotted against
MgO. The incompatible elements Ba, Sr, Nb and La
show different levels of enrichment relative to the tholeiites, and Type 1 clearly splits up in two groups (a and
b). Types 0, 1b and 3 are the most strongly enriched in
Ba, Sr and La, Type 1a is intermediate, and Type 2 and
Stordal are the least enriched. Types 0, 1 and 3 have
extremely high Nb/Zr ratios in the range 0·5–0·7. In
Fig. 8. CIPW-normative character of the alkaline rocks and matrix glasses of the Manı̂tdlat Member. Norms calculated with wt % Fe2O3/FeO
adjusted to 0·15. The parameter plotted on the horizontal axis is calculated as hy − ne. Crustally contaminated rocks (grey in Fig. 6) are not
plotted. The fields of the tholeiites are shown as outlined areas.
18
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 9. Trace elements (XRF data) vs MgO for the alkaline rocks of the Manı̂tdlat Member. The contemporaneous tholeiites are shown as
fields labelled ‘tholeiites’. Some alkaline magma types in some diagrams have been outlined for clarity. Grey symbols in all diagrams denote
samples considered to be crustally contaminated; these are the same as in Fig. 6.
(Ellam & Cox, 1989, 1991). Carbonatites and allegedly
carbonatite-metasomatized mantle xenoliths show variable degrees of enrichment and have spiky patterns with
similarities to those of the Manı̂tdlat Member, showing
deep troughs for Rb, K, Zr and Ti, and strong relative
enrichment in Ba, Nb and LREE (Nelson et al., 1988;
19
264165
Stordal
38·9
410
25·9
149
28·4
0·90
562
46
90
10·4
40
6·9
2·06
5·86
0·91
5·04
0·92
2·34
0·33
2·10
0·32
3·76
1·54
4·76
4·43
1·29
0·70579
0·70556
14·27
0·512349
0·512310
−4·87
18·489 ± 0·127
15·252 ± 0·114
38·175 ± 0·240
18·330
15·244
37·995
Sample:
Type:
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
( 87Sr/86Sr)m
( 87Sr/86Sr)60
Sr60
( 143Nd/144Nd)m
( 143Nd/144Nd)60
Nd60
( 206Pb/204Pb)m
( 207Pb/204Pb)m
( 208Pb/204Pb)m
( 206Pb/204Pb)60
( 207Pb/204Pb)60
( 208Pb/204Pb)60
Alkaline rocks
16·6
590
27·1
140
74·3
0·28
530
68
134
15·3
58
9·2
2·71
6·74
1·03
5·32
0·97
2·46
0·34
2·12
0·32
3·51
3·14
3·53
4·95
1·60
0·70409
0·70402
−7·52
0·512357
0·512320
−4·72
19·994 ± 0·058
15·405 ± 0·046
39·005 ± 0·103
19·717
15·392
38·724
264167
Stordal
3·5
458
18·8
122
69·9
0·52
577
92
157
16·5
58
7·9
2·15
4·87
0·76
3·86
0·68
1·64
0·23
1·41
0·22
2·84
3·31
4·45
2·54
1·72
0·70372
0·70370
−12·06
0·512407
0·512370
−3·64
20·234 ± 0·028
15·558 ± 0·021
38·560 ± 0·052
19·998
15·547
38·446
327099
0
5·6
629
19·4
125
73·5
0·23
865
98
169
17·7
61
8·3
2·27
4·91
0·79
3·89
0·70
1·70
0·24
1·39
0·23
2·98
3·44
4·94
2·66
1·75
0·70376
0·70374
−11·53
0·512413
0·512380
−3·51
20·177 ± 0·035
15·574 ± 0·028
38·561 ± 0·090
19·962
15·564
38·453
327097
0
7·3
647
21·4
147
99·2
0·50
534
88
156
16·5
58
8·4
2·28
5·21
0·85
4·21
0·76
1·90
0·27
1·67
0·26
3·61
5·11
3·72
3·25
2·61
0·70446
0·70443
−1·68
0·512363
0·512330
−4·52
21·051 ± 0·122
15·550 ± 0·090
38·619 ± 0·244
20·620
15·530
38·442
326788
1a
264124
1b
6·5
5·1
695
1001
21·9
25·3
152
142
99·3
104·0
0·17
0·11
643
2286
87
170
154
290
16·4
29·9
58
102
8·4
13·0
2·31
3·49
5·58
6·92
0·85
1·10
4·33
5·35
0·78
0·93
1·98
2·19
0·28
0·32
1·66
1·93
0·27
0·29
3·74
3·61
5·09
5·06
3·68
9·77
3·26
4·86
2·60
2·59
0·70438
0·70376
0·70436
0·70375
−2·75
−11·40
0·512406
0·512417
0·512370
0·512380
−3·69
−3·52
21·004 ± 0·051
19·849 ± 0·011
15·512 ± 0·039
15·561 ± 0·010
38·520 ± 0·111
38·524 ± 0·056
20·570
19·689
15·492
15·553
38·341
38·425
326787
1a
Table 5: Trace element and isotope analyses of Manı̂tdlat Member and contemporaneous tholeiitic rocks
12·0
1281
27·3
155
120·7
0·16
2380
191
327
33·6
114
14·4
3·81
7·51
1·18
5·78
0·98
2·34
0·34
2·03
0·31
3·84
5·79
10·45
5·35
3·02
0·70379
0·70377
−11·12
0·512412
0·512380
−3·61
19·868 ± 0·034
15·549 ± 0·026
38·496 ± 0·067
19·693
15·541
38·394
264122
1b
29·3
335
20·7
161
63·6
0·39
266
25
52
6·4
26
4·9
1·53
4·70
0·70
3·90
0·73
1·84
0·27
1·62
0·26
4·39
4·26
2·71
2·12
0·80
0·70572
0·70550
13·54
0·512478
0·512430
−2·52
18·823 ± 0·068
15·307 ± 0·062
38·296 ± 0·166
18·648
15·299
38·143
264113
2
JOURNAL OF PETROLOGY
VOLUME 44
20
NUMBER 1
JANUARY 2003
326785
Dyke
21
136943
Below
0·5
0·9
116
133
18·7
19·8
57
65
3·6
3·5
0·01
0·02
17
26
4·0
3·8
9·9
10·3
1·48
1·62
7·6
8·2
2·37
2·59
0·87
0·94
3·16
2·89
0·53
0·51
3·20
3·26
0·66
0·70
1·72
1·90
0·26
0·27
1·68
1·66
0·26
0·24
1·57
1·85
0·32
0·31
0·73
0·64
0·31
0·30
0·09
0·07
0·70322
0·70321∗
0·70321
0·70319
−19·02
−19·25
0·512971
0·513031∗
0·512900
0·512970
6·54
8·01
17·993 ± 0·142
15·353 ± 0·122
37·728 ± 0·301
17·919
15·350
37·648
326783
Below
264109
In MMb
264137
Above
1·1
2·7
2·1
169
195
135
18·1
19·5
18·0
74
74
63
6·5
4·8
3·3
0·03
0·14
0·03
29
35
24
5·6
4·8
3·7
14·1
12·7
10·0
2·12
1·92
1·59
10·7
10·0
8·5
3·02
2·99
2·62
1·07
1·09
0·96
3·57
3·84
3·30
0·57
0·60
0·55
3·30
3·59
3·32
0·64
0·69
0·66
1·71
1·79
1·73
0·24
0·26
0·24
1·44
1·60
1·55
0·24
0·25
0·24
2·01
1·97
1·75
0·53
0·38
0·28
0·81
0·64
1·13
0·43
0·31
0·24
0·12
0·11
0·08
0·70332
0·70347
0·7033
0·70330
0·70344
0·70326
−17·70
−15·82
−18·29
0·512975
0·513036
0·513083
0·512910
0·512960
0·512980
6·76
7·85
8·11
18·113 ± 0·097
18·317 ± 0·269 18·532 ± 0·055
15·343 ± 0·084 15·432 ± 0·221 15·529 ± 0·050
37·679 ± 0·211
37·811 ± 0·563
37·901 ± 0·157
18·026
18·215
18·489
15·339
15·427
15·527
37·577
37·719
37·860
327100
Between
Tholeiitic rocks of Ordlingassoq Member
Trace elements in ppm. All trace elements by ICP-MS analysis at University of Durham, except for 136943, which is from GEUS, Copenhagen. Subscripts: m, as
measured; 60, age corrected to 60 Ma. For the Sr isotopes, 2 SE = 0·00003. For the Nd isotopes, 2 SE = 0·00005. For the Pb isotopes, 2 SE are given in the table.
For stratigraphic position of tholeiites, see Table 3.
∗Sr and Nd isotopes from Holm et al. (1993).
36·6
31·0
37·9
2·4
374
1161
1481
240
23·5
31·1
35·9
20·3
197
215
255
88
72·4
147·7
195·0
15·9
0·36
0·39
0·36
0·16
267
1241
1896
78
25
154
198
11·8
53
284
360
26·1
6·5
30·8
38·6
3·42
27
111
135
15·0
5·5
15·3
18·2
3·60
1·72
4·22
4·96
1·24
5·17
9·53
10·27
4·04
0·81
1·36
1·61
0·65
4·45
6·59
7·73
3·80
0·84
1·13
1·33
0·72
2·11
2·60
3·08
1·88
0·30
0·36
0·44
0·28
1·92
2·22
2·65
1·67
0·29
0·33
0·40
0·27
5·35
4·75
5·37
2·30
4·99
7·17
9·58
1·13
3·64
5·32
7·30
1·69
2·20
5·55
7·45
0·70
0·69
4·14
5·66
0·53
0·70575
0·70367
0·70371
0·70402
0·70551
0·70360
0·70365
0·70400
13·60
−13·43
−12·82
−7·87
0·512556
0·512391
0·512399
0·512715
0·512500
0·512360
0·512370
0·512660
−1·17
−3·96
−3·79
1·91
19·034 ± 0·065
21·723 ± 0·023 21·743 ± 0·048 19·808 ± 0·155
15·326 ± 0·051
15·595 ± 0·016 15·622 ± 0·034 15·563 ± 0·135
38·491 ± 0·136
38·946 ± 0·217 39·015 ± 0·098 38·729 ± 0·351
18·921
21·238
21·258
19·619
15·321
15·572
15·599
15·554
38·372
38·732
38·806
38·646
264106
3
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
( 87Sr/86Sr)m
( 87Sr/86Sr)60
Sr60
( 143Nd/144Nd)m
( 143Nd/144Nd)60
Nd60
( 206Pb/204Pb)m
( 207Pb/204Pb)m
( 208Pb/204Pb)m
( 206Pb/204Pb)60
( 207Pb/204Pb)60
( 208Pb/204Pb)60
264093
3
264114
2
Sample:
Type:
Alkaline rocks
Table 5: continued
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
most other rocks, and Types 0, 1b, and 3 lie at the
opposite extreme with lower K/Ba and Rb/Sr and higher
La/Nb and Ba/Nb than almost all other rocks. The
significance of the spread in Fig. 12 and the good correlation within the Manı̂tdlat Member is discussed below.
Isotopes
The alkaline rocks of the Manı̂tdlat Member have very
unusual isotope compositions (Table 5). The Sr–Nd–Pb
results for the two samples analysed of each type are
mutually consistent, and the results for the tholeiitic
picrites of the Ordlingassoq Member, analysed simultaneously with the alkaline rocks, are in complete
agreement with earlier results by Holm et al. (1993),
Lightfoot et al. (1997) and Graham et al. (1998).
The tholeiitic rocks have positive Nd and negative
Sr (Fig. 13) and plot within the field for the Iceland
mantle plume (e.g. Stecher et al., 1999). They have been
interpreted by earlier workers (Holm et al., 1993; Lightfoot
et al., 1997; Graham et al., 1998) as produced by melting of
the asthenospheric mantle in the proto-Icelandic mantle
plume, and the data fields for the Ordlingassoq Member
shown in Figs 13 and 14 thus conceivably represent the
local contemporaneous asthenospheric mantle.
The alkaline rocks (excepting the dyke) all have negative
Nd, and all except Type 2 and the high-Si Stordal lava
also have negative Sr (Fig. 13). Types 0, 1b and 3 have
nearly identical Nd–Sr isotope compositions. Except for
Type 2, all the alkaline rocks plot well below the oceanic
mantle array in the Nd–Sr isotope diagram, in an area
of the lower left quadrant occupied by very few uncontaminated mantle-derived rocks. None are known
from the North Atlantic Igneous Province; the few we
have noted are some potassic rocks from the Elkhead
Mountains, Colorado (Leat et al., 1988; Thompson et
al., 1989), some carbonatites and kimberlites from the
Archangelsk region, NW Russia (Mahotkin et al., 2000),
and some nephelinites from the Napak and Mt Elgon
volcanoes in East Africa (Simonetti & Bell, 1994, 1995).
The Pb isotope compositions (Fig. 14) of the tholeiites
of the Ordlingassoq Member cluster around the Northern
Hemisphere Reference Line (NHRL) and fall within the
Iceland field. In contrast, the alkaline rocks have high
206
Pb/204Pb ratios and plot below the NHRL. The basement rocks in the area have low Pb isotope ratios (Fig.
14), and the high-Si Stordal sample is clearly displaced
towards basement values. In terms of the vertical distance
from NHRL as defined by Hart (1984), the Manı̂tdlat
Member rocks have 7/4 = −8 to −25, 8/4 = >0
for Type 2 and 8/4 = −100 to −257 for the other
types, outside the range of all modern mid-ocean basalts
(MORB) and ocean island basalts (OIB) (Thirlwall, 1997).
Fig. 10. Chondrite-normalized REE contents for the alkaline rocks of
the Manı̂tdlat Member and average contemporaneous tholeiite. For
clarity only one sample of each type is shown. The samples have MgO
contents in the range 8–22 wt %, and if the data are recalculated to
a common MgO value the spectra of Types 0, 1b and 3 attain closely
similar levels. Normalization values from McDonough & Sun (1995).
O’Reilly & Griffin, 1988; Yaxley et al., 1991; Larsen &
Rex, 1992; Le Roex & Lanyon, 1998; Coltorti et al.,
2000). The Manı̂tdlat Member rocks have low to very
low Th/U (down to 1·2 in Type 3), whereas most
carbonatites have high Th/U although the ratios are
very variable.
A different element enrichment pattern is seen in the
Type 2 rocks. The incompatible-element enrichment is
less extreme, Rb is enriched, Ba much less so, the troughs
for Th, K, Pb and Sr are very small, and there is a large
Nb–Ta peak and lesser peaks for Zr–Hf and Ti. This
pattern is similar to those of some basanites and alkali
basalts from eastern Australia (O’Reilly & Zhang, 1995;
Zhang et al., 2001), and remarkably similar to patterns
of amphibole from some mantle xenoliths (Moine et al.,
2001).
The Stordal type has an enrichment pattern most
similar to those of Types 0, 1 and 3, but with smaller
troughs for Rb, Th and K. The high-Si Stordal sample
has higher Rb and Pb, and lower U, Nb, Ta and REE
than the other Stordal sample.
The dyke is the least enriched of the alkaline rocks.
Its incompatible element contents all lie between those
of the tholeiites and the other alkaline rocks, and the
trace-element pattern, with low Rb and high Nb/La,
mostly resembles that of Type 1a.
The uniqueness of the Manı̂tdlat Member magmas
may be illustrated by comparison with incompatible
element ratios in other strongly enriched rocks (Fig. 12).
In melilitites, nephelinites and meimechites, K/Ba, Rb/
Sr, Ba/Nb and La/Nb ratios vary over about one order
of magnitude. In the Manı̂tdlat Member rocks these
ratios vary over about two orders of magnitude, and
whereas Type 1a and Stordal often plot together with
other rocks, Type 2 lies at one extreme end with higher
K/Ba and Rb/Sr and lower La/Nb and Ba/Nb than
22
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 11. Primitive mantle normalized multi-element diagrams for the alkaline rocks of the Manı̂tdlat Member and average contemporaneous
tholeiite. The different levels of Types 0 and 3 are mainly an effect of olivine accumulation and fractionation. For the comparisons in the two
lower diagrams, data sources are as follows. Carbonatites: OKU-18, Damaraland (Le Roex & Lanyon, 1998); GGU 265186, Sarfartoq, West
Greenland (Larsen & Rex, 1992, and unpublished data, 2002). Melilitites: BISC-1 and ZHC-1, Namaqualand (Rogers et al., 1992); Götzenbrühl,
Germany (Hegner et al., 1995). Nephelinites: AS-002, Chyulu Hills, Kenya (Späth et al., 2001); GGU 421301-1, Nunatak, East Greenland
(Bernstein et al., 2000). Meimechite: G3-100, Maymecha, Russia (Arndt et al., 1998). Nuanetsi picrite: N163 (Ellam & Cox, 1989, 1991). Basanite:
BR-11, Barrington, East Australia (O’Reilly & Zhang, 1995). Mantle amphibole: xenolith MG91-143.4, 34080 m (Moine et al., 2001).
Normalization values from McDonough & Sun (1995).
Africa (Hawkesworth et al., 1990) and in metasomatic
oxide (mathiasite) from such xenoliths (Fig. 14; Griffin
et al., 1999).
All the samples analysed for Os isotopes have low to
very low 187Os/188Os ratios (0·1342–0·1067, Table 6) and
all except the Type 1a sample have negative Os (Fig.
15). There is an inverse correlation between the Os
isotope ratios and the amounts of Os present in the
samples (Table 6). One sample of Type 0 has exceptionally high Os, 44 ppb. These data are as unusual
as the other isotope data: 187Os/188Os ratios below 0·110
have previously not been reported from igneous rocks
but only from peridotite xenoliths from old subcontinental
We do not know of any other mantle-derived uncontaminated igneous rocks with Pb isotope ratios similar
to those of the Manı̂tdlat Member. Many alkaline rocks
have similarly high 206Pb/204Pb ratios, but for a given
206
Pb/204Pb they all have higher 207Pb/204Pb and 208Pb/
204
Pb ratios, most of which lie close to or above the
NHRL (e.g. Nelson et al., 1988; Simonetti & Bell, 1994,
1995; Hegner et al., 1995; Wilson et al., 1995; Kalt et al.,
1997; Le Roex & Lanyon, 1998; Bell & Tilton, 2001;
Späth et al., 2001, and compilation therein). Pb isotope
ratios similar to those of the Manı̂tdlat Member are,
however, reported from some highly metasomatized peridotite xenoliths with metasomatic oxides from South
23
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
Fig. 13. (a) Sr–Nd isotope compositions at 60 Ma of the alkaline rocks
of the Manı̂tdlat Member and contemporaneous tholeiites of the
Ordlingassoq Member. The mixing curve is for hypothetical mixing
between tholeiitic and Type 0 melts, and the numbers denote the
fraction of alkaline melt in the mix. Data for the tholeiites of the
Ordlingassoq Member are from Table 4, Holm et al. (1993), Lightfoot
et al. (1997) and Graham et al. (1998). (b) Comparison with volcanic
rocks from other provinces and mantle and crustal components. Areas
of continental volcanic rocks spread over most of the lower right
quadrant; see, for example, compilation by Zindler & Hart (1986).
Elkhead Mts from Leat et al. (1988) and Thompson et al. (1989), Leucite
Hills from Vollmer et al. (1984), African carbonatites from Bell &
Blenkinsop (1989), Iceland after Stecher et al. (1999), OIB after Hofmann
(1997), and the mantle components after Hart (1988).
Fig. 12. Incompatible element ratios in Manı̂tdlat Member rocks
compared with melilitites, nephelinites, meimechites, alkali picrites and
basanites. The comparison data are shown as one group (‘Others’),
and the central area of this group is outlined for clarity. Data sources:
Brooks et al. (1979), Anthony et al. (1989), Rogers et al. (1992), Arndt
et al. (1995, 1998), Hegner et al. (1995), O’Reilly & Zhang (1995),
Wilson et al. (1995), Bernstein et al. (2000), Mahotkin et al. (2000), Späth
et al., (2001).
Lightfoot et al., 1997). Contamination with either Precambrian basement or Mesozoic sediments leads to increase in SiO2, K2O, Rb and Pb, and decrease in FeO∗,
CaO, TiO2 and Nb (but not appreciably Zr). Sr isotope
ratios increase, and Nd and Pb isotope ratios decrease.
Compositional shifts of this kind are seen in the three
high-Si Stordal samples, one of which is analysed isotopically, and to a lesser extent in four Type 1a samples
(Figs 6, 7, 9, 13 and 14). We therefore consider these
samples to represent crustally contaminated alkaline
magmas. Modelling of the crustal contamination process
is outside the scope of the present paper.
On the other hand, the character of the alkaline rocks
as such, with their strong enrichment in Ba, U, Nb,
lithospheric mantle (Pearson et al., 1995; Hanghøj et al.,
2001).
DISCUSSION
Modification of the primary alkaline
magmas
Identification of crustal contamination
The effects of crustal contamination on the tholeiitic
volcanic rocks on Disko and Nuussuaq are well described
(Pedersen, 1979, 1985a; Pedersen & Pedersen, 1987;
Goodrich & Patchett, 1991; Pedersen et al., 1996;
24
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
batches and their phenocrysts, of which the tholeiites
contained only olivine + chromite. The heterogeneous
olivine populations within single samples suggest that
such mixing has taken place (Fig. 3). Low-CaO zoned
olivines such as phenocryst 1 in Fig. 3 are explicable as
crystals that originally formed in tholeiitic magmas and
later were mixed into the alkaline magmas where they
partly re-equilibrated and continued their growth. The
‘tholeiitic’ levels of Ca, Cr and Ni in the cores are
preserved, whereas Mg and Fe are re-equilibrated to
lower mg-numbers. The low CaO/Al2O3 in the glass
inclusions in the low-CaO olivines (Fig. 5) confirms the
tholeiitic character of their parental melts. These low-Ca
olivines have mg-numbers of 82–92·3 (Fig. 3). Phenocrysts
2–4 in Fig. 3 are normal phenocrysts in the alkaline
magma.
Because of the evidence for mixing it is necessary to
evaluate the extent to which this process has modified
the composition of the original alkaline magmas. The
main constraint on this comes from the near-constant
Nd isotope ratios in Types 0, 1b and 3; the total variation
in 143Nd/144Ndi for these rocks is only 0·00002 (0·51236–
0·51238, Table 5). Because the Nd contents are strongly
dominated by the alkaline component, moderate amounts
of tholeiitic magma can be mixed into the alkaline magma
before the Nd isotope ratio changes significantly. A Type
0 alkaline end-member magma may be mixed with up
to 25% tholeiitic magma before the Nd isotope ratio
increases by more than 0·00002 (mixing curve in Fig.
13), and 25% is therefore considered a maximum amount
of in-mixed tholeiitic magma. We have no means of
quantitatively constraining the amounts of tholeiitic
magma further, but based on the relative scarcity of the
tholeiitic olivine crystals we consider that the amount of
tholeiite in the alkaline magmas was normally 10% or less.
It is possible that the tholeiitic component is dominated by
olivine crystals picked up in the mush zones in the conduit
systems, with very little accompanying tholeiitic melt.
With an upper limit of 25% tholeiitic magma in the
alkaline magmas, the major-element composition of the
unknown pure alkaline end-member melt will not be
very different from that of the erupted magmas. SiO2 in
the end-member will be lower than in the erupted rocks
by 0·5 wt % or less; CaO/Al2O3 will be higher but
still <1·26. TiO2 and Na2O are invariably low. The
incompatible trace elements in the alkaline end-member
will be higher by a factor of 1·3 or less; ratios of
more incompatible elements (e.g. Nb/La, Th/U) will be
virtually unaffected, whereas ratios of more incompatible
to less incompatible trace elements (e.g. Nb/Y, Ba/Ti)
will be lowered by the tholeiitic component. With an
upper limit of 10% in-mixed tholeiite the alkaline magmas
are practically unchanged. In conclusion, there is undoubtedly a small tholeiitic component in the alkaline
Fig. 14. Pb isotope compositions of the alkaline rocks of the Manı̂tdlat
Member and contemporaneous tholeiites. Data as measured; short
lines at data points show the size of a 60 Ma age correction. For the
tholeiites, the correction is the size of the symbol or less. Data for the
tholeiites of the Ordlingassoq Member are from Table 4, Lightfoot et
al. (1997), Graham et al. (1998) and unpublished data (2002). NHRL
is the Northern Hemisphere Reference Line (Hart, 1984). The highSi Stordal sample is labelled ‘cont.’. The continental crust in West
Greenland has low Pb isotope ratios and the main data fields fall
outside the diagram areas to the lower left (Kalsbeek et al., 1988;
Kalsbeek & Taylor, 1999; compilation by Lightfoot et al., 1997). The
Iceland fields are after Stecher et al. (1999), the OIB fields after Hofmann
(1997), and the mantle components after Hart (1988). Alkaline volcanic
rocks after compilation by Späth et al. (2001), Elkhead Mts from
Thompson et al. (1989), and mathiasite (×) from Griffin et al. (1999).
LREE and P, and low Sr isotope ratios and high Pb
isotope ratios, cannot be explained in terms of contamination with any known crustal components in the
region.
Evidence for mixing of tholeiitic magma into alkaline
magma
The alkaline magmas were erupted within a regional
zone of eruption centres that mainly produced tholeiitic
magmas, and it is likely that the magmas utilized the
same conduit systems. The conduits would be possible
sites of mixing between alkaline and tholeiitic magma
25
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
Table 6: Os isotope analyses of Manı̂tdlat Member alkaline rocks, West Greenland
Sample:
264167
327097
327099
326787
264124
264113
264093
7649
7818
Type:
Stordal
0
0
1a
1b
2
3
Tholeiite∗
Tholeiite∗
( 187Re/188Os)m
1·508
0·028
7·080
1·917
2·645
1·868
0·295
0·814
0·505
( 187Os/188Os)m
0·12225
0·10850
0·12494
0·13416
0·11504
0·12570
0·10673
0·13178
0·13279
Re (ppb)
0·309
0·260
0·792
0·105
0·621
0·326
0·382
0·269
0·199
Os (ppb)
0·973
44·278
0·53
0·258
1·112
0·826
6·140
1·594
1·899
( 187Os/188Os)60
0·12074
0·11785
0·13224
0·11240
0·12383
0·10643
0·13097
0·13228
0·10847
Os60
−4·64
−14·33
−2·20
−15·94
T-RD (Ma)
927
2706
1351
−6·92
−788
2143
−11·23
472
2996
−595
−795
T-MD (Ma)
−258
2898
−19
283
−321
−53
10401
692
3278
4·45
3·44
4·48
∗Analyses of West Greenland Ordlingassoq Member tholeiites for comparison, from Schaefer et al. (2000).
m, as measured; 60, age corrected to 60 Ma. T-RD, Re depletion ages; T-MD, mantle depletion ages (Shirey & Walker, 1998).
(Fig. 13). The low Rb and normalized Ta/La>1 suggest
a relation to Type 1a, and the dyke is actually situated
in an area where the alkaline rocks are solely of Type
1a. Simple mixing calculations between tholeiite (sample
326783) and Type 1a (sample 326787) give consistent
results for most trace elements and the Nd isotopes,
suggesting that the dyke is a tholeiitic magma that contains
around 15% alkaline component.
Primary alkaline magmas
The composition of the most magnesian, possibly primary, alkaline magmas may be estimated from the most
magnesian cognate olivines present. The most magnesian
high-CaO olivine has mg-number 90 (Fig. 5c), which
corresponds to a melt calculated to have around 15 wt %
MgO, somewhat dependent on the oxidation state. Thus,
samples with >15 wt % MgO most probably contain
accumulated olivines whereas samples with lower MgO
may represent the erupted and more or less fractionated
magmas. At 15 wt % MgO the tholeiitic melts had
temperatures close to 1400°C (Larsen & Pedersen, 2000).
The parental alkaline melts would have had lower temperatures, loosely estimated around 1300°C.
Fig. 15. Os isotopic compositions at 60 Ma of the alkaline rocks of
the Manı̂tdlat Member and contemporaneous tholeiites. Fields of the
various mantle components and plume melts from Shirey & Walker
(1998). Data for the tholeiites of the Ordlingassoq Member from
Schaefer et al. (2000) and D. G. Pearson (unpublished data, 2002).
magmas, but the compositional influence of this is negligible for most elements.
The dyke: mixing of alkaline magma into tholeiitic magma
Whereas the composition of an alkaline magma is fairly
robust against addition of minor amounts of tholeiitic
magma, the opposite is not the case. This is illustrated
by the investigated dyke that cuts the entire Manı̂tdlat
Member succession. The dyke can be interpreted in
terms of mixing of alkaline magma into tholeiitic. Its
major-element composition is tholeiitic (Fig. 6), whereas
its incompatible trace elements (including K and P)
are intermediate between tholeiitic and alkaline values,
higher than those in the tholeiites by a factor of 2–4 (Fig.
11). The Nd–Sr isotope composition of the dyke is also
intermediate between that of tholeiitic and alkaline rocks
Mantle sources and melting processes
The rocks of the Manı̂tdlat Member have incompatible
trace element concentrations that are enriched by up to
100–300 times primitive mantle for elements such as Ba,
U, Nb and La (Fig. 11). Similar enrichment levels in
basic igneous rocks are normally found in melilitites,
nephelinites and meimechites, and also some kimberlites.
These rock types are all strongly silica undersaturated,
sometimes larnite normative, plagioclase free or plagioclase poor, and are considered to be formed by small
26
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
enriched alkali picrites (Milholland & Presnall, 1998), but
they will not be like the alkali-poor Manı̂tdlat Member
melts. Dilution of an extreme small-degree alkaline melt
with tholeiitic melt is another possibility, but this requires
large amounts of tholeiitic melt (>80%), which, as discussed above, is very unlikely. A more straightforward
explanation is that the melting involved unusual mineral
assemblages in enriched and metasomatized parts of the
mantle.
Plume mantle or subcontinental lithospheric mantle?
In continental areas, the enriched mantle sources for
basic alkaline rocks have been envisaged to be situated
in the subcontinental lithospheric mantle, although the
Nd–Sr isotopes often indicate that the lithosphere was
metasomatized by asthenospheric melts shortly before the
melting (e.g. Wilson, 1989). In oceanic islands, enriched
mantle components such as EM1 and EM2 are considered to reside within the mantle plume that brings
them to the surface and into the melting regime (e.g.
Hofmann, 1997). The idea of melting of enriched plume
components has also been applied to some continental
basic alkaline rocks, particularly the meimechites of Siberia (Arndt et al., 1995, 1998), and the melilitites and
nephelinites of central East Greenland (Bernstein et al.,
2000, 2001). Indications of the location of the enriched
mantle source for the rocks of the Manı̂tdlat Member
come from the isotope compositions and the field relations.
The isotope compositions of the Manı̂tdlat Member
rocks of Types 0, 1, 3 and Stordal plot outside the fields
of the known oceanic mantle components for all of the
isotope systems investigated (Figs 13–15). The few other
rocks that also plot within the lower left quadrant of the
Nd–Sr isotope diagram are interpreted to be either melts
from enriched subcontinental lithosphere, or asthenospheric melts that have reacted with such lithosphere
(Thompson et al., 1989; Simonetti & Bell, 1995; Mahotkin
et al., 2000). As noted above, the Pb and Os isotope data
are comparable only with data from peridotite xenoliths
representing the subcontinental lithospheric mantle. The
melilitites and nephelinites from East Greenland that are
suggested to be derived from an enriched component
intrinsic to the Iceland mantle plume have Nd–Sr isotopes
that plot on the mantle array in extension of the Iceland
field, and Pb isotopes that plot within to slightly above
the Iceland fields (Bernstein et al., 2001, and unpublished
data, 2002). If the enriched mantle source for the
Manı̂tdlat Member was an intrinsic part of the protoIcelandic plume, then it is a new and unique component
of that plume.
The geological setting of the Manı̂tdlat Member is
significant in this context. The alkaline magmas were
generated only during a brief period in the middle of a
Fig. 16. Al2O3 and total alkalis in the Manı̂tdlat Member rocks
compared with melilitites, nephelinites, meimechites, alkali picrites
and basanites (‘Others’). Also shown is a suite of camptonitic and
monchiquitic (chemically, nephelinitic) dykes from Ubekendt Ejland
>100 km north of Disko (Fig. 1). These dykes have low TiO2 and
P2O5 and high Al2O3, comparable with the Manı̂tdlat Member, but
they have high alkalis. Data sources for ‘Others’ as in Fig. 12. Ubekendt
Ejland: Larsen (1981, 1982), Clarke et al. (1983). Analyses are not
recalculated volatile-free because of the variable and often high contents
of primary volatiles.
degrees of melting of enriched, volatile-bearing mantle
(e.g. Nelson et al., 1988; Wilson, 1989; Rogers et al., 1992;
Taylor et al., 1994; Arndt et al., 1995, 1998; Hegner et
al., 1995; Wilson et al., 1995; Mahotkin et al., 2000). In
comparison with these rocks, those of the Manı̂tdlat
Member are not highly undersaturated and have relatively high contents of SiO2 and Al2O3, and low TiO2,
alkalis and P2O5 (Fig. 16). The effusive eruption style
and the anhydrous mineralogy do not suggest that the
volatile contents of the Manı̂tdlat Member magmas were
higher than in the tholeiites.
Manı̂tdlat Member melts such as those of Types 0 and
1 have CaO/Al2O3 around 1·2 and, at the 15 wt %
MgO level, >11·5 wt % Al2O3 and >1·0 wt % Na2O.
According to the melting model of Herzberg & Zhang
(1996), such melts cannot be produced near the solidus
of an ordinary dry lherzolite at any pressure. Olivine
fractionation from a deep near-solidus melt may produce
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Metasomatized mantle lithologies
much longer period of formation of high-degree melts
from the plume (the tholeiitic picrites), and they were
erupted within only a 20 km by 30 km area. If the enriched
source was a localized, unique component within the
upwelling plume, being the most fusible part it would
have started melting at deeper levels than the depleted
surroundings. This requires that the alkaline melts ascended through the overlying melting zone producing
the tholeiites (and through the lithosphere) while retaining
their identity without being mixed into the contemporaneous tholeiitic melts. This is contrary to the
concept of pooling of melts in the melting column.
On the other hand, if the enriched mantle source
was a limited volume situated within the subcontinental
lithospheric mantle beneath northern Disko, the particular setting of the Manı̂tdlat Member volcanic rocks
is easily explicable. The regional tholeiitic feeder systems
progressed with time from NW to SE, and only when
they moved to the Stordal–Maniillat area (Fig. 1) did
the conduits start to traverse the enriched parts of the
lithosphere. The passing hot tholeiitic picrites provided
the necessary heat for the melting of the enriched lithosphere material, and when the low-melting parts had
been removed the production of alkaline magma ceased
whereas the tholeiitic magma production from the plume
continued.
Minerals known from metasomatized mantle xenoliths
are (besides olivine) clinopyroxene, amphibole, phlogopite, apatite, carbonate, ilmenite, rutile and a number
of exotic oxides. Orthopyroxene and spinel or garnet are
often rare or absent because they have reacted to form
clinopyroxene, amphibole or mica (O’Reilly & Griffin,
1988; Yaxley et al., 1991, 1998; Van Achterbergh et al.,
2001). If the metasomatizing agent was carbonate rich,
much of the CO2 released by the decarbonation reactions
will subsequently have been lost (Yaxley et al., 1991,
1998).
Melting of clinopyroxene-rich lithologies has been invoked to explain the high CaO and CaO/Al2O3 of
nephelinites and melilitites (e.g. Francis & Ludden, 1990;
Francis, 1991; Gibson et al., 1999). Clinopyroxenes in
metasomatized mantle xenoliths may contain up to 100–
600 ppm Sr, 10–100 ppm La and 20–200 ppm Ce
(Erlank et al., 1987; O’Reilly et al., 1991; Hauri et al.,
1993; Ionov et al., 1997; Ionov, 1998; Zanetti et al., 1999;
Laurora et al., 2001; Van Achterberg et al., 2001) and
may thus yield a significant contribution to the Sr and
REE budget of the melt; however, elements such as Ba
and Nb do not exceed 1–2 ppm in clinopyroxene, which
even at 1% melting will not yield the large degrees of
enrichment seen in the Manı̂tdlat Member.
Amphibole in metasomatized mantle xenoliths is often
pargasitic. Mantle pargasite contains 0·4–2·0 wt % K2O,
typically around 1 wt % K2O (O’Reilly & Griffin, 1988;
O’Reilly et al., 1991; Ionov & Hofmann, 1995; Chazot
et al., 1996; Ionov et al., 1997; Ionov, 1998; Zanetti et al.,
1999; Laurora et al., 2001; Moine et al., 2001). Therefore,
despite the presence of deep K troughs in the trace
element patterns of most Manı̂tdlat Member rocks (Fig.
11) their source may have contained significant proportions of pargasitic amphibole. Amphiboles are capable
of accommodating higher amounts of incompatible trace
elements than clinopyroxene, particularly Nb and Ta
(see the references cited above), and multi-element spectra
of mantle pargasites show distinct peaks at Nb–Ta (Ionov
& Hofmann, 1995; Laurora et al., 2001; Moine et al.,
2001), rather similar to the patterns of the Manı̂tdlat
Member Type 2 melts (Fig. 11). A large fraction of
pargasite in the melting assemblage will lead to high Al
contents in the melts, as seen in the Manı̂tdlat Member.
Phlogopite contains about 8 wt % K2O and has high
concentrations of Rb and Ba. The deep troughs for Rb
and K in the trace element patterns of all the Manı̂tdlat
Member melts except those of Type 2 suggest that no
or only little phlogopite contributed to the melts. The
huge amounts of Ba must be provided by other phases.
Mantle apatite may accommodate high amounts of
Th, U, Sr and REE (O’Reilly et al., 1991; Chazot et al.,
1996; Ionov et al., 1996, 1997; O’Reilly & Griffin, 2000),
Age of the enriched components
It is clear from the radiogenic isotope compositions of
the Manı̂tdlat Member rocks that the ultimate derivation
of the enriched source components from a depleted
asthenospheric mantle cannot be related to the 60 Ma
melting event but must be older. As a consequence of
radiogenic ingrowth in the enriched source the isotope
ratios of the Manı̂tdlat Member rocks at 60 Ma correlate
broadly with the appropriate parent/daughter ratios; in
particular, the Type 2 rocks, which have higher Sm/Nd
and Rb/Sr and lower U/Pb than the other types, also
have higher 143Nd/144Nd and 87Sr/86Sr and lower 206Pb/
204
Pb (Figs 11, 13 and 14). The age of the enrichment
is, however, very uncertain. If the Type 2 source at some
initial stage had Nd, Sr and Pb isotope ratios similar to
the sources for the other magma types this could have
been at around 300–500 Ma. Assuming that the Sm/
Nd ratios were not significantly changed during the
melting event at 60 Ma, the Nd isotopes give Nd extraction ages from a depleted reservoir of 1·1–1·4 Ga.
The oldest age indication comes from the Os isotopes,
where the Os in the two samples with 187Os/188Os =
0·106–0·108 must have evolved in a practically Re-free
environment for nearly 3 Gyr (Table 5); this is, however,
a depletion age and could represent melting of sulphides
in the ancient side wall to the enriched areas.
28
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
but apatite does not fractionate Th from U to the degree
seen in the Type 0, 1 and 3 melts (Fig. 11). Moreover,
Types 1a and 1b have similar concentrations of P2O5
but very different levels of enrichment in Sr and REE
(Figs 6 and 11), suggesting that an additional phase
delivered these elements to the Type 1b melts.
Metasomatic oxides such as the Ba–K titanates lindsleyite and mathiasite (LIMA phases) have very high
concentrations of Ba, Sr, LREE, Nb, Zr, Pb, U and Th,
very high U/Th ratios, and Pb isotope ratios similar to
those of some Manı̂tdlat Member melts ( Jones, 1989;
review by Haggerty, 1991; Griffin et al., 1999). A LIMA
phase in the melting assemblage would explain some of
the unusual trace element ratios such as the high U/Th,
and also the unusual Pb isotopes of the Manı̂tdlat Member
Type 0, 1 and 3 melts.
The lower contents of Ni in the Manı̂tdlat Member
rocks compared with tholeiites with similar MgO contents
(Fig. 6) suggest that olivine and orthopyroxene constituted
a smaller part of the melting assemblage than in an
ordinary lherzolite, and this is consistent with a metasomatized source.
In conclusion, the most probable mantle sources for the
Manı̂tdlat Member melts are amphibole–clinopyroxenerich lithologies with apatite and, for Types 0, 1 and 3, a
LIMA phase. Mantle xenoliths with such lithologies (but
without LIMA) have been described from eastern Australia by Wass (1979), O’Reilly & Griffin (1988) and
O’Reilly et al. (1991), and have been invoked as source
for the nephelinites and basanites in that region (O’Reilly
& Zhang, 1995), which show several similarities to the
Manı̂tdlat Member rocks, namely, high Al2O3 and low
TiO2, P2O5, Rb/Sr and La/Nb. In the Australian xenoliths, amphibole and apatite occur both disseminated
through lherzolite and concentrated in centimetre-sized
veins.
enriched Type 2 melts could have resulted from higher
degrees of melting than the other alkaline melts, but
then the phases such as LIMA that contributed to the
peculiarities of the other melts should also be discernible
in the Type 2 melt, which they are not. In all, we consider
that differences in the melting assemblages are far more
important than variations in the degree of melting for
production of the compositional spread observed.
Many basic alkaline rocks have deep K troughs in
their trace element spectra, and this has been interpreted
as a result of the presence of a residual potassic phase
during melting, either phlogopite (Foley & Wheller, 1990;
Rogers et al., 1992) or amphibole (Späth et al., 2001).
This interpretation is necessitated by the assumption that
the enriched mantle source for the melts had a smooth
trace element spectrum with no K trough. However, this
assumption may in many cases be unwarranted; indeed
the trace-element spectra of enriched mantle xenoliths
are usually far from smooth and very often have deep
troughs at K and sometimes at Zr and Ti (Fig. 11;
Menzies et al., 1987; O’Reilly & Griffin, 1988; Yaxley et
al., 1991; Hauri et al., 1993; Ionov et al., 1997; Gorring
& Kay, 2000). The incompatible trace elements reside
mainly in the metasomatic phases, which are often volatile-bearing and easily fusible. During melting of metasomatized mantle the metasomatic phases will melt
preferentially, transferring the incompatible trace elements and thereby the spiky pattern to the melt, and
leaving the depleted residual mantle with a smoothed
trace element pattern and, in many cases, no residual
phlogopite or amphibole but only the breakdown products from incongruent melting of these minerals. We
consider that the spiky trace element spectra of the
Manı̂tdlat Member magmas were inherited in this way
from similar spectra of the enriched source.
Efficient preferential melting of metasomatic phases
was demonstrated experimentally by Foley et al. (1999)
for mantle assemblages with amphibole, apatite, clinopyroxene, mica and ilmenite. At 15 kbar, the solidus for
pargasite-bearing assemblages is in the range
1050–1075°C, and amphibole and oxide melt completely
within a few tens of degrees above the solidus. Apatite
survives to slightly higher temperatures, whereas mica
melts over a larger temperature range up to 1170°C. In
West Greenland, the tholeiitic melts originated in the
asthenosphere with temperatures of >1550°C; within
the lithosphere their temperatures dropped to
1400–1300°C for melts with 15–12 wt % MgO (Larsen
& Pedersen, 2000). We envisage that these melts provided
sufficient heat rapidly enough to raise the temperatures
in the adjacent lithosphere to 1100–1170°C, thereby
inducing complete melting of amphibole, oxide, apatite
and mica. The restite then consisted of clinopyroxene +
olivine ± orthopyroxene and would have a sharply raised
melting temperature. Farther from the heat source the
Melting conditions and residual phases
Pargasitic amphibole is stable to >1050°C and 30 kbar,
>96 km (Gilbert et al., 1982). In South Africa, LIMA
phases were formed together with richteritic amphibole
in the garnet stability field at estimated depths of 75–100
km (Haggerty, 1991). At 60 Ma, just before breakup, the lithosphere in West Greenland was strongly
attenuated and only around 100 km thick (Herzberg
& O’Hara, 1998). The metasomatized mantle volumes
beneath West Greenland were probably situated at deep
levels in the attenuated lithosphere but precise depth
estimates cannot be made.
It is virtually impossibe to detect variations in the
degree of melting between types because all the usual
chemical indicators, such as CaO/Al2O3 and Ce/Y,
reflect the source enrichment more than the degree of
melting. If phlogopite was a refractory phase the less
29
JOURNAL OF PETROLOGY
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Table 7: Fractions (wt) of minerals in the melting assemblages for the Manı̂tdlat Member
Amphibole
Mica
Apatite
Cpx
Oxide
Carb?
Olivine
Sum
Type 2
0·60
0·04
0·0074
0·27
0·09
1·01
Stordal
0·60
0·0087
0·26
opx 0·13
1·00
Type 1a
0·60
0·0112
0·30
0·005
0·10
1·02
Type 1b
0·60
0·0114
0·30
0·005
0·040
0·10
1·06
Type 0
0·60
0·0128
0·30
0·003
0·010
0·09
1·02
Type 3
0·60
0·0174
0·30
0·007
0·010
0·09
1·02
The sums of the fractions are not adjusted to unity because of the uncertainty of the method; for example, the carbonate
may as well be smaller amounts of an unknown phase. The oxide is probably one or both of the Ba-titanates lindsleyite
and hawthorneite (LIHA). The Stordal type has orthopyroxene instead of olivine.
metasomatic minerals may not have melted completely,
but it is conceivable that melts from such areas would
make only a small contribution to the total melt volume
produced.
different from those of the other types. The Stordal type
can be modelled only with low clinopyroxene (26%), and
orthopyroxene instead of olivine because of the relatively
high SiO2. All the Stordal magmas may be crustally
contaminated to some extent.
For the incompatible trace elements, the known inventory of these in the primary alkaline melts was distributed between the contributing phases as detailed in
the Appendix. It must be stressed that these calculations
do not provide unique solutions. They characterize possible and reasonable solutions and highlight the differences between the various melt types involved, but other
solutions are not thereby invalidated.
As shown in Fig. 17, amphibole contains a significant
part of the trace-element inventory (and K), particularly
Nb, Ta, Zr and Hf. The clinopyroxene contribution is
small and that of olivine is, by our definition, nil.
A LIMA phase is required in the melting assemblages
for Types 0, 1a, 1b and 3, and it must be one or both
of the Ba end-members lindsleyite, Ba(Ti,Cr,Fe,
Zr,Mg)21O38, or hawthorneite, Ba(Ti,Cr,Fe,Mg)12O19
(LIHA). These minerals have very high concentrations
of Ba, U, Nb, Ta, REE and Sr (e.g. Jones, 1989; Haggerty,
1991; Griffin et al., 1999). They also have high TiO2
(20–60 wt %), but because only a small fraction of the
mineral is required the contribution to the total Ti
budget is small compared with those from amphibole
and clinopyroxene. The required LIHA phase cannot be
Zr rich, which could suggest it is hawthorneite. The
LIHA phase is the main phase responsible for the extreme
Th–U fractionation and the high Pb isotope ratios in the
Type 0, 1 and 3 rocks.
In the Type 2 melting assemblage, mica is required
by both the major and trace element constraints and
contributes significant amounts of Rb, Ba and K to the
melt. The amphibole is almost solely responsible for the
large Nb–Ta peak (Fig. 17). Apatite is the main phase
Melting assemblages for generation of the Manı̂tdlat
Member melts
Quantification of the melting process is difficult because
of the number of unknowns involved. The source may
be veined, and both veins and wall rocks would melt but
to very different extents (Foley, 1992; Foley et al., 1999),
making parameters such as the degree of melting complex
if not meaningless. But if we take the existing melts
as starting points, viewing them as the sums of the
contributions from the various melting minerals, we can
calculate the bulk melting mineral assemblages. When
we make the simplifying assumption, justified above, of
complete melting of the enriched phases except clinopyroxene, simple mass-balance equations can be used in
the calculations of the melt compositions. The details of
these calculations are described in the Appendix and
given in the Electronic Appendix B; the modelled melting
assemblages are shown in Table 7, and Fig. 17 shows
the incompatible trace element inventory of the various
Manı̂tdlat Member magma types distributed between the
contributing melting phases.
For the major elements, combinations of analysed
mantle minerals were used in mixing equations to approximate the major-element composition of the primary
alkaline melts with 15 wt % MgO. This worked surprisingly well and showed that melts calculated as consisting of roughly 60 wt % amphibole, 30 wt %
clinopyroxene, 10 wt % olivine and 1 wt % apatite are
fairly close to the major-element compositions of the
Manı̂tdlat Member melts. In detail, the various melt
types require slightly different melting assemblages. Type
2 requires small amounts of mica (>4%), less clinopyroxene (27%), and minerals with compositions slightly
30
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 17. Distribution of the trace element inventory of the six types of melt in the Manı̂tdlat Member between the contributing melting phases.
The spectra are normalized to primitive mantle abundances as in Fig. 11, but the scale is linear to show the correct proportions between the
phases. The contribution from each phase is stacked on top of the other in the succession shown in the legends, so that the sum of the
contributions equals the content in the melt. (For details of calculation, see text.) The melting assemblages also comprise 9–13% olivine and
orthopyroxene (see Table 7). It should be noted that amphibole and apatite in Type 2 are compositionally different from those in the other
types.
responsible for the contents of Th and U (and P), whereas
amphibole and apatite contribute equally to the REE.
In the Stordal melting assemblage, Th and U come
mainly from apatite whereas the REE and Sr are shared
between apatite and amphibole, with a smaller contribution from clinopyroxene. The apatite has higher Th,
U and REE than the apatite in Type 2.
For the Type 1a and 1b melts, the modelled melting
assemblages are in the first round closely similar. However, compared with Type 1a, Type 1b is significantly
more enriched in Ba, Sr, Pb and LREE–MREE, but not
in U, Nb, Ta and Ti (Fig. 11). Thus, either the oxide
phase has another composition or an additional phase is
involved, which concentrates Ba, Sr, Pb and LREE–
MREE. The most Ba rich of the metasomatic oxides is
hawthorneite, with c. 13 wt % BaO, which occurs in
association with lindsleyite (Erlank et al., 1987; Haggerty
et al., 1989). It is possible that the oxide in Type 1a is
lindsleyite and in Type 1b is hawthorneite, but this makes
the similar U–Nb–Ta and Ti in the two types fortuitous
and is difficult to model satisfactorily. Carbonates with
high contents of Ba, Sr and LREE–MREE have been
reported from carbonatites (Knudsen, 1991; HornigKjarsgaard, 1998), in metasomatized mantle xenoliths
( Jones et al., 2000), in melt inclusions in lherzolites (Van
Achterbergh et al., 1999), and in doped experiments
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JOURNAL OF PETROLOGY
VOLUME 44
(Veksler et al., 1998). We have tentatively assigned 4 wt %
carbonate with high carbonate–amphibole partition coefficients for Ba, Sr, Pb and REE to the Type 1b melt.
A third possibility is a yet unknown phase, in which case
much smaller amounts may be required.
The Type 0 and Type 3 melting assemblages are
closely related to that of Type 1b but require slightly
different proportions of the minor phases (Table 7).
NUMBER 1
JANUARY 2003
mixing of the melts. Type 1a is the volumetrically dominant melt type; this suggests that the corresponding
metasomatic zone is the widest of the successive zones.
The volume of the Manı̂tdlat Member alkaline melts
is at least 30 km3, which requires that the metasomatic
zones are much wider than the few decimetres envisaged
by O’Reilly et al. (1991), and that the original intrusion
was on a larger scale than just ‘veins’. In a rough
calculation, 30 km3 of primary melt produced by on
average 10% melting corresponds to 300 km3 metasomatized mantle. The length of the feeder systems for
the Manı̂tdlat Member, considered to be NE-directed
fissures, is c. 30 km (Fig. 1). Thus, if the metasomatized
areas were 30 km in length and had a depth extension
of, for example, 10 km, then the total width of the
metasomatic zones must be around 1 km to encompass
300 km3 of material. This allows room for the contemporaneous tholeiitic melts that were erupted within
the area covered by the Manı̂tdlat Member lavas. It
does, however, involve the corollary that a substantial
amount of crystallized mafic melt must be situated at
depth in the lithosphere; in eastern Australia xenolith
data support such a situation (O’Reilly et al., 1988).
Structure of the metasomatized mantle
As discussed above, a number of metasomatized source
areas with different mineralogies must have been present
in the lithospheric mantle to give rise to the various magma
types of the Manı̂tdlat Member. The source areas were
present within the same limited part of the lithosphere
beneath northern Disko, and only there, and they melted
simultaneously. Although separate they must have been
closely associated, and most probably they were genetically
related. The existence of the magma types suggests an
ordered structure of the metasomatized mantle domains
rather than a random medley of veins and patches.
O’Reilly & Griffin (1988) and O’Reilly et al. (1991)
suggested that metasomatism in the lithospheric mantle is
caused by volatile components released from crystallizing
veins of alkaline basic magma (ultimately asthenosphere
derived). The released volatile components will be rich
in incompatible elements, which migrate through the
mantle sidewall and precipitate in a sequence depending
on the mineral stabilities and the mineral–fluid partition
coefficients, as in a chromatographic column. In particular, mica will form only close to the veins, trapping
Rb, K and Ti, and thereby depleting these elements in
the migrating fluid, which becomes more carbonate rich
with distance from the vein. Elements such as Ba, Th,
U and LREE will move farther before they are precipitated, e.g. in apatite. The result is a number of
metasomatic zones.
If the metasomatized mantle source for the Manı̂tdlat
Member magmas had a similar zoned structure, both
the occurrence of the different magma types and the
complementary relationship between Type 2 and Type
1b seen in Fig. 12 would be explicable. Figure 18 shows
the suggested model, which is slightly modified from that
of O’Reilly et al. (1991) with regard to the successions of
amphibole and apatite. We also assume precipitation of
an oxide phase, and an additional stage with oxide +
carbonate or an unknown phase. The Type 2 melts
would then be derived from vein-near micaceous mantle,
the Stordal and Type 1a melts from intermediate zones,
and Type 0, 1b and 3 melts would be derived from the
zones farthest away. The magma conduits must cross all
the zones, and the zones must extend vertically to avoid
Metasomatized mantle in West Greenland
The Precambrian basement in the whole of West Greenland is intruded by several occurrences of small-volume
ultramafic–alkaline magmas such as lamprophyres, kimberlites and carbonatites, with ages ranging from Archaean to Eocene (Larsen & Rex, 1992). Thus, there is
abundant evidence that the lithosphere contains metasomatized domains, but the extent of these is not known
because their possible surface expressions, the alkaline
magmas, are strongly controlled by the tectonic state of
the lithosphere. The Manı̂tdlat Member itself is an example that metasomatized areas of the lithosphere may
lie ‘dormant’ with no magmatic surface expression for
hundreds of millions of years until a new tectonic regime
allows melt generation and ascent to the surface.
The basement within the Nuussuaq Basin and in the
stable areas east of it consists of Archaean crust of >2800
Ma age that was reworked during the Proterozoic around
1750 Ma (Kalsbeek, 1999). The region hosts two occurrences of alkaline rocks in addition to the Manı̂tdlat
Member. The basement at Eqi (Fig. 1) contains a small
swarm of weakly deformed, Palaeoproterozoic ([1780
Ma) carbonate-rich, ultramafic lamprophyre dykes
(Larsen & Rex, 1992). On Ubekendt Ejland (Fig. 1), a
small swarm of lamprophyre dykes dated at >34 Ma
cut the >20 Myr older volcanic succession (Larsen, 1981,
1982; Clarke et al., 1983). The dykes are strongly enriched
in incompatible trace elements, alkalis (up to 6·3 wt %,
Fig. 16) and volatiles (5–9 wt % H2O + CO2), and they
may be small-degree melts of metasomatized mantle
32
LARSEN et al.
PALAEOCENE ALKALI PICRITES, WEST GREENLAND
Fig. 18. Metasomatism in the lithosphere envisaged as successive reaction zones around a crystallizing alkaline basic melt that releases traceelement-enriched residual liquid and fluid into the surrounding mantle, which acts as a chromatographic column. Upper half of diagram modified
from O’Reilly et al. (1991), based on data from southeastern Australia. Lower part of diagram shows the scheme envisaged for the old enrichment
event in the source for the Manı̂tdlat Member melts. The similarities between the trace-element spectra from Australia and Greenland
should be noted. The succession of trace elements in the spectra is that used by O’Reilly & Griffin (1988) and O’Reilly et al. (1991):
Cs–Rb–Ba–Th–U–K–Na–Ta–Nb–La–Ce–Sr–Nd–P–Sm–Zr–Hf–Ti–Tb–Y–Yb.
with similarities to that of the source for the Manı̂tdlat
Member. The age of the enrichment event is unknown.
These three occurrences of alkaline rocks within 150 km
of each other could perhaps be viewed as being derived
from one extensive metasomatized zone in the Archaean
lithosphere, mobilized at different times and by different
tectonic events. However, as discussed above, the par-
ticular setting of the Manı̂tdlat Member clearly indicates
that its enriched mantle source is an isolated domain.
Thus, at least in the Disko–Svartenhuk region, the metasomatized mantle domains are individual separate entities
1000 km2 or less in horizontal extent, which lends support
to the theory that the enrichment events were associated
with discrete intrusions of asthenospheric magmas.
33
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 1
JANUARY 2003
In the Paleocene, large volumes of hot, asthenospherederived tholeiitic magmas traversed the thinned and
fractured lithosphere. When the NW–SE migrating active
conduits traversed the enriched domains in the lithosphere, extensive melting of the incompatible-elementrich low-melting phases took place. The volume of enriched mantle was small and the low-melting component
was rapidly extracted, making the alkaline event very
short-lived.
CONCLUSIONS
The enriched alkaline picrites and basalts of the Manı̂tdlat
Member represent a discrete, short-lived melting event
during the formation of the tholeiitic picrite melts of
the Vaigat Formation. The ascending alkaline magmas
utilized the tholeiitic conduit systems, as shown by the
presence of occasional tholeiitic (low-Ca) olivine xenocrysts; however, the compositional effect of in-mixing of
small amounts of tholeiitic magma is very small.
The alkaline rocks have anhydrous primary mineral
assemblages of mainly olivine, chromite, clinopyroxene
and plagioclase; this may be explicable by high temperatures, with matrix glasses quenched at 1150–1190°C,
above the stability limit of amphibole.
Six alkaline magma types can be defined (Stordal and
Types 0, 1a, 1b, 2 and 3), all showing variable enrichment
in Ba, U, Nb–Ta and LREE. Five types show relative
depletion in Rb, Th, K, Zr–Hf and Ti, whereas the least
enriched Type 2 shows no depletion in these elements
and particular enrichment in Nb–Ta and Zr–Hf. Unusual
Sr–Nd–Pb–Os isotope compositions, with low 87Sr/86Sr,
143
Nd/144Nd and 187Os/188Os, and high Pb isotope ratios,
are outside the range of any modern asthenospheric
mantle components and have resulted from melting of
source rocks with long-term low Rb/Sr, Sm/Nd and
Re/Os, and high U/Pb and U/Th.
The most likely origin for the alkaline melts is in
old metasomatized lithospheric mantle domains rich in
amphibole, clinopyroxene and apatite, some parts also
containing mica and other parts containing small amounts
of metasomatic oxides (lindsleyite or hawthorneite,
LIHA), and possibly carbonate. The amounts of the
mantle phases going into the melt can be assessed by
simple mass-balance calculations, which show that the
bulk of the melts is made up of >60% pargasitic amphibole, 26–30% clinopyroxene, >9% olivine and >1%
apatite. The melt types most enriched in Ba, U, Nb–Ta
and LREE in addition require 0·3–0·7% LIHA phase,
and some types (0, 1b and 3) require an additional phase
(carbonate?) with these elements. The least enriched
Type 2 melt requires 4% mica in the melting assemblage.
The enriched mantle domains must be highly structured to give rise to the various melt types. Good correlations between some incompatible element ratios
suggest a common link between the melt types. A mantle
with metasomatic zones produced during an old event
of migration of incompatible elements away from a
crystallizing alkaline intrusion and precipitation of these
elements in successive zones, as in a chromatographic
column, can explain the partly complementary relation
between the melt types. The zone closest to the contact
selectively trapped Rb, K, Zr–Hf and Ti and later gave
rise to the melts of Type 2, whereas the most distal zones
gave rise to the most Ba–U–Ta–LREE-enriched melts.
ACKNOWLEDGEMENTS
We are grateful to J. Rønsbo for assistance with the setup of the microprobe, and particularly for the highprecision analyses. S. Bernstein and W. L. Griffin put
unpublished data at our disposal. T. Andersen, S. Bernstein, W. L. Griffin, H. Hansen, P. M. Holm and D.
Peate provided constructive comments and discussions,
as did M. Wilson and the reviewers N. Arndt, A. C.
Kerr and S. Y. O’Reilly. The Geological Survey of
Greenland and the Arctic Station in Godhavn provided
extensive support during fieldwork. The Danish Natural
Science Research Council provided the electron microprobe and the X-ray fluorescence spectrometer at the
University of Copenhagen. This paper materialized as a
result of a grant from the Carlsberg Foundation to L.M.L.
and is published with the permission of the Geological
Survey of Denmark and Greenland.
SUPPLEMENTARY DATA
Supplementary data are available on Journal of Petrology
online.
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VOLUME 44
NUMBER 1
JANUARY 2003
For the incompatible trace elements, the known inventory of these in the primary alkaline melt was distributed between the contributing phases with a simple
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Ci,1 = Ci,x1p1 + Ci,x2p2 + Ci,x3p3 + . . .
where C is concentration, i is element i, l is the liquid,
x1, x2, x3 are the phases, and p1, p2, p3 are the fractions
of the phases making up the melt. Further, the element
concentrations in all phases are related to that in the
amphibole:
Ci,x = Ci,ampDx/amp
where Dx/amp is the mineral–amphibole distribution coefficient for element i. Then
Ci,1 = Ci,amppamp + Ci,ampDcpx/amppcpx +
Ci,ampDmica/amppmica + . . .
or
Ci,1 = Ci,amp(pamp + Dx/amppx).
Distribution coefficients Dx/amp for the various minerals
were taken from the literature or interpolated or estimated. For each of the six types of Manı̂tdlat Member
melt Ci,l is the measured trace element concentration
recalculated at 15 wt % MgO. The fractions of the
various phases in the melt, px, can then be adjusted to
give realistic element concentrations in the melting mantle
minerals. The calculated concentrations of trace elements
in the melting phases are given in Electronic Appendix
B.
The fractions of amphibole, clinopyroxene and olivine
contributing to the melt are taken from the major-element
calculations.
The fraction of apatite is accurately determined by the
P2O5 content in the melt and the assumption that apatite
contains 40 wt % P2O5.
The fraction of mica is determined by the K2O and
Rb contents in the melt and assumptions of 0·7–1 wt %
K2O in amphibole, >8 wt % K2O in mica, and Rb
contents that must be realistic. Only the Type 2 melt
has enough K and Rb to include a mica component.
The fraction of Ba–Ti-oxide (LIHA) is determined by
the concentrations of Ba and Nb in the amphibole.
Maximum concentrations were put at 800 ppm Ba and
100 ppm Nb because very few measured mantle amphiboles contain more than this and most are well below
that level (O’Reilly et al., 1991; Ionov & Hofmann, 1995;
Chazot et al., 1996; Ionov et al., 1997; Ionov, 1998;
Zanetti et al., 1999; Moine et al., 2001), although some
higher values have been reported (Ionov & Hofmann,
1995; Raffone et al., 2001). When the concentrations of
Ba and Nb in the amphibole exceeded the maximum, an
LIHA phase was introduced into the melting assemblage.
APPENDIX: CALCULATION OF
MELTING ASSEMBLAGES IN THE
GENERATION OF THE MANÎTDLAT
MEMBER MAGMAS
The primary alkaline melts are considered to have
15 wt % MgO, and the bulk-rock analyses are recalculated to this value by addition or subtraction of
olivine; elements incompatible in olivine are simply diluted or concentrated by the recalculation. For the Type
3 (ankaramitic) melts we used a sample with 10 wt %
MgO and minimal evidence for clinopyroxene fractionation. The Stordal sample has, however, fractionated
clinopyroxene.
The modelling assumes total consumption of all the
enriched mantle phases except clinopyroxene during
melting; then the calculations can be made using simple
mass-balance equations. The calculations are available
as spreadsheets in Electronic Appendix B, which may be
downloaded from the Journal of Petrology web site at http://
www.petrology.oupjournals.org.
For the major elements, combinations of analysed
mantle minerals were used in mixing equations to approach the major-element composition of the primary
alkaline melts. FeO and MgO were added up to make
allowance for Fe/Mg fractionation during melting.
38