JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 1
PAGES 169–190
2005
doi:10.1093/petrology/egh067
The Role of Continental Crust and
Lithospheric Mantle in the Genesis of
Cameroon Volcanic Line Lavas: Constraints
from Isotopic Variations in Lavas and
Megacrysts from the Biu and Jos Plateaux
K. RANKENBURG1,2*, J. C. LASSITER1 AND G. BREY2
1
MAX-PLANCK-INSTITUT FÜR CHEMIE, ABT. GEOCHEMIE, POSTFACH 3060, 55020 MAINZ, GERMANY
2
INSTITUT FÜR MINERALOGIE, SENCKENBERGANLAGE 28, 60054 FRANKFURT/MAIN, GERMANY
RECEIVED AUGUST 26, 2002; ACCEPTED AUGUST 3, 2004
ADVANCE ACCESS PUBLICATION OCTOBER 1, 2004
We present a combined Sr, Nd, Pb and Os isotope study of lavas and
associated genetically related megacrysts from the Biu and Jos
Plateaux, northern Cameroon Volcanic Line (CVL). Comparison of
lavas and megacrysts allows us to distinguish between two contamination paths of the primary magmas. The first is characterized by
both increasing 206Pb/204Pb (1982–2033) and 87Sr/86Sr
(070290–070310), and decreasing eNd (70–60), and involves
addition of an enriched sub-continental lithospheric mantle-derived
melt. The second contamination path is characterized by decreasing
206
Pb/204Pb (1982–1903), but also increasing 87Sr/86Sr
(070290–070359), increasing 187Os/188Os ( 0130–
0245) and decreasing eNd (70–46), and involves addition of up to
8% bulk continental crust. Isotopic systematics of some lavas from the
oceanic sector of the CVL also imply the involvement of a continental
crustal component. Assuming that the line as a whole shares a common
source, we propose that the continental signature seen in the oceanic sector
of the CVL is caused by shallow contamination, either by continentderived sediments or by rafted crustal blocks that became trapped in the
oceanic lithosphere during continental breakup in the Mesozoic.
The Cameroon Volcanic Line (CVL) comprises a genetically related series of Cenozoic intraplate volcanoes that
extend for 1600 km from the island of Annobon (formerly
known as Pagalu) in the South Atlantic Ocean to the
continental interior of West Africa (Fig. 1). The northern
end of the continental part of the CVL is marked by the
Cenozoic volcanism of the Biu Plateau, Nigeria. Fitton &
Dunlop (1985) showed that basaltic rocks in the oceanic
and continental sectors of the CVL are geochemically
and isotopically (87Sr/86Sr) similar and suggested that a
line or zone of hot asthenospheric mantle is upwelling
underneath the region, partial melting of which has generated parental magmas without any substantial involvement of the overlying lithosphere.
This simple picture was challenged by combined Nd,
Sr, Pb and O isotope studies of Halliday et al. (1988,
1990), in which those workers found a distinctive
206
Pb/204Pb anomaly in CVL lavas focused at the
continent–ocean boundary (c.o.b.), which diminishes
over a distance of 400 km to either side. Halliday et al.
considered this HIMU Pb isotope signature (high m high 238U/204Pb, leading to time-integrated high
206
Pb/204Pb) to be inherited from relatively recent U/
Pb fractionation at 125 Ma during impregnation of the
uppermost mantle by the St. Helena hotspot when the
Equatorial Atlantic opened. The observed Pb isotope
heterogeneity of the CVL lavas was therefore proposed
to be derived from remelting of variably metasomatized
lithosphere rather than reflecting primary asthenospheric
source heterogeneity. From a study of peridotite xenoliths
*Corresponding author. Telephone: þ1 281 244 1084. Fax: þ1 281
483 1573. E-mail: kai.rankenburg1.jsc.nasa.gov
Journal of Petrology vol. 46 issue 1 # Oxford University Press 2004;
all rights reserved
KEY WORDS:
crustal contamination; CVL; megacrysts; ocean floor;
osmium isotopes
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 1
JANUARY 2005
Fig. 1. Geological map showing the eruption ages of the major volcanic centres of the Cameroon Volcanic Line and the Gulf of Guinea [adapted
from Fitton & Dunlop (1985)]. Ages compiled from Fitton & Dunlop (1985), Halliday et al. (1990), Lee et al. (1994) and Ngounouno et al. (1997).
The Jos volcanics are located 400 km to the NW of the line axis and are usually not included in CVL magmatism. However, no occurrence of
continental Cenozoic volcanism has been recorded west of the Jos Plateau. Sample locations are indicated by grey triangle (Biu Plateau) and black
square ( Jos Plateau).
Lee et al. (1996) provided evidence that portions of the
lithospheric mantle beneath the CVL are isotopically
enriched. There is also qualitative evidence for interaction with the continental crust in some evolved lavas of
the continental sector based upon large variations in Hf
isotopes (Ballentine et al., 1997), 87Sr/86Sr as high as
0705–0714 (Marzoli et al., 1999) and the Sr–Nd isotope
systematics of lavas and genetically related megacrysts
(Rankenburg et al., 2004).
In this study, we examine the respective contributions
of crustal contamination and assimilation of subcontinental lithospheric mantle (SCLM) by comparing
the isotopic (Sr, Nd, Pb and Os) and trace element variations of Biu and Jos Plateau lavas with the compositions
of genetically related megacrysts that grew at mantle
depth. We have analysed Sr, Nd and Pb isotopes in
36 whole rocks and 13 megacrysts collected from the
Biu and Jos Plateaux, as well as osmium isotopes of a
subset of 17 rock samples. The Re–Os isotope system
provides an excellent tool for discrimination between
assimilation of continental crust or the SCLM. Unlike
Sr, Nd and Pb isotope compositions, which may overlap
in both continental crust and the SCLM, there is generally a strong contrast in osmium isotopes between the
continental crust and the peridotitic SCLM as a result of
the compatible behaviour of Os during mantle melting.
Whereas continental crust generally has developed
variable but high 187Os/188Os ratios over time
170
RANKENBURG et al.
CAMEROON VOLCANIC LINE LAVAS
(e.g. Esser & Turekian, 1993; Esperanca et al., 1997), the
SCLM generally has complementary unradiogenic
187
Os/188Os ratios (e.g. Walker et al., 1989). Thus, if a
melt is contaminated by old crust-derived material, it
should have an unusually radiogenic Os isotope
signature. In contrast, contaminants derived from the
peridotitic SCLM should have unradiogenic Os isotope
compositions.
Pyroxenite xenoliths derived from the SCLM may also
have a radiogenic Os isotope signature (e.g. Reisberg et al.,
1991; Roy-Barman et al., 1996; Lassiter et al., 2000;
Pearson & Nowell, 2004). Thus melting of pyroxenite
layers or veins in the SCLM has been invoked to explain
the ubiquity of elevated Os isotope ratios in ocean island
basalt (OIB) (e.g. Hauri & Hart, 1993; Schiano et al.,
1997; Lassiter et al., 2000; Hauri, 2002; Kogiso et al.,
2004). However, contamination with pyroxenite-derived
melts may be distinguished from crustal material based
upon other geochemical tracers, such as, for example, Pb
isotope and trace element signatures.
Geological setting: the Benue Trough and
Biu and Jos Plateaux
The continental sector of the CVL has a Y-shaped form
(see Fig. 1). Whereas most previous studies considered the
Biu Plateau as the end of the NNW branch of the continental sector of the CVL (e.g. Turner, 1978; Fitton,
1980; Halliday et al., 1988; Poudjom-Djomani et al.,
1995; Lee et al., 1996; Ballentine et al., 1997; Barfod et al.,
1999; Marzoli et al., 2000), the Jos Plateau (located
c. 400 km to the NW of the central CVL axis, see Fig. 1)
is usually not assigned to CVL volcanism. However, the
timing of the Jos Plateau volcanism is very similar to
that of the other CVL volcanic centres (Grant et al.,
1972). The Biu and Jos Plateau lavas have similar
major and trace element chemistry, and Jos Plateau
lavas also span a similar range in isotopic compositions,
overlapping the data of the CVL as a whole (Rankenburg
et al., 2004). We therefore consider the Jos Plateau to
be associated with CVL volcanism in the following
discussion.
According to Turner (1978), the Biu Plateau was constructed in three stages during two periods of volcanism:
(1) an early fissure type eruption; (2) formation of relatively large tephra ring volcanoes and building up of
localized thick lava piles (up to 250 m) in the southern
part of the plateau. Lavas of this plateau-building stage
range in composition from hy-normative basalt to basanite, with K/Ar ages from 535 to 084 Ma (Grant et al.,
1972; Fitton & Dunlop, 1985). Extensive weathering and
laterite formation suggest a hiatus after this episode. (3)
Resumption of igneous activity with the formation of over
80 NNW–SSE-aligned cinder cones with similar chemistry to the earlier basalts. A rough estimate of the age of
the last magmatic period is <50 ka based on diffusional
constraints of He in mantle xenoliths of the CVL (Barfod
et al., 1999) and >25 ka based on pollen dating of maar
sediments from the Biu Plateau (Salzmann, 2000).
As with the Biu Plateau, volcanic activity on the Jos
Plateau occurred in two periods and thus the basalts from
this region have been divided into an earlier and a more
recent group (McLeod et al., 1971). There are no isotopic
age determinations available for the older basalts, but
Wright (1976) suggested a Paleocene age, roughly synchronous with Benue Trough folding and uplift. The
more recent activity formed a group of 22 cinder cones.
Radiometric K–Ar ages (Grant et al., 1972) suggest,
unlike on the Biu Plateau, continuous volcanism between
21 and 09 Ma.
The younger volcanics of both the Biu and Jos Plateaux
are characterized by abundant inclusions of mantle xenoliths and megacrysts. The megacryst suites of the Biu and
Jos Plateaux were described in detail by Wright (1970)
and Frisch & Wright (1971), and comprise chemically
homogeneous crystals of clinopyroxene (cpx), garnet
(gnt), plagioclase (plag) and ilmenite (ilm) with diameters
of up to several centimetres, whereas crystals of olivine
(ol), amphibole (amph), spinel (sp), apatite (apa), zircon
(zr) and blue corundum (cor) are extremely rare.
SAMPLING AND ANALYTICAL
TECHNIQUES
Major and trace element data were obtained for 27 volcanic rocks from the younger Biu Plateau suite, four rocks
from the older, plateau-building suite of the Biu Plateau,
and five rocks from the younger Jos Plateau suites (Table 1).
The lavas were first coarsely crushed in steel mortars.
Selected chips free of obvious xenocrysts or alteration
were then powdered in an agate ring-disc mill. The powders were analysed by X-ray fluorescence spectroscopy
(XRF) with a Philips PW 1404 instrument at the University
of Frankfurt using Li-borate glass discs for major elements
and at the University of Mainz using pressed powder pellets
for trace elements. Rock powders were commercially analysed at the University of Goettingen, Germany (all samples) and at the Memorial University of Newfoundland,
Canada (subset of 17 samples) by inductively coupled
plasma mass spectrometry (ICP-MS) following HF–
HNO3 acid dissolution [analytical details have been given
by Jenner et al. (1990)]. A subset of 20 samples was additionally analysed for rare earth element (REE) concentrations by inductively coupled plasma atomic emission
spectrometry (ICP-AES) following sinter dissolution at the
GeoForschungsZentrum in Potsdam (Zuleger & Erzinger,
1988). Comparison of all the datasets revealed problems of
the Goettingen ICP-MS laboratory with respect to accurate determination of high field strength element (HFSE)
171
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 1
JANUARY 2005
Table 1: Major (wt %) and trace (ppm) element analyses of Biu and Jos Plateau lavas
Sample:
ZAGU
JIGU 1
JIGU-M
X
PELA JUNG
KOROKO
PELA ALT
DAM
DAM2
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
SiO2
TiO2
Al2O3
FeOt
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
Rb
Ba
Th
U
46.42
2.61
49.17
2.17
44.34
3.08
47.82
2.33
49.44
2.19
59.88
0.35
51.05
2.26
44.33
2.99
47.80
2.60
14.39
10.49
14.74
9.71
12.66
11.28
14.43
10.40
14.74
9.71
20.50
3.10
15.25
9.27
12.50
11.48
14.12
10.20
0.18
9.15
0.15
8.64
0.16
8.81
0.15
8.67
0.18
0.37
0.14
7.16
0.20
0.16
9.23
9.26
3.71
1.82
9.18
3.32
0.19
10.78
9.89
9.89
3.16
8.41
3.42
2.02
8.33
7.18
3.60
1.38
3.75
1.77
0.80
98.83
0.47
98.92
0.99
98.74
1.29
0.54
98.84
1.62
0.56
98.92
4.81
0.12
99.65
2.45
0.59
98.97
50.4
786
7.93
1.99
32.4
469
4.36
1.06
Nb
92
51
Ta
5.34
58.8
3.43
30.2
La
Ce
Pb
Pr
Sr
Nd
Zr
Hf
Sm
Eu
Gd
Tb
Dy
Ho
Y
Er
Tm
Yb
Lu
Sc
110
3.08
11.7
973
48.7
272
59.4
2.21
6.74
610
28.1
187
6.93
9.38
5.27
6.00
3.18
7.72
2.08
5.36
1.19
6.07
0.86
4.48
1.11
27.3
0.84
20.4
2.71
0.36
2.11
0.27
2.03
0.29
1.57
0.23
20.8
21.3
46.2
598
8.38
2.43
104
6.19
67.8
138
3.55
15.3
1009
62.4
352
8.66
12.4
3.90
9.67
1.41
6.96
1.22
29.5
3.10
0.37
2.20
0.31
20.4
36.1
39.0
432
500
5.30
1.28
6.06
1.54
62
n.m.
37.5
72.9
2.34
668
339
40.6
77.5
120
703
33.2
35.0
183
230
5.78
7.02
2.37
n.m.
7.01
2.32
5.95
5.94
0.97
0.93
4.88
4.87
0.87
0.87
23.3
21.8
2.17
2.26
0.29
0.26
1.61
1.74
0.25
0.21
19.4
24.2
36.75
10.04
67
3.96
2.87
8.57
8.10
196
1135
19.25
206
16.90
20
1012
64
808
17.84
11
3.1
8.5
1.2
7.0
1.3
34
3.7
0.48
3.5
0.51
0.75
66.8
830
7.57
1.98
93
n.m.
53.8
101
3.60
10.7
805
42.8
346
n.m.
8.05
2.67
6.44
0.94
4.36
0.76
20.2
1.86
0.23
1.22
0.18
16.0
11.50
10.08
3.28
1.56
0.80
98.72
40.7
736
7.06
1.89
85
n.m.
55.2
113
2.95
12.6
858
51.3
305
n.m.
10.6
3.32
8.30
1.25
6.25
1.11
27.1
2.93
0.36
2.07
0.31
22.0
8.71
3.16
2.15
0.72
98.86
55.8
747
7.86
1.94
89
5.46
52.7
104
3.88
11.2
872
45.4
313
7.28
9.19
2.93
7.28
1.11
5.39
0.95
24.2
2.36
0.30
1.69
0.23
18.6
V
155
164
201
183
166
9
184
192
Cr
231
299
379
287
279
10
214
401
246
Ni
202
191
278
176
210
3
34.3
202
291
222
Zr/Nb
28.9
2.96
19.2
3.67
30.8
3.38
21.6
2.95
25.2
3.43
Ce/Pb
35.7
26.9
38.8
31.2
27.0
La/Yb
K/U
7612
10792
6061
8400
8737
172
2.38
12.2
3974
183
44.1
3.72
26.6
3.59
31.1
3.52
28.2
38.2
26.8
10263
6838
9208
RANKENBURG et al.
CAMEROON VOLCANIC LINE LAVAS
Sample:
BUGOR
SE BUGOR
HILIA 1
HILIA 2
TAMZA
GUFKA
MIR
GULD-UMBUR
PELA 2
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
46.97
2.43
46.35
2.45
46.41
2.59
49.26
2.19
46.51
3.06
51.17
1.93
53.24
1.92
45.98
2.59
45.54
2.88
13.95
10.40
14.16
10.42
14.56
10.48
14.69
9.96
14.17
10.62
15.34
9.03
16.19
7.51
12.91
10.65
13.23
10.98
0.17
9.16
0.15
8.59
0.17
9.28
0.13
7.35
0.12
5.91
9.23
9.27
3.17
1.16
8.86
8.90
3.36
2.00
3.49
1.14
SiO2
TiO2
Al2O3
FeO
MnO
0.16
0.17
MgO
10.82
8.94
10.15
9.57
3.13
1.46
3.43
1.53
CaO
Na2O
K2O
3.72
1.77
0.18
0.18
6.80
11.87
9.35
10.18
10.05
3.94
2.93
2.93
1.61
3.35
1.62
P2O5
0.58
0.62
0.74
0.46
0.79
0.52
0.59
0.75
0.75
Total
98.84
37.9
98.84
38.3
98.83
50.2
98.89
28.3
98.82
43.5
98.99
24.3
99.16
73.2
98.81
45.5
98.78
45.1
Rb
Ba
Th
U
441
5.20
1.39
438
5.65
1.51
628
6.49
1.67
Nb
65
73
83
Ta
n.m.
38.1
n.m.
39.5
n.m.
50.0
La
Ce
Pb
Pr
Sr
Nd
Zr
Hf
Sm
Eu
Gd
Tb
Dy
Ho
Y
Er
Tm
Yb
Lu
Sc
72.9
2.13
8.34
668
33.9
211
76.6
2.16
8.67
714
34.5
206
96.5
2.64
10.2
860
42.1
259
n.m.
7.03
n.m.
7.01
n.m.
8.49
2.37
6.02
2.43
6.11
2.72
6.87
0.99
5.12
0.95
4.93
1.05
5.39
379
604
4.07
0.65
51
8.69
2.36
103
3.40
30.2
60.1
1.78
6.19
58.2
115
3.67
12.3
6.85
567
1013
28.2
183
49.7
338
5.20
6.17
2.20
10.0
3.20
8.29
5.33
0.86
7.82
1.20
5.75
1.01
440
4.24
1.01
52
2.87
30.6
58.7
1.86
6.67
667
27.4
173
4.10
5.81
2.03
4.95
0.78
3.93
864
10.29
2.67
120
7.48
62.9
120
3.91
12.5
1075
47.2
355
531
7.37
2.12
86
4.51
50.6
97.8
2.76
10.6
804
44.4
263
8.97
6.03
8.7
2.83
8.72
2.87
6.75
0.92
6.78
1.08
5.50
0.98
579
6.39
1.52
82
n.m.
51.0
99.2
2.84
10.9
824
46.5
279
n.m.
9.16
3.01
7.20
1.09
5.64
0.93
0.89
0.95
4.49
0.82
0.69
4.03
0.64
23.9
2.39
23.0
2.33
26.7
2.48
20.3
2.04
25.4
2.60
17.8
1.67
16.0
1.55
25.1
2.54
29.3
2.54
0.30
1.74
0.31
1.76
0.32
1.82
0.27
1.53
0.32
1.86
0.22
1.23
0.17
0.99
0.33
1.90
0.33
1.84
0.25
22.1
0.26
21.6
0.26
22.2
0.22
0.27
21.1
18.3
0.17
17.4
0.14
11.7
0.27
21.4
0.99
0.28
24.2
V
184
181
160
171
197
141
122
190
195
Cr
368
290
196
273
224
242
122
387
397
Ni
307
231
183
220
202
196
123
346
243
La/Yb
Zr/Nb
Ce/Pb
K/U
21.9
3.25
34.3
8736
22.4
2.82
35.4
8416
27.5
3.12
36.6
8814
19.8
3.59
31.3
3.28
33.7
14900
31.3
7038
173
24.8
3.33
31.5
9324
63.8
2.96
30.8
9120
26.7
3.06
35.5
6295
27.7
3.40
34.9
8845
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 1
JANUARY 2005
Table 1: continued
Sample:
WIGA
GWARAM
ZUMTA
HIZSHI
TUM
ETUM
GUMJA
TILA 1
TILA STR
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
45.66
3.04
50.10
2.38
50.03
2.20
47.77
2.67
46.06
2.40
48.28
2.57
47.49
2.69
46.30
3.09
47.40
2.77
13.55
10.88
15.17
9.28
15.42
8.89
14.25
10.22
13.20
10.58
14.34
10.16
14.46
10.02
13.48
11.03
14.22
11.14
MnO
0.17
0.14
7.78
0.17
8.81
0.18
9.16
0.15
9.43
0.16
10.69
9.71
0.15
7.04
0.16
MgO
0.15
9.33
7.27
4.55
2.33
7.67
3.77
2.45
8.10
4.48
1.59
8.03
3.55
1.88
9.57
3.13
1.37
0.70
98.97
61.0
0.67
0.81
98.86
83.8
0.72
98.87
49.2
0.58
0.59
99.01
66.9
98.88
32.7
98.77
35.8
SiO2
TiO2
Al2O3
FeO
CaO
Na2O
K2O
2.67
1.59
P2O5
0.85
Total
98.79
33.3
Rb
Ba
Th
U
519
7.36
1.96
858
10.5
2.73
Nb
88
115
Ta
La
n.m.
54.9
n.m.
72.0
Ce
106
139
Pb
Pr
Sr
Nd
Zr
Hf
Sm
Eu
Gd
Tb
Dy
Ho
Y
Er
Tm
Yb
Lu
Sc
2.94
11.5
895
46.7
264
n.m.
9.57
4.46
15.1
1044
59.7
429
n.m.
11.5
805
8.22
2.00
96
6.47
54.5
102
3.39
11.0
918
42.8
282
8.16
8.04
12.89
8.71
2.83
1.41
0.57
98.82
42.1
723
498
8.98
2.39
5.78
1.43
109
73
4.40
n.m.
67.9
40.2
78.1
133
3.75
14.4
2.22
8.55
1017
692
56.7
34.2
386
220
n.m.
11.3
518
7.83
2.08
94
n.m.
54.7
108
3.31
12.1
453
6.57
1.72
72
3.95
42.7
78.6
2.60
8.43
918
48.8
748
367
200
34.5
6.06
6.89
n.m.
9.55
5.15
6.77
3.02
7.49
3.67
8.79
2.63
6.33
3.49
8.68
2.29
5.94
3.09
7.90
2.31
5.80
1.15
5.61
1.30
6.10
0.95
4.33
1.26
6.09
0.89
4.64
1.24
6.70
0.92
4.88
0.94
24.4
1.01
26.2
0.70
18.1
1.03
27.4
0.82
20.9
1.22
28.2
0.86
21.7
2.44
0.31
2.55
0.29
1.78
0.20
2.67
0.30
2.09
0.27
2.94
0.39
2.18
0.29
1.67
0.24
1.64
0.22
1.21
0.16
1.73
0.25
1.61
0.24
2.38
0.33
1.70
0.24
20.0
14.3
15.7
17.3
23.7
17.2
20.4
10.34
9.83
2.55
1.41
384
4.54
1.11
61
n.m.
36.6
76.7
2.19
9.01
671
38.9
252
n.m.
9.21
2.42
1.60
0.53
98.76
40.7
383
4.14
1.10
58
n.m.
31.1
64.2
1.88
7.49
708
31.1
216
n.m.
8.48
2.71
7.04
2.38
6.93
1.09
5.88
0.95
5.51
0.96
4.91
0.85
25.7
2.41
21.3
2.13
0.30
1.76
0.25
1.50
0.25
0.20
16.7
23.8
V
205
145
153
162
195
149
197
210
185
Cr
300
187
217
265
452
333
229
286
283
Ni
269
174
202
233
435
241
23.0
229
230
246
20.7
Zr/Nb
32.8
3.00
43.9
3.73
45.0
2.94
39.3
3.54
24.9
3.01
Ce/Pb
36.2
31.3
30.1
35.3
35.1
La/Yb
K/U
6742
7065
10167
5523
8237
174
3.90
32.6
7476
25.1
2.78
20.8
4.13
30.2
35.1
6638
10499
3.72
34.1
12093
RANKENBURG et al.
CAMEROON VOLCANIC LINE LAVAS
Sample:
DAI
KERANG
AMPANG
PIDONG-M
PIDONG-S
Biu4
Biu5
Biu8
Biu9
melt incl.
Group:
Jos
Jos
Jos
Jos
Jos
Biu old
Biu old
Biu old
Biu old
Biu young
(mean of 5)
SiO2
TiO2
Al2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
Rb
Ba
Th
U
46.26
2.44
44.84
2.66
46.69
2.30
45.63
2.43
47.82
2.67
46.18
2.96
45.93
3.15
48.01
2.21
46.63
2.77
53.22
2.43
13.78
10.45
13.68
11.29
13.16
10.48
13.46
10.68
16.30
10.63
13.30
10.89
13.13
11.13
13.95
10.21
13.84
10.54
17.56
6.24
0.17
10.77
9.77
0.19
9.49
0.18
0.16
9.97
0.19
10.14
9.74
2.92
1.69
3.99
1.81
3.33
1.80
3.60
1.79
10.71
2.69
10.70
9.68
0.07
1.49
8.64
3.47
2.32
0.17
10.48
10.31
0.17
10.79
9.54
0.17
6.13
9.91
0.18
11.05
9.12
2.38
1.50
2.41
1.13
2.63
1.62
0.60
98.84
0.89
98.74
0.72
98.83
0.71
98.81
0.66
98.82
0.58
98.76
41.8
0.40
0.73
0.99
98.86
24.6
98.83
44.9
95.56
45.1
554
6.08
1.47
Nb
72
Ta
n.m.
42.5
La
Ce
Pb
Pr
Sr
Nd
Zr
Hf
Sm
Eu
Gd
Tb
Dy
Ho
Y
Er
Tm
Yb
Lu
Sc
84.1
2.84
9.24
803
37.9
231
n.m.
7.74
2.57
6.44
0.99
4.98
0.89
23.3
2.27
0.28
1.64
0.24
22.8
54.5
734
8.31
2.01
102
5.43
67.5
128
4.25
13.6
1024
53.5
293
6.58
10.3
3.51
8.58
1.24
5.91
0.98
27.7
2.54
0.30
1.67
0.23
20.2
53.0
726
7.46
1.79
86
4.57
56.5
107
4.56
11.3
858
44.2
247
54.5
615
7.21
1.80
56.7
795
7.74
1.95
86
97
n.m.
52.4
n.m.
62.4
101
125
3.81
10.8
841
43.7
257
4.21
13.4
1128
52.2
331
5.55
8.5
n.m.
8.90
2.83
6.81
2.93
6.96
3.15
7.47
1.00
4.92
1.07
5.23
1.08
5.21
0.84
22.0
0.92
24.7
0.91
21.4
2.13
0.26
2.35
0.29
2.33
0.29
1.49
0.22
1.57
0.24
1.62
0.24
18.6
20.7
n.m.
9.6
13.7
1.32
0.61
98.79
29.0
498
4.21
1.12
63
4.16
34.4
70.6
1.87
8.37
712
35.1
240
7.29
7.59
2.53
390
4.66
1.16
63
n.m.
36.8
78.1
2.13
8.97
635
38.3
257
n.m.
351
3.08
0.76
43
n.m.
25.3
51.8
1.90
6.14
497
25.1
164
n.m.
8.11
2.75
5.52
1.84
6.50
1.07
6.96
1.08
4.63
0.74
5.27
0.91
5.20
0.93
3.97
0.72
24.1
2.32
24.0
2.34
18.6
1.78
0.29
1.65
0.28
1.54
0.23
1.44
0.23
21.2
0.22
21.8
0.21
20.8
535
6.49
1.73
4.50
5.93
3.13
n.m.
1179
n.m.
n.m.
88
190
n.m.
53.1
n.m.
79.5
105
3.53
11.4
888
47.5
297
n.m.
9.61
2.99
7.61
1.16
5.98
1.06
29.1
2.68
0.35
2.04
0.30
22.9
156
n.m
18.4
1598
66.6
361
8.2
11.8
3.75
7.48
0.78
3.54
0.57
11.9
1.05
b.d.
0.09
b.d.
n.m.
V
175
167
156
157
145
206
211
174
179
n.m.
Cr
323
205
417
347
45
355
305
396
358
b.d.
Ni
230
161
345
260
75
244
230
242
226
b.d.
Zr/Nb
25.9
3.21
40.5
2.87
37.8
2.87
33.3
2.99
38.6
3.41
20.9
3.81
24.0
4.08
17.6
3.81
26.1
3.38
Ce/Pb
29.6
30.2
23.5
26.6
29.7
37.8
36.7
27.2
29.8
La/Yb
K/U
9521
7445
8326
8276
9854
n.m., not measured; b.d., below detection limit.
175
9791
10697
12237
7755
880
1.9
JOURNAL OF PETROLOGY
VOLUME 46
concentrations, most probably because of precipitation of
insoluble fluorides from the sample solution (Yokoyama
et al., 1999). XRF data are, therefore, reported in Table 1
for Ba, Sr, Nb, Zr, V, Cr and Ni with errors of <5%. All
other data in Table 1 are from ICP-MS and ICP-AES
analyses with errors on REE, Cs, Rb, Th, U and Pb of
<10%, and on Cs, Lu, Ta and Hf of <18%.
Pb isotope analyses were carried out on all 36 lavas and
13 megacrysts (Table 2). The lavas were coarsely crushed,
sieved and washed in 1N HCl and distilled water to
improve surface quality. Handpicked rock chips
( 100 mg of the 075–15 mm fraction) were leached in
hot 6N HCl for 1 h and washed ultrasonically in deionized water before dissolution in HF–HNO3. Lead was
extracted from the same sample solution that was used for
Sr–Nd isotope analyses by anion exchange in mixed
HBr–HNO3 media (Abouchami et al., 2000).
For Re–Os analysis, 2 g of clean rock chips were
handpicked. Complete avoidance of steel tools (which
may have considerable concentrations of Os and Re) is
not practically possible during the first steps of coarse
rock splitting. We used steel hammers in the field and a
hydraulic press made of hardened steel to coarsely crush
the rock to chips. To reduce possible metal contamination, rock samples were wrapped in thick plastic foil
before crushing in the hydraulic press. If small metal
particles are abraded during the crushing process, we
expect them to reside in the fine-grained fraction of the
crushed rocks or alternatively to be plated onto the surface of the rock chips. For Re–Os isotope analyses we
therefore used handpicked rock chips as raw material for
the subsequent rock processing in the agate ring mill, and
discarded the dust fraction. As an additional test, we
prepared a duplicate powder for one rock sample with
low Os concentration and highly radiogenic 187Os/188Os
(Biu 8). The low Os concentration of this sample makes it
particularly susceptible to possible contamination during
sample preparation. We thoroughly examined the rock
chips used to prepare the duplicate powder individually
under a binocular microscope to ensure clean, metal-free
surfaces before pulverizing the chips in an agate mill. The
187
Os/188Os value for the duplicate analysis is similar to,
although somewhat higher than, the original analysis,
confirming that the measured radiogenic Os signature is
a true sample feature.
For Re–Os analyses, the rock powder was digested in a
sealed quartz vessel together with a mixed 185Re/190Os
isotope tracer and concentrated HCl–HNO3 (2:3) for
16 h in a Perkin Elmer high-pressure asher operating at
ugman et al.,
10 MPa N2 overpressure and 300 C (Br€
1999). Osmium was extracted into liquid bromine and
purified by micro-distillation following the method of
Birck et al. (1997). Rhenium was separated and purified
from the residue using ion exchange extraction (Morgan
& Walker, 1989). Os and Re were subsequently loaded
NUMBER 1
JANUARY 2005
onto Pt filaments with a mixed Na(OH)–Ba(OH)2
emitter. The concentrations and isotopic compositions
reported in Table 2 were measured at the Max-PlanckInstitut, Mainz, by thermal ionization mass spectrometry
in negative ion mode (N-TIMS) using a Finnigan MAT262 system. The effects of fractionation during Os runs
were corrected for by normalizing the Os isotope ratios to
192
Os/188Os ¼ 30827 (Luck & Allegre, 1983). Six
procedural blanks for Os ranged from 016 pg to
145 pg with 187Os/188Os between 023 and 039, resulting in corrections on sample 187Os/188Os of <2%, and
corrections on sample Os concentrations of <2%.
Measured Re procedural blanks were 11–47 pg Re,
resulting in corrections to Re sample concentrations
of up to 20%. The Mainz in-house Os standard yielded
187
Os/188Os of 010703 20 (2s error, n ¼ 5). The
reproducibility of 187Os/188Os ratios based upon duplicate sample dissolutions was <27% of the mean, with
lower errors associated with higher Os concentrations of
the samples, but still significantly worse than the standard
reproducibility. This variability may have several
reasons. Osmium is a strongly chalcophile element and
the Os budget in a sample may be dominated by Os
contained in small sulphide globules (Roy-Barman et al.,
1998), which may have heterogeneous isotopic compositions. The specific distribution of sulphide globules within
aliquots of the sample powder therefore may account
for the poor sample reproducibility observed. The same
effect may also be responsible for the relatively poor
reproducibilities of Os and Re concentrations, which
were <24% and 54% of the mean. Similar variability
in Os concentrations and isotopic compositions in
individual flows has previously been reported by, for
example, Alves et al. (1999). Minor sample contamination
by disaggregated xenoliths may also contribute to sample
heterogeneity.
Analytical details and results for major element, trace
element and Sr–Nd isotope analyses of the megacrysts
have been given by Rankenburg et al. (2004). For Pb
isotope analyses of megacrysts (10 cpx and three feldspars; Table 3) 300 mg of cpx or 100 mg of plagioclase, respectively, were handpicked under a binocular
microscope in dark and bright field. Grains were then
leached twice in hot 25N HCl for 20 min, then in cold
5% HF for 15 min in an ultrasonic bath, rinsed with cold
25N HCl to remove fluoride complexes and finally
rinsed in deionized water. During a microscopic reexamination all grains with visible reaction rims were
removed to ensure 100% optically pure separates. Grains
were then dissolved in Teflon beakers using HF–HNO3.
Samples were run on a Finnigan MAT 261 multicollector
TIMS instrument in static mode. All data are reported
after fractionation correction of typically 0116% per
a.m.u. as determined by contemporaneous runs of the
NBS981 standard. External 2s reproducibility (n ¼ 29) of
176
RANKENBURG et al.
CAMEROON VOLCANIC LINE LAVAS
Table 2: Isotope analyses of Biu and Jos Plateau lavas, Os and Re concentrations measured by ID-TIMS
Sample:
ZAGU
JIGU 1
JIGU-M
X
PELA JUNG
KOROKO
PELA ALT
DAM
DAM 2
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
0.703351
0.703193
0.512928
5.66
0.702995
0.512963
6.34
0.703171
0.512961
6.30
87
Sr/86Sr
143
Nd/144Nd
eNd
206
Pb/204Pb
207
Pb/204Pb
208
Pb/204Pb
0.702951
0.512976
6.59
0.703504
0.512880
4.71
19.677
15.640
19.095
15.623
39.380
39.208
Os (ppt)
95
Re (ppt)
144
0.1423
187
Os/188Os
0.702892
0.513015
7.35
19.832
15.649
39.493
131 [101]
124
0.1290
[0.1311]
0.703183
0.512936
5.81
0.512901
5.12
19.384
15.620
19.481
15.647
20.103
15.693
39.259
39.505
39.717
0.703192
0.512928
5.66
19.860
15.681
39.953
19.517
15.629
19.654
15.655
39.306
39.694
19
350
51 [83]
113
244
61 [68]
0.1494
0.2450
0.1234
[0.1448]
Sample:
BUGOR
SE BUGOR
HILIA 1
HILIA 2
TAMZA
GUFKA
MIR
GULDUMBUR
PELA 2
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
0.702922
0.702927
0.512986
6.79
0.702923
0.512987
6.81
0.702888
0.702890
0.703079
0.512989
6.85
0.512968
6.43
0.512968
6.44
0.703061
0.512947
6.03
87
Sr/86Sr
143
Nd/144Nd
eNd
206
Pb/204Pb
207
Pb/204Pb
208
Pb/204Pb
0.512962
6.32
19.670
15.655
19.623
15.637
19.775
15.655
19.641
15.637
39.484
39.374
39.540
39.420
20.039
15.674
19.416
15.634
39.747
39.261
20.326
15.692
40.354
0.702914
0.512988
6.83
20.006
15.669
19.657
15.643
39.712
39.362
Os (ppt)
85
48
101
Re (ppt)
112
87
138
0.1284
187
Os/188Os
0.703090
0.512957
6.21
0.1384
0.1268
Sample:
WIGA
GWARAM
ZUMTA
HIZSHI
TUM
ETUM
GUMJA
TILA 1
TILA STR
Group:
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
Biu young
0.703140
0.512949
6.07
0.702900
0.512997
7.00
87
Sr/86Sr
143
Nd/144Nd
eNd
206
Pb/204Pb
207
Pb/204Pb
208
Pb/204Pb
Os (ppt)
0.702941
0.513004
7.14
19.781
15.640
39.512
185
Re (ppt)
187
Os/188Os
0.1333
0.702981
0.512973
6.53
20.018
15.672
39.798
0.703105
0.512958
6.24
20.162
15.684
40.281
0.702904
0.512963
6.34
0.703003
0.512968
6.44
20.109
15.670
39.779
0.702856
0.702890
0.512984
6.75
0.512991
6.89
20.018
15.674
19.878
15.643
39.844
39.520
19.711
15.656
19.557
15.644
19.773
15.640
39.470
39.422
39.506
111
160
133
217
111
116
197
151
396
0.1305
123
0.1290
0.1302
0.1266
0.1389
Sample:
DAI
KERANG
AMPANG
PIDONG-M
PIDONG-S
Biu4
Biu5
Biu8
Biu9
Group:
Jos
Jos
Jos
Jos
Jos
Biu old
Biu old
Biu old
Biu old
87
Sr/86Sr
143
Nd/144Nd
eNd
206
Pb/204Pb
207
Pb/204Pb
0.703379
0.512881
4.74
19.326
15.646
0.703393
0.512896
5.02
19.654
15.672
0.703588
0.512875
4.62
19.264
15.660
0.703370
0.703173
0.512872
4.56
0.512902
5.15
19.622
15.662
19.684
15.669
177
0.702884
0.512996
6.98
19.849
15.648
0.702934
0.512991
6.89
19.742
15.635
0.703341
0.512923
5.56
19.031
15.632
0.702942
0.512972
6.52
19.705
15.642
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 1
JANUARY 2005
Table 2: continued
Sample:
DAI
KERANG
AMPANG
PIDONG-M
PIDONG-S
Biu4
Biu5
Biu8
Biu9
Group:
Jos
Jos
Jos
Jos
Jos
Biu old
Biu old
Biu old
Biu old
208
39.319
39.477
39.616
39.514
Pb/204Pb
39.304
39.489
39.643
Os (ppt)
218
103
Re (ppt)
305
136
0.1373
0.1580
187
Os/188Os
39.175
66
70 [49]
155
75
0.1400
39.417
0.1601
[0.1687]
Os, Re concentrations and
187
Os/188Os blank corrected, duplicate analyses in square brackets.
Table 3: Lead isotope analyses of Biu and Jos Plateau megacrysts
Sample:
Mir þ Grt
Mir þ Plag
Mir 15
Mir 21
Dam þ
Dam þþ
Mineral:
cpx
cpx
cpx
cpx
cpx
cpx
cpx
Group:
Biu
Biu
Biu
Biu
Biu
Biu
Biu
206
Pb
20.658
15.723
20.663
15.736
20.621
15.719
20.724
15.737
20.048
15.708
20.103
15.711
20.902
15.757
Pb/204Pb
40.818
40.835
40.688
40.862
39.771
39.850
41.073
Pb/204Pb
207
Pb/
204
208
Pela alt
Sample:
Ker 2
Pid M
Ampang
Mir a
Mir b
Mir c
Mineral:
cpx
cpx
cpx
plag
plag
plag
Group:
Jos
Jos
Jos
Biu
Biu
Biu
206
Pb
19.823
15.672
19.924
15.685
20.039
15.701
20.702
15.725
20.678
15.714
20.752
15.726
Pb/204Pb
39.560
39.681
39.791
40.830
40.795
40.950
Pb/204Pb
207
Pb/
208
204
the standard was 387 ppm, 614 ppm and 861 ppm for
206
Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb ratios,
respectively. Total procedural lead blanks were <50 pg.
RESULTS
Major and trace elements
Most of the samples from the Biu and Jos Plateaux are
classified as basalts, and include both nepheline- and
hypersthene-normative compositions. The older plateaubuilding rocks of the Biu Plateau range from basanites to
trachybasalts. In contrast, the younger cinder cones have
more variable chemistry ranging from basanite to
phonolite. Only the younger suites of the Biu and Jos
Plateaux contain xenoliths (peridotites, pyroxenites and
crustal rocks) and megacrysts. Although the cinder cones
are only small volcanic structures that produced <2 km3
lava ( Turner, 1978), rock samples from a single location
can be chemically heterogeneous. For example, two
samples from a single cone (Dam) are basanite and
trachybasalt, respectively. Ilmenite megacrysts from a
rock of basaltic trachyandesitic composition (Miringa)
contain melt inclusions of a more evolved trachyandesitic
composition. This highlights the importance of magma
mixing processes in the genesis of the younger rocks.
Typical primitive mantle normalized trace element
patterns for the Biu and Jos Plateau samples are
shown in Fig. 2. Although all patterns are similar to
typical OIB patterns with positive Nb and negative Pb
anomalies, they are overall more enriched in incompatible elements than typical OIB [see the St. Helena
basalt SH68 (Thirlwall, 1997) for comparison in
Fig. 2]. The melt inclusions found in ilmenite megacrysts are highly enriched in incompatible trace elements but depleted in the heavy rare earth elements
(HREE), possibly indicating garnet fractionation in the
genesis of the melts. Two samples seem to be affected by
late-stage alteration because of their unusual low U/Nb
(Hilia 2), or low K/Rb (Hizshi), respectively. The phonolite pattern (not shown) has a large negative Ti
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Sample/Primitive Mantle
100
10
1
RbBaTh U Nb K LaCePb Pr Sr P Nd ZrSmEu Ti Dy Y YbLu
Fig. 2. Primitive mantle normalized trace element patterns for representative primitive samples of the younger and older Biu Plateau suites
and younger Jos Plateau suite, along with a basanite pattern from St.
Helena (Thirlwall, 1997), and the HREE-depleted pattern of the mean
of five melt inclusions found in ilmenite megacrysts.
anomaly pointing to fractionation of Fe–Ti-oxides,
whereas the lack of a europium anomaly indicates that
plagioclase was not a major fractionating phase in phonolite genesis. Excluding the two altered samples and
the phonolite, La/Yb and Zr/Nb ratios are similar for
Biu and Jos Plateau lavas and are in the range of 176–
638 and 28–41, respectively (Table 1). La/Yb and
Zr/Nb ratios are comparable with ratios reported for
HIMU-type OIB worldwide (e.g. Sun & McDonough
1989; Weaver, 1991; Thirlwall, 1997).
Sr, Nd, Pb and Os isotopes of lavas and
megacrysts
Figure 3 shows a compilation of Sr–Nd isotope compositions of the more primitive rocks of the CVL along with
our own analyses of volcanic rocks and associated megacrysts from the Biu and Jos Plateaux (Rankenburg et al.,
2004; Table 2). We chose an arbitrary cutoff at MgO
>5 wt % to exclude the rare continental CVL phonolites
and trachytes that fractionated within the continental
crust (Marzoli et al., 1999) and to highlight the variations
within the more primitive group of the lavas.
206
Pb/204Pb ratios of older and younger lavas from the
Biu Plateau overlap and range from 1903 to 2033 with a
mean of 1975 (Fig. 4). 206Pb/204Pb ratios in lavas from
the Jos Plateau are more restricted and range from 1926
to 1968. 207Pb/204Pb ratios in older and younger Biu
lavas range from 1562 to 1569. 207Pb/204Pb in Jos lavas
are slightly higher for a given 206Pb/204Pb than Biu lavas.
208
Pb/204Pb ratios in older and younger Biu lavas range
from 3918 to 4035, whereas Jos lavas range from 3930
to 3964. In contrast, megacrysts from both the Biu and
Jos Plateaux range to more radiogenic Pb isotope
Fig. 3. Sr–Nd isotope compositions of Biu and Jos Plateau lavas and
megacrysts along with literature data for oceanic and continental CVL
rocks with MgO >5 wt % [compiled from Halliday et al. (1988, 1990),
Lee et al. (1994), Marzoli et al. (1999, 2000) and Rankenburg et al.
(2004)]. Rocks from the Biu and Jos Plateaux span much of the range
of CVL lavas as a whole. Most lavas from the oceanic islands of St.
Helena, Tubuai and Mangaia plot in the field labelled ‘HIMU’. Isotopic systematics may be explained by mixing between a DMM
(depleted MORB mantle) component and a CC (continental crust)
and/or EM (enriched mantle) component.
compositions than their associated host lavas. Biu megacrysts are in the range of 2005–2090 in 206Pb/204Pb,
1571–1576 in 207Pb/204Pb and 3977–4107 in
208
Pb/204Pb (n ¼ 10). Jos megacrysts are in the range of
1982–2004 in 206Pb/204Pb, 1567–1570 in
207
Pb/204Pb and 3956–3979 in 208Pb/204Pb (n ¼ 3).
In combined plots of 207Pb/204Pb, 208Pb/204Pb, D7/4,
D8/4, eNd and 87Sr/86Sr vs 206Pb/204Pb (Figs 4 and
5) whole-rock samples from the Biu and Jos Plateaux
fall within the fields for oceanic and continental CVL
rocks previously reported by Halliday et al. (1988, 1990),
Lee et al. (1994) and Marzoli et al. (1999, 2000). Samples
with 206Pb/204Pb >198 lie close to the northern hemisphere reference line (NHRL) as defined by Hart (1984).
As far as can be concluded from the restricted dataset,
megacrysts from a single location cluster on a general
trend towards enriched compositions. For example, all
four cpx and three plagioclase megacrysts analysed from
the Miringa volcano (Biu Plateau) plot within a narrow
range only slightly greater than analytical uncertainty.
The isotopic trends defined by the megacrysts as a
whole overlap with the isotopic compositions of the
basalts, but extend to considerably more radiogenic Pb
isotope compositions. However, the megacrysts are different from typical HIMU OIB compositions such as St.
Helena (SH), Tubuai (T) or Mangaia (M), because of
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Fig. 4. Variation of 207Pb/204Pb (a) and 208Pb/204Pb (b) vs 206Pb/204Pb. Also shown are compositions of typical HIMU basalts from St. Helena
(SH), Tubuai (T) and Mangaia (M), and a mean of literature data for local continental crust (grey cross labelled CC). (a) Megacrysts (stars, cpx;
crosses, plag) extend to more radiogenic Pb isotope compositions than associated host lavas (squares, Jos; triangles, Biu young; diamonds, Biu old)
and lie close to the present-day NHRL. The bold dashed line gives location of the NHRL at 147 Ma, calculated using a simple two-stage model
starting at 443 Ga with Canyon Diablo lead and m0 ¼ 926. Internal differentiation at 177 Ga accounts for the NHRL. Fine dashed line
represents a reference line of 147 Ma age. (b) The source of the megacrysts evolved with similar k (¼ 232Th/238U) to the NHRL, and therefore is
different from typical HIMU compositions such as St. Helena, Tubuai or Mangaia. (c, d) Same data as in (a) and (b) but plotted using the delta
notation of Hart (1984), which represents the vertical deviation of a given data point in 207Pb/204Pb and 208Pb/204Pb from the NHRL multiplied
by a factor of 100.
their higher 208Pb/204Pb (Fig. 4b) and 87Sr/86Sr (Fig. 5b)
ratios for a given 206Pb/204Pb.
Osmium concentrations in the lavas vary from 19 to
350 (mean 125) ppt (Table 2), and are typical of those
found in OIB (Shirey & Walker, 1998). Rhenium concentrations, on the other hand, vary from 64 to 396
(mean 158) ppt and are low when compared with average mid-ocean ridge basalt (MORB) (926 ppt) or OIB
(377 ppt) (Righter & Hauri, 1998). Cu/Re ratios range
between 106 105 and 796 105. Cu/Re ratios higher
than the primitive mantle value of 107 105
(McDonough & Sun, 1995) and low Re concentrations
might be related to degassing of subaerial erupted lavas,
as suggested by Bennett et al. (2000) and Lassiter (2003)
for Hawaiian tholeiites. Re and Os concentrations do not
correlate with silicate-compatible elements such as Mg or
Ni, or with chalcophile elements such as Co or Cu. This
indicates that there is no direct relationship between Os
abundance and the degree of sample differentiation.
However, Re and Os concentrations are fairly well correlated (r2 ¼ 057).
The measured 187Os/188Os ratios of the lavas range
from 01234 to 02450. 187Re/188Os ranges from 33 to
274. Given the young age of the rocks (535–084 Ma for
older Biu Plateau lavas, other samples <50 ka), age corrections are generally small and within the analytical reproducibility. On a 187Os/188Os vs 1/[Os] diagram, the
volcanics do not form a significant correlation. However,
low concentrations (<70 ppt) are generally associated with
the highest 187Os/188Os ratios (01384–02450). There is a
broad negative correlation between 187Os/188Os ratios
and 206Pb/204Pb (Fig. 6), with radiogenic Os associated
with unradiogenic Pb compositions. Gannoun et al. (2001)
recently found a similar trend in Os–Pb isotope space in
CVL lavas from the continental sector. There is no significant trend in Os isotope composition in lavas with
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Fig. 6. Variation of 187Os/188Os vs 206Pb/204Pb. Samples with very
low Os concentrations (<50 ppt) are shown by dashed symbols. Samples with 206Pb/204Pb <198 (filled symbols) have 187Os/188Os up to
024, whereas samples with 206Pb/204Pb >198 (open symbols) scatter around typical values for OIB. The grey shaded area indicates
possible compositions for crustally (06–35 Ga, 187Re/188Os ¼ 10–
100) contaminated lavas. Assuming [Os] ¼ 50 ppt for average continental crust (Esser & Turekian, 1993), our Os data are best explained
by assimilation of material with 187Os/188Os 1.
Fig. 5. Variation of eNd (a) and 87Sr/86Sr (b) vs 206Pb/204Pb. In Pb–Sr
and Pb–Nd isotope space, the megacrysts extend the trend formed by
the lavas with 206Pb/204Pb >198. Other lavas are displaced towards
the composition of Pan-African continental crust. In subsequent diagrams we divide Biu and Jos Plateau rocks into two sets with different
coding: lavas that lie on a mixing trajectory from ‘A’ to ‘B’ are shown
by open symbols, whereas lavas that do not overlap with megacryst
compositions are shown by filled symbols.
206
Pb/204Pb >198. 187Os/188Os ratios range between
01266 and 01400, similar to typical OIB values. The
two samples from this subset with the highest
187
Os/188Os (01384 and 01400) are again those with the
lowest Os concentrations (48 and 66 ppt, respectively).
DISCUSSION
Origin of megacrysts
To use the composition of the megacrysts as a tracer for the
composition of CVL magmas at depth, we have to demonstrate the genetic link between them. Megacrysts have been
argued to be either genetically related (e.g. Green &
Ringwood, 1967; Irving, 1974; Irving & Frey, 1984; Liu
et al., 1992; Schulze et al., 2001) or xenocrystic phases (e.g.
Righter & Carmichael, 1993; Davies et al., 2001) in the past
and much of the discussion in the literature is based
upon comparison of isotope systematics or trace element
concentrations of lavas and megacrysts. However, petrological and petrographic evidence may be better suited to
constrain the origin of megacryst suites, because isotopic
compositions and trace element budgets of the host lavas
may be modified subsequent to megacryst formation.
Therefore simple comparison of the isotopic range of
megacrysts and host lavas does not provide a suitable
means to prove or disprove their cogeneity.
A complete discussion of our model for the genesis
of megacrysts in alkaline basalts from the Biu and Jos
Plateaux has been given by Rankenburg et al. (2004) and
the reader is referred to that work for more petrological
detail. In summary, major and trace element covariations
of the megacrysts are consistent with their derivation via
fractional crystallization from an evolving alkali basaltic
liquid. Megacrysts do not represent xenocrystic phases
derived from disaggregated peridotite- or pyroxenitexenoliths, because the latter are more calcic and
have significantly more depleted (lherzolite), or more
enriched (pyroxenite) trace element patterns for a given
mg-number than cpx megacrysts.
Megacrysts could potentially represent phenocryst
phases from an earlier magmatic event unrelated to
recent CVL volcanism. However, cpx–garnet intergrowths preserve magmatic textures and record magmatic temperatures of 1400 C (Rankenburg et al.,
2004). The absence of cooling features, such as recrystallization, diffusional gradients or exsolution lamellae,
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commonly found in pyroxenite xenoliths rules out the
interpretation that the megacrysts were precipitated by
an earlier magmatic event and stored in cooler
lithosphere for a significant amount of time. Based
upon modelling of concentration profiles of fast diffusing
species observed in cpx–garnet megacryst intergrowths
(Rankenburg et al., 2004), we concluded that the
megacrysts were not stored within the SCLM for more
than a few hundred years. Therefore, we propose that
the megacrysts derive from magmas related to the
recent ( 5 Ma to present) magmatism on the Biu and
Jos Plateaux.
Pressure estimates for crystallization of primitive cpx
and garnet megacrysts based upon phase relations in
alkaline magmas (Bultitude & Green, 1971), the pMelts
code (Ghiorso et al., 2002) and cpx–liquid thermobarometry (Putirka et al., 1996) are 17–23 GPa. Some
ilmenite megacrysts intergrown with more evolved cpx
compositions contain tiny trachyandesite melt inclusions.
Assuming that the melt inclusions and evolved cpx were
in equilibrium, we calculate Peq and Teq of 136 GPa and
1160 C (Putirka et al., 1996). This estimate is consistent
with the temperature calculated from the most evolved
cpx–garnet intergrowth, which yields 1100 C. The
crust–mantle boundary beneath the Biu and Jos Plateaux
is constrained from seismic data to depths of 28–30 6 km or 09 02 GPa (Poudjom-Djomani et al., 1995).
This requires that the megacrysts grew well below the
crust–mantle boundary, and we can therefore take the
megacrysts as probes of melt evolution within the subcontinental lithospheric mantle.
Although plagioclase is typically considered a lowpressure phase, there is one experiment on a natural
amphibolite that produced garnet þ cpx þ ab-rich
plagioclase coexisting with an andesitic melt at 18 GPa
(Rushmer, 1993). The compositions of experimentally
produced crystals (cpx, garnet, plag) in that experiment
match those of the megacrysts closely, thus permitting
plagioclase precipitation within the mantle. One feature
of such high-pressure plagioclase is its low anorthite
content, consistent with the observed composition of the
plagioclase megacrysts.
Mixing arrays in Sr–Nd–Pb isotope space
In plots of Nd–Pb and Sr–Pb isotopes (Fig. 5) the lavas
define triangular fields and can be explained by mixing
of three different end-member compositions forming
two linear arrays. The junction of both arrays is
defined by a cluster of lavas with 206Pb/204Pb
1982, 207Pb/204Pb 1564, 208Pb/204Pb 3953,
eNd 70 and 87Sr/86Sr 070290. We will refer to
this composition in the following discussion as component ‘A’. CVL lavas with a similar Sr–Nd–Pb isotope
composition have mantle-like d 18O values of 55%,
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whereas samples with higher 87Sr/86Sr extend to
higher d 18O values (Halliday et al., 1988). La/Yb,
Zr/Nb, Ce/Pb and K/U ratios of component ‘A’
lavas closely match the ratios reported for HIMUtype OIB (e.g. Thirlwall, 1997). Because component
‘A’ is prominent in the most primitive samples and is
common to both trends, we suggest that component ‘A’
has a sub-lithospheric, plume-like origin.
The second end-member (component ‘B’) is best represented by the most radiogenic cpx megacryst of each
plateau. An important observation from Fig. 5 is that
the megacrysts extend the mixing array towards component ‘B’ defined by the lavas to considerably higher
206
Pb/204Pb, 207Pb/204Pb, 208Pb/204Pb, 87Sr/86Sr and
lower eNd. Another observation is that cpx megacrysts
from the Biu and Jos Plateaux do not overlap in all
isotopic systems. Jos cpx megacrysts are characterized
by a lower 206Pb/204Pb for a given eNd and 87Sr/86Sr
when compared with Biu megacrysts (Fig. 5).
All rocks from the Biu and Jos Plateaux with
206
Pb/204Pb <198 seem to fan out to compositions that
lie within the field defined by literature data for Nigerian
basement rocks [granulites, gneisses, migmatites and
granites analysed by Halliday et al. (1988), Dickin et al.
(1991) and Dada et al. (1995)]. We will refer to this
composition in the following discussion as component
‘CC’ (see Figs 4 and 5). It is important to note that the
megacrysts do not plot on this trend. In the following
section, we evaluate the nature of the three components ‘A’, ‘B’ and ‘CC’ by comparison of the lava
compositions of the Biu and Jos Plateaux with the isotopic
compositions of genetically related megacrysts that grew
within the lithospheric mantle.
Comparison of megacrysts and host lavas
As pointed out above, the megacrysts in Figs 4 and 5
extend to more radiogenic lead isotope compositions
than their host lavas. In principle, this might be related
to radiogenic ingrowth after formation of the megacrysts,
or alternatively diffusional re-equilibration of the megacrysts with more radiogenic wallrocks. If the more radiogenic signature of the megacrysts was related to
radiogenic ingrowth, we would expect a strong correlation of Pb isotope ratios with their parental U/Pb ratios.
However, we observe that cpx and plagioclase megacrysts from a single volcano are indistinguishable in Pb
isotope composition despite their very different m values
( 0 and 8; see Lee et al., 1996). We therefore rule out
the possibility that the radiogenic Pb signature of the
megacrysts is due to late-stage radiogenic ingrowth.
Diffusional re-equilibration with surrounding radiogenic wallrock could in theory explain the observed offset
in lead isotope space. However, given the relatively short
time of possible storage in the SCLM, the slow diffusion
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CAMEROON VOLCANIC LINE LAVAS
coefficient of Pb in cpx (Cherniak, 2001) and the large
size of the megacrysts (up to several centimetres across),
we have to invoke 206Pb/204Pb ratios of the assimilated
component that are far outside the range reported for
mantle xenoliths from the CVL (Lee et al., 1996). Diffusive re-equilibration of the cpx and plagioclase megacrysts with radiogenic lithospheric mantle therefore
cannot account for their higher Pb isotope signatures
than the lavas. Because neither radiogenic ingrowth nor
diffusive re-equilibration subsequent to megacryst formation can account for the radiogenic Pb isotope signatures
of these phases, the Pb isotope compositions of cpx and
plagioclase megacrysts most probably reflect the source
compositions of the magmas from which they grew.
Therefore, we conclude that either the megacrysts grew
from melts that had been contaminated by a radiogenic
component, or the host lavas were contaminated with an
unradiogenic component after megacryst crystallization.
Evidence for assimilation of SCLM in the
genesis of lavas from the Biu and Jos
Plateaux
There is evidence for a combined assimilation–fractional
crystallization (AFC) process in the lavas lying along the
trend from component ‘A’ to ‘B’, i.e. in lavas with
206
Pb/204Pb >198 (Fig. 7). In a plot of Ni concentration
vs CaO/Al2O3 ratio (Fig. 7a) we identified two samples
with elevated Ni contents. Because these two samples also
show high modal abundance of disequilibrium [KDol/liq
(Fe–Mg) <03] olivine phenocrysts, we conclude that
these two melts have accumulated olivine phenocrysts
in a late stage of their history. After correcting the MgO
contents of the lavas for olivine accumulation, we see
clear correlations between decreasing MgO and increasing 206Pb/204Pb ratios (Fig. 7b) and decreasing eNd,
respectively (Fig. 7c). This suggests that the temporal
AFC progression proceeded from ‘A’ to ‘B’ and not vice
versa.
Assimilation and fractional crystallization are almost
always coupled because the heat needed to generate an
anatectic wallrock melt must be balanced by cooling and
the latent heat of crystallization (cumulate formation) of
the magma that is being contaminated (e.g. DePaolo,
1981). More sophisticated energy-constrained AFC calculations (Bohrson & Spera, 2001; Spera & Bohrson,
2001) all yield ratios of anatectic melt added to mass of
cumulates formed <1. Although assimilation may not
lead to a decrease in MgO (depending on the MgO
content of the partial melt), fractional crystallization
will, and this is likely to be the dominating process seen
in Fig. 7.
The P–T conditions of formation of the Biu and Jos
Plateaux megacrysts indicate that the megacrysts formed
within the underlying SCLM (Rankenburg et al., 2004).
Fig. 7. Evidence for AFC in the basalts with 206Pb/204Pb >198 (open
symbols). (a) Ni (ppm) vs CaO/Al2O3 weight ratio. Although most
samples lie on a high-pressure cpx control line, two samples have
elevated Ni concentrations, consistent with late-stage olivine accumulation. Assuming that cumulus olivine has 3000 ppm Ni and 498 wt %
MgO, we calculate 145 and 28 wt % accumulation, which will not
change the isotopic composition of the magmas. (b) 206Pb/204Pb and
(c) eNd vs MgO (wt %). The MgO-corrected compositions (dashed
lines) are consistent with assimilation of a high 206Pb/204Pb–low eNd
component, and MgO decreases as a result of fractionation of cpx
megacrysts.
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We therefore conclude that component ‘B’ is also derived
from the SCLM and not from the continental crust. This
assumption is further supported by the Os isotope systematics of the lavas. The SCLM and melts derived
therefrom on average should have subchondritic to chondritic 187Os/188Os ratios. Average continental crust, on
the other hand, is highly radiogenic [e.g. 187Os/188Os of
105 (Peucker-Ehrenbrink & Jahn, 2001)]. Assimilation
of SCLM or continental crust therefore should lead to
different trends in 187Os/188Os when plotted versus
indices of magma differentiation. Figure 6 shows that
187
Os/188Os ratios of lavas with 206Pb/204Pb >198
scatter around typical OIB values of 0130. The sample
with the highest 206Pb/204Pb (Mir) also shows somewhat elevated 187Os/188Os of 01384. However, this sample
also has very low Os concentration (48 ppt). Therefore it
remains unclear whether Mir defines a general trend
towards more radiogenic 187Os/188Os, possibly because
of minor incorporation of a HIMU component, or if Mir
is affected by small amounts of crustal contamination, too
small to be detected by the other isotopic systems. The
latter conclusion is supported by the fact that Mir shows
small fragments of felsic material in thin section.
However, all samples with 206Pb/204Pb >198 are within
the range of typical OIB basalts and do not extend to the
radiogenic values seen in samples with 206Pb/204Pb
<198 (187Os/188Os up to 02450). The lack of a
significant trend towards ‘crustal’ 187Os/188Os further
argues against a crustal origin for component ‘B’.
Component ‘B’ also does not appear to be a typical
HIMU melt, because this component is different from
the latter in that it has higher 208Pb/204Pb and 87Sr/
86
Sr, and lower 187Os/188Os for a given 206Pb/204Pb
(Figs 4–6). Lee et al. (1996) reported radiogenic Pb
(206Pb/204Pb ¼ 210) in a harzburgitic SCLM xenolith
from the Biu Plateau. Therefore, we propose that component ‘B’ represents metasomatically enriched SCLM.
Plume-derived melts (component ‘A’) assimilated variable
quantities of enriched SCLM or SCLM-derived melts or
fluids en route to the surface. Megacrysts extend to more
contaminated compositions than the magmas because
they derive from magmas that ponded in the lithosphere
for longer periods, thus allowing greater assimilation of
SCLM-derived material. The tiny melt inclusions found
in ilmenite megacrysts provide a suitable parental liquid
for the most evolved megacrysts in terms of their
highly enriched major and trace element chemistry
(Rankenburg et al., 2004). However, an isotopic study of
the melt inclusions was not possible because of their small
size. We previously suggested that the compositions of the
melt inclusions are not sampled at the surface as individual lavas because they mix with fresh batches of primitive magma, which trigger ascent and eruption of the
magma–megacryst mixture to the surface (Rankenburg
et al., 2004).
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We have shown in Fig. 5 that cpx megacrysts from
the Biu and Jos Plateaux do not overlap in all isotopic
systems. However, the similar trace element patterns of
primitive cpx megacrysts from the Biu and Jos Plateaux
(Rankenburg et al., 2004) suggest a common source
magma for both areas. Therefore, the three Jos Plateau
cpx megacrysts seem to define a different contamination
vector in combined Nd–Sr–Pb isotope space than Biu
megacrysts. Assuming that the process of megacryst genesis is similar in both areas, we propose that the lithospheric mantle underlying the Biu Plateau is more
radiogenic in Pb isotopes compared with that beneath
the Jos Plateau, but similar in Sr–Nd isotopes. Additional
analyses of peridotite xenoliths from the Biu and Jos
Plateaux are needed to better constrain the regional
variations in lithospheric mantle composition beneath
the continental sector of the CVL.
Models for the origin of enriched SCLM
Lee et al. (1996) reported enriched isotopic signatures
(87Sr/86Sr ¼ 0704156, eNd ¼ 04, 206Pb/204Pb ¼
2100) in an SCLM-derived harzburgite xenolith collected on the Biu Plateau. Models for the enrichment of
the SCLM as a result of underplated plumes or through
metasomatism by asthenosphere-derived melts have been
previously proposed by, for example, Ringwood (1982),
Hawkesworth et al. (1984), Stein & Hofmann (1992) and
Halliday et al. (1995). Halliday et al. (1990) suggested that
the high 206Pb/204Pb anomaly focused at the CVL c.o.b.
was inherited from relatively recent U/Pb fractionation
at 125 Ma during impregnation of the uppermost
mantle by the St. Helena hotspot when the Equatorial
Atlantic opened. A better estimate of the timing of
enrichment of the lithosphere underlying the Biu Plateau
might be 147 Ma, based upon the earliest period of magmatic activity in the northern Benue Trough (Coulon
et al., 1996). The overall CVL HIMU signature was
therefore proposed to be derived from recent radiogenic
ingrowth in the lithospheric mantle with variable but
high U/Pb.
If the radiogenic Pb isotope signature of the CVL was
derived from a lithospheric mantle source that was variably fractionated in U/Pb at 147 Ma, we would expect
the present-day samples to lie along a corresponding
147 Ma isochron in 206Pb/204Pb–207Pb/204Pb isotope
space. However, we observe that component ‘A’ and
the enriched lithospheric component ‘B’, as represented
by megacryst compositions, both cluster close to the
NHRL, which defines an apparent ‘age’ of 177 Ga
[using a gradient of 01084 and an intercept of 13491
from Hart (1984)]. The 147 Ma reference line indicated
in Fig. 4a has a considerably shallower slope than the
NHRL. Internal differentiation of the SCLM at 147 Ma
therefore cannot account for the Pb isotope signatures
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CAMEROON VOLCANIC LINE LAVAS
observed in the lavas and megacrysts of the Biu and Jos
Plateaux. An alternative model is that the radiogenic Pb
isotopes in the sub-CVL SCLM could derive from
HIMU-type magmas derived from the fossil St. Helena
hotspot, which at the time of opening of the South
Atlantic was located in the region occupied by the present
c.o.b. However, as already pointed out by Halliday et al.
(1990), there is a discrepancy between the isotopic data
for typical HIMU and other CVL lavas, in that the
former have lower 208Pb/204Pb and 87Sr/86Sr for a
given 206Pb/204Pb and do not form an appropriate mixing end-member for formation of the megacryst source
magmas.
We propose that the enriched isotopic signature of the
SCLM simply represents older (Proterozoic?) lithosphere
that was subjected to multiple metasomatic events rather
than a single overprint in the Mesozoic. On average,
long-term metasomatism produced Pb isotope signatures
lying along the NHRL. Sr and Nd isotopes, however,
may be more radiogenic compared with typical MORB
because of the elevated Rb/Sr and U/Pb and low Sm/
Nd ratios of the percolating melts. Regardless of the
origin of the radiogenic Pb isotopes in the sub-CVL
SCLM, the presence of mantle xenoliths with elevated
206
Pb/204Pb indicates that portions of the SCLM
beneath the Cameroon Volcanic line do possess isotopic
characteristics similar to our proposed component ‘B’.
The correlation of Pb and Nd isotopes with MgO in the
lavas with 206Pb/204Pb >198 strongly suggests that this
component is related to assimilation of such material in
the SCLM.
Evidence for crustal contamination of lavas
from the Biu and Jos Plateaux
Lavas that do not lie on a mixing trend between primary
composition ‘A’ and the inferred lithospheric mantle
contaminant ‘B’ form broad trends pointing towards a
third end-member characterized by high 87Sr/86Sr, low
eNd, low 206Pb/204Pb, high D7/4, high D8/4 and high
187
Os/188Os (Figs 4–6). Previous publications reported
evidence for interaction with the continental crust in
some evolved phonolites and trachytes of the continental
sector based upon 87Sr/86Sr ratios as high as 0705–
0714 (Marzoli et al., 1999) and large variations in Hf
isotopes (Ballentine et al., 1997). The Biu Plateau lies
mainly on granite, gneisses and charnockitic rocks of the
Pan-African basement complex of Nigeria, and on the
Cretaceous Bima sandstones of the Benue trough in
the north and west. Radiometric ages of Nigerian basement rocks have been presented by, for example, Van
Breemen et al. (1977), Dickin et al. (1991), Dada (1998)
and Kr€
oner et al. (2001). Ages cluster around 35, 31–30,
27–25, 21–18 Ga and there is a major tectonometamorphic imprint on most Nigerian basement rocks in
Pan-African ( 600 Ma) times. On average, these rocks
are characterized by high 87Sr/86Sr, low eNd, low
206
Pb/204Pb, high D7/4 and high D8/4, and therefore
provide a suitable candidate for composition ‘CC’. Megacrysts or xenoliths with these compositions have not been
observed. We therefore suggest that the trend towards
component ‘CC’ observed in the lavas was imposed on
the magmas at shallower depths, after formation of the
megacrysts. This may be due to assimilation of very
shallow SCLM or the continental crust.
Os isotopes provide a means for distinguishing if component ‘CC’ derives from the SCLM or the continental
crust. If the SCLM is responsible for the trends observed,
we expect a trend towards lower or constant 187Os/188Os
with increasing 87Sr/86Sr or decreasing 206Pb/204Pb,
respectively. Contamination with continental crust, on
the other hand, should lead to significantly higher
187
Os/188Os. Os isotopes are generally correlated with
206
Pb/204Pb in lavas with 206Pb/204Pb <198 (Fig. 6),
trending towards higher 187Os/188Os with decreasing
206
Pb/204Pb. The 187Os/188Os ratios in lavas falling
along the ‘A–CC’ trend extend to ratios much higher
than observed in the vast majority of OIB and mantle
derived xenoliths. Because Nigerian basement rocks are
also characterized by high D7/4 ratios, crustal contamination should also result in increasing D7/4 in the lavas
falling along the ‘A–CC’ trend. Figure 4 shows our data
(symbols) along with the mean composition of local PanAfrican continental crust. The lavas with 206Pb/204Pb
<198 clearly point towards the composition of continental crust. The trend towards elevated 187Os/
188
Os and D7/4 with decreasing 206Pb/204Pb strongly
suggests that component ‘CC’ derives from assimilation
of continental crust.
To quantify the amount of crustal contamination in
the lavas, we used the mean composition of Nigerian
basement rocks found in the literature (Halliday et al.,
1988; Dickin et al., 1991; Dada et al., 1995), which is
similar to mean upper continental crust (Taylor &
McLennan, 1995), but significantly more enriched in
the REE. Sr, Nd, Pb and Os concentrations of the
uncontaminated end-member were taken as the mean
of type ‘A’ lavas. Mixing can successfully be modelled
via bulk-rock assimilation. The results for Pb, Sr and Nd
isotopes are given in Fig. 8 and are consistent with assimilation of 8% crust in the most contaminated lavas.
There have been only a limited number of Re–Os
isotopic studies on lower- and upper-crustal rocks. As
yet, there are no 187Os/188Os data available for Nigerian
basement rocks. Therefore, modelling of Os isotopes is
not as straightforward as for Sr, Nd and Pb isotopes. The
continental crust, however, generally has 187Re/188Os
ratios ranging between 10 and 100 (e.g. Esser &
Turekian, 1993; Saal et al., 1998; Peucker-Ehrenbrink &
Jahn, 2001). For the purposes of this paper, we modelled
185
JOURNAL OF PETROLOGY
8
VOLUME 46
8
(a)
(b)
6
2
εNd
2
6
4
6
5
4
6
5
8
4
0.7028
0.7032
0.7034
87Sr/86Sr
0.7036
0.7038
70
(c)
0.7036
87Sr/86Sr
0.7030
8
0.7034
0.7032
0.7030
0.7028
19.0
4
19.0
8
50
6
30
4
10
2
19.5
19.5
20.0
20.5
20.0
20.5
206Pb/204Pb
(d)
∆ 8/4
0.7038
JANUARY 2005
7
7
εNd
NUMBER 1
8
6
4
2
-10
20.0
20.5
206Pb/204Pb
-30
19.0
19.5
206Pb/204Pb
Fig. 8. Quantitative Sr–Nd–Pb isotope model of crustal contamination of composition ‘A’ using a mean literature dataset for the composition of
the local crust with [Sr] ¼ 351 ppm, [Nd] ¼ 667 ppm, [Pb] ¼ 214 ppm, 87Sr/86Sr ¼ 07256, eNd ¼ 162, 206Pb/204Pb ¼ 1813, D7/4 ¼ 151
and D8/4 ¼ 128 (see text for references). Symbols as in Figs 6 and 7. The results are mutually consistent in combined Sr–Nd–Pb isotope systems,
with maximum amounts of 8% assimilation.
present-day 187Os/188Os ratios of Nigerian basement
rocks using this range of Re/Os ratios and mean ages of
the basement rocks ranging from 06 to 35 Ga (Dada,
1998). As a result we expect present-day 187Os/188Os
ratios of the Nigerian basement rocks ranging from
0225 to 61. If we assume a mean age of 2 Ga for the
basement rocks and a mean upper-crustal 187Re/188Os
ratio for these rocks of 345 (Peucker-Ehrenbrink & Jahn,
2001), we calculate a present-day 187Os/188Os ratio of
128. A mixing scenario for these three cases is included
in Fig. 6. In Pb–Os isotope space the lavas from the Biu
and Jos Plateaux with 206Pb/204Pb ratios <198 fall
within the field of crustal contamination. Most of the
samples in Fig. 6 scatter around a mixing trajectory
calculated for an assimilant with 187Os/188Os of 128.
Although mixing calculations including Os isotopes do
not allow quantitative estimations of the amount of
crustal contamination, the negative correlation of Os
and Pb isotopes, with Os isotopes extending to radiogenic
values greatly exceeding those found in primitive OIB,
strongly supports the hypothesis that lavas with
206
Pb/204Pb <198 were contaminated
Pan-African continental crust.
by
local
Implications for crustal contamination in
the oceanic CVL
The above discussion strongly suggests that crustal contamination has affected some continental CVL lavas,
particularly those lavas with 206Pb/204Pb <198. However, it has previously been noted that the ranges of Sr,
Nd, Hf and Pb isotopes in CVL lavas from the
continental and oceanic sectors are similar. This observation has been used to argue against a significant role for
crustal contamination in the continental CVL lavas
(Fitton & Dunlop, 1985; Halliday et al., 1988; Ballentine
et al., 1997). Most lavas from the oceanic portion of the
CVL have Pb isotope compositions that plot near the
NHRL. However, a subset of samples from Principe
and Annobon extend to anomalously high D7/4 and
D8/4 values (Fig. 4c and d). If the high D7/4 and D8/4
values (and high 187Os/188Os) in the Biu and Jos Plateau
186
RANKENBURG et al.
CAMEROON VOLCANIC LINE LAVAS
lavas reflect crustal contamination, this raises the question of how such features can be present in samples from
the oceanic portion of the CVL.
Historically, the ‘continental’ signature present in lavas
from both the continental and oceanic portions of the
CVL has been interpreted as reflecting the presence of an
EM-type component in the plume source of these lavas.
However, most EM-type lavas such as those from the
Society and Cook–Austral Islands are characterized
by relatively unradiogenic Os isotopes (187Os/188Os
0135), in contrast to the crust-contaminated Biu and
Jos Plateau lavas. It is possible that the mantle source of
CVL lavas contains an EM component that is coincidentally similar in composition to the local Pan-African crust
and that this component is responsible for the ‘continental’ signature in some oceanic CVL lavas. However, we
propose that blocks of Pan-African crust may have
become embedded in the oceanic crust during initiation
of continental breakup. Consequently, contamination of
mantle-derived magmas at shallow levels by continental
crust may be responsible for the ‘continental’ signature in
lavas from the oceanic portion of the CVL as well as in
those from the continental sector.
There is growing evidence that fragments of continental crust may be more prevalent in the ocean basins than
previously believed. Fragments of continental crust have
been drilled or dredged at several locations along the
Mid-Atlantic Ridge (e.g. Bonatti et al., 1996; Belyatsky
et al., 1997; Pilot et al., 1998). Schaltegger et al. (2002) have
also reported ancient continental zircon xenocrysts in
basalts from Iceland and Mauritius, and suggested the
presence of either rafted or shallowly subducted continental crust under these islands. In the Indian Ocean,
fragments of continental crust have been recovered from
portions of the Kerguelen Plateau (Ingle et al., 2002), and
the isotopic and trace element compositions of many
Kerguelen Plateau basalts appear to record assimilation
of continental crust (Frey et al., 2002; Ingle et al., 2003).
Thus even in the ocean basins the possibility of continental crustal contamination of lavas cannot be excluded
a priori. Not all isotopic variations in OIB necessarily
reflect variations in mantle composition.
At present, there is no direct evidence for the presence
of continental fragments underneath the islands of
Principe or Annob
on. Future Os isotope studies of lavas
from these islands should provide one test for the crustal
contamination model. Gannoun et al. (2001) reported
187
Os/188Os values up to 01876 in basalts from the
oceanic portion of the CVL. If high 187Os/188Os values
correlate with high D7/4 and D8/4 values on these
islands, as in the case for the Biu and Jos Plateau lavas,
this would further strengthen our contention that local
crustal contamination is the source of the anomalous Pb
isotope signature in some Principe and Annobon
lavas. Additional petrographic studies should be used to
determine the presence or absence of xenocrystic zircons
or other fragments of continental crust in oceanic CVL
lavas.
Finally, we note that the refractory ‘LOMU’ component in Mid-Atlantic Ridge basalts described by Douglass
et al. (1999) and Douglass & Schilling (2000) is isotopically
similar to Pan-African continental crust. Douglass et al.
(1999) proposed the ‘LOMU’ component to be delaminated subcontinental lithospheric mantle dispersed into
the upper mantle during the breakup of Gondwana.
However, the most pronounced ‘LOMU’ compositions
sampled in oceanic basalts are found in tholeiites from the
Aphanasey Nikitin Rise (Douglass et al., 1999), which
were recently proposed by Borisova et al. (2001) to be
contaminated via shallow assimilation of continental
crust derived from cratonic Gondwanian lithosphere
based upon combined major and trace element and isotopic data. We propose that the South Atlantic ‘LOMU’
signature in general may not correspond to refractory
delaminated SCLM, but instead reflects shallow assimilation of continental crust that became trapped in the
oceanic lithosphere during continental breakup in the
Mesozoic.
CONCLUSIONS
We have demonstrated in this paper that the isotopic
compositions of megacrysts, which are argued to be
genetically related to recent CVL volcanism, allow us to
identify and distinguish the lithospheric modifications
imprinted on two suites of CVL alkaline intraplate volcanics. Jos and Biu Plateau lavas seem to have a
homogeneous asthenospheric source with 206Pb/204Pb
198, D7/4 and D8/4 0, eNd 7, 87Sr/86Sr 07029, 187Os/188Os 0129 and d 18O values of
55%. Magmas subsequently interacted with either
enriched SCLM via melt–melt mixing, leading to
increasing 206Pb/204Pb, and/or continental crust, leading to decreasing 206Pb/204Pb and increasing
187
Os/188Os. The SCLM underlying the Biu Plateau is
characterized by high 206Pb/204Pb 210, whereas the
Jos Plateau SCLM probably has 206Pb/204Pb 20.
Although quantitative modelling of lithospheric contamination is hampered by too many unknown parameters,
crustal contamination is well constrained and is of the
order of 8% for the most contaminated lavas of both the
Biu and Jos Plateaux. Assuming that the continental and
oceanic sector of the CVL are fed by a common and
relatively homogeneous asthenospheric source, we infer
that the similar contamination trends seen in some oceanic CVL lavas are also caused by shallow assimilation of
crustal material. Furthermore, we suggest that the South
Atlantic ‘LOMU’ signature of Douglass et al. (1999) may
also be caused by assimilation of rafted blocks of continental crust rather than refractory delaminated SCLM.
187
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VOLUME 46
ACKNOWLEDGEMENTS
We are grateful to Al Hofmann and John Snow for their
constructive criticism on this paper. Wafa Abouchami
and Sieglinde Bederke-Raczek are acknowledged for
their assistance at the Max-Planck-Institut, Mainz.
Sincere thanks are given to Godfrey Fitton for sending
me some unpublished rock analyses of the Biu Plateau.
Gareth Davies, Der-Chuen Lee, Godfrey Fitton and Dan
Barfod are thanked for their constructive comments on
an earlier version of the manuscript. This work was
financially supported by Deutsche Forschungsgemeinschaft project no. BR 1012/11-1 and by the MaxPlanck-Institut, Mainz.
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