JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
PAGES 1309–1344
2005
doi:10.1093/petrology/egi016
Volcanism in the Vitim Volcanic Field,
Siberia: Geochemical Evidence for a Mantle
Plume Beneath the Baikal Rift Zone
J. S. JOHNSON1*, S. A. GIBSON1, R. N. THOMPSON2 AND
G. M. NOWELL2
1
DEPARTMENT OF EARTH SCIENCES, UNIVERSITY OF CAMBRIDGE, DOWNING STREET, CAMBRIDGE CB2 3EQ , UK
2
DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF DURHAM, SOUTH ROAD, DURHAM DH1 3LE, UK
RECEIVED JUNE 23, 2003; ACCEPTED JANUARY 18, 2005
ADVANCE ACCESS PUBLICATION MARCH 4, 2005
The Baikal Rift is a zone of active lithospheric extension adjacent to
the Siberian Craton. The 6–16 Myr old Vitim Volcanic Field
(VVF) lies approximately 200 km east of the rift axis and consists
of 5000 km3 of melanephelinites, basanites, alkali and tholeiitic
basalts, and minor nephelinites. In the volcanic pile, 142 drill core
samples were used to study temporal and spatial variations. Variations in major element abundances (e.g. MgO ¼ 33–146 wt %)
reflect polybaric fractional crystallization of olivine, clinopyroxene
and plagioclase. 87Sr/86Sri (07039–07049), 143Nd/144Ndi
(05127–05129) and 176Hf/177Hfi (02829–02830) ratios
are similar to those for ocean island basalts and suggest that the
magmas have not assimilated significant amounts of continental
crust. Variable degrees of partial melting appear to be responsible
for differences in Na2O, P2O5, K2O and incompatible trace element
abundances in the most primitive (high-MgO) magmas. Fractionated heavy rare earth element (HREE) ratios (e.g. [Gd/Lu]n >
25) indicate that the parental magmas of the Vitim lavas were
predominantly generated within the garnet stability field. Forward
major element and REE inversion models suggest that the
tholeiitic and alkali basalts were generated by decompression melting
of a fertile peridotite source within the convecting mantle beneath
Vitim. Ba/Sr ratios and negative K anomalies in normalized multielement plots suggest that phlogopite was a residual mantle phase
during the genesis of the nephelinites and basanites. Relatively high
light REE (LREE) abundances in the silica-undersaturated melts
require a metasomatically enriched lithospheric mantle source.
Results of forward major element modelling suggest that melting
of phlogopite-bearing pyroxenite veins could explain the major
element composition of these melts. In support of this, pyroxenite
xenoliths have been found in the VVF. High Cenozoic mantle
potential temperatures ( 1450 C) predicted from geochemical
modelling suggest the presence of a mantle plume beneath the Baikal
Rift Zone.
*Corresponding author. Present address: British Antarctic Survey,
High Cross, Madingley Road, Cambridge CB3 0ET, UK. Telephone:
þ44 (0)1223 221313. Fax: þ44 (0)1223 361616. E-mail: [email protected]
# The Author 2005. Published by Oxford University Press. All
rights reserved. For Permissions, please email: journals.permissions@
oupjournals.org
Baikal Rift; mafic magmatism; mantle plume; metasomatism; partial melting
KEY WORDS:
INTRODUCTION
Magmatic activity frequently accompanies continental
rifting, and is generally considered to be the consequence
of adiabatic decompression melting of the convecting
mantle as it upwells beneath the thinned lithosphere
(McKenzie & Bickle, 1988). Melting processes beneath
continental rifts are not, however, as well understood as
those beneath mid-ocean ridges where upwelling of the
mantle is a passive response to lithospheric extension.
Beneath continental rifts, it is sometimes unclear as to
whether asthenospheric mantle upwelling is an active or
passive process. Active rifting is initiated and driven by
the impingement of hot asthenospheric mantle (i.e. a
mantle plume) on the base of the overlying continental
lithosphere. This can cause lithospheric thinning and
uplift. Passive rifting is driven by tensional forces that
extend the lithosphere and thin it, resulting in upwelling
of the asthenospheric mantle. In reality, however, there
are difficulties in differentiating between active and passive rifting processes because passive rifting may eventually produce the same geophysical and geological effects
(e.g. alkaline magmatism, lithospheric extension and
JOURNAL OF PETROLOGY
VOLUME 46
asthenospheric upwelling) as active rifting (Ruppel,
1995). It is probable that processes occurring in the
evolution of most continental rifts fall between these two
extremes, and there may also be a temporal change from
passive to active rifting in some regions (e.g. Delvaux
et al., 1997).
The asthenospheric and lithospheric mantles are both
known to be important melt source regions during continental rifting (e.g. Leat et al., 1988; Thompson et al.,
1990). The volume and composition of the magmas generated during rifting have been shown to be directly
related to the degree and duration of extension, the
temperature of the underlying asthenospheric mantle
(McKenzie & Bickle, 1988) and the presence or absence
of volatile/hydrous phases within the mantle source
(Gibson et al., 1993). Extension and thinning of the lithosphere cause adiabatic decompression melting of the
asthenospheric mantle. Additionally, melting within the
lithospheric mantle may occur if it has been previously
metasomatized. Heat conduction and/or heat advection
by melts derived from the convecting mantle (McKenzie,
1989) may cause partial melting of volatile and K-rich
veins within the lithosphere, resulting in magmas that are
enriched in incompatible trace elements. Evidence for
such a process is widely documented in continental rifts.
For example, in the Rio Grande Rift of the western USA,
strongly potassic magmatism (lamproites and minettes)
on the rift flanks have been interpreted as the products
of re-melting small melt fractions that had previously
infiltrated the lithospheric mantle and solidified as veins
(Gibson et al., 1993). Wholesale melting of the lithospheric mantle could, potentially, occur if small melt
fractions were to react substantially with the surrounding
anhydrous lithosphere (e.g. Hawkesworth et al., 1992).
The Baikal Rift is one of the least studied regions of
currently active major lithospheric extension. In this
study, we focus on the petrography, mineral and wholerock chemistry of extension-related lavas from the Vitim
Volcanic Field (VVF) of the Baikal Rift, and develop a
petrogenetic model for magma generation processes operating during the formation of this major continental rift.
BAIKAL RIFT ZONE
The Baikal Rift is located within the Palaeozoic Sayan–
Baikal fold belt, near its boundary with the Siberian
craton (Fig. 1). The fold belt is a collage of terranes (e.g.
Precambrian microcontinents, fragments of oceanic crust
and island arcs) that were accreted onto the craton in the
Late Riphean and Early to Late Palaeozoic (Logatchev &
Zorin, 1992). The onset of rifting occurred in the Oligocene (35–30 Ma), with a major increase in the rate of
lithospheric extension in the Late Pliocene (<3 Ma).
This increased rate of extension was largely responsible
NUMBER 7
JULY 2005
for the present-day topography (Logatchev & Zorin,
1987). The S-shape of the rift zone appears to be controlled by the edge of the Siberian craton and the Sayan–
Baikal foldbelt (e.g. Molnar & Tapponnier, 1975); Lake
Baikal itself follows the southeastern edge of the craton
(Fig. 1).
Cenozoic igneous activity in the Baikal Rift Zone
(BRZ) is restricted to several small (<7000 km2) lava fields
to the east and south of Lake Baikal, with a total volume
of approximately 5000 km3 (Logatchev & Florensov,
1978). To the south and west of the southern end of
Lake Baikal, there is a diffuse zone of volcanism in East
Sayan and Tuva, extending into the regions of Hangai,
Gobi and the Mongolian Altai (Fig. 1). The igneous
activity in these regions has been described by Rasskazov
(1994) and Barry & Kent (1998). Within the BRZ, there
are four regions of Cenozoic volcanism, located to the
south and east of Lake Baikal (Fig. 1): Udokan Plateau,
Vitim, Hamar-Daban and Bartoy. The largest of the four
volcanic fields is Vitim, which is located approximately
200 km east of Lake Baikal. Hamar-Daban and Bartoy
are the smallest in area, and are found close to the
axial part of the rift that runs through Lake Baikal in a
NE–SW direction. Cenozoic volcanism is absent from the
central part of the rift system and most of its basins
(Ionov, 2002).
The BRZ has been the focus of numerous geophysical
studies. The results of a recent seismic investigation
(Suvorov et al., 2002) indicate that the depth to the
Moho along the rift axis varies between 35 and 50 km;
the crust is thinnest beneath the southern part of the rift.
Furthermore, a 200 km wide zone of crustal thinning
appears to extend NE from the rift axis to the VVF,
where the crust is 35 km thick (Suvorov et al., 2002). A
teleseismic study by Gao et al. (1994) indicates that there is
a broad zone of asthenospheric upwelling beneath Lake
Baikal and the region to the east. This was not observed
in the studies of Burov et al. (1994), Petit et al. (1998) or
Zhang (1998); Artemieva & Mooney (2001), however,
have shown a region of lithospheric thinning extending
200 km to the east of Lake Baikal and suggested that the
base of the thermal lithosphere (defined by them as the
depth to the 1300 C adiabat) is between 100 and 125 km
depth beneath the VVF.
Two hypotheses have been put forward to explain the
origin of the forces responsible for extension in the BRZ:
(1) Oligocene rifting in the Baikal region began 30–
35 Myr ago and was contemporaneous with the early
stages of the India–Asia collision. Molnar & Tapponnier
(1975) suggested that the collision was responsible for
most of the large-scale tectonics of Asia, and that this
may have caused rifting in the Baikal area. Recently,
Polyansky (2002) suggested that the rifting was caused by
the northward movement of the Indian plate into Eurasia,
the east–west convergence of the North American and
1310
JOHNSON et al.
96O
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
104O
100O
108O
LEGEND
112O
faults
120
O
Vitim river
r
ive
r
na
116O
Le
An
56
ga
ra
towns
O
Udokan
riv
er
volcanic fields
sedimentary basins
Vitim
SIBERIAN CRATON
EAST SAYAN
Chita
Irkutsk
L
KA
I
BA LT
N- BE
A
Y LD
SA FO
52O
Ulan-Ude
Hamar-Daban
Bartoy
Lake Baikal
Se
l
riv eng
er a
TUVA
RUSSIA
48O
Hangai
ALTAI
MONGOLIA
N
Dariganga
0
44O
400 km
Fig. 1. Map showing the location of the major Cenozoic volcanic fields of the BRZ and Mongolia (after Kiselev, 1987). The locations of Lake
Baikal, the Siberian craton, Sayan–Baikal fold belt and the late Oligocene to Quaternary sedimentary basins (Rasskazov, 1994) are also indicated.
The inset is a general map of East Asia; the white square indicates Lake Baikal and the surrounding area.
Eurasian plates and the southeastward extrusion of the
Amur plate into NE Asia.
(2) Several workers have discussed the possibility of
a mantle plume or plumes lying beneath the Baikal
region during the Cenozoic (Zorin, 1981; Logatchev &
Zorin, 1987, 1992; Kiselev & Popov, 1992; Windley &
Allen, 1993; Petit et al., 1998). Balijinnyam et al. (1993)
proposed that the Baikal and Mongolia regions are
underlain by a number of small plumes or ‘diapirs’ (with
diameters of <100 km), which were partly responsible
for the formation of the rifts. As well as reporting features
that suggest interaction of a mantle plume with the
lithosphere beneath Lake Baikal (e.g. high heat flow,
uplifted topography, lithospheric thinning and alkaline
magmatism), Windley & Allen (1993) discussed how
other observations are not consistent with stresses linked
to the India–Asia collision. For example, alkalic magmatism and high heat flow are not confined to localized
rifts, but extend across the Mongolian Plateau, implying
a laterally extensive heat source. A variety of rift
orientations are observed in the plateau. These are
inconsistent with east–west extension produced by the
India–Asia collision, as only a single rift orientation would
be expected.
VITIM VOLCANIC FIELD
A simplified geological map of the VVF is shown in
Fig. 2a, indicating the distribution of the main volcanic
centres. The igneous activity of the VVF may be divided
into two phases: the earlier, more voluminous, phase
occurred during Miocene to Pleistocene times (Kiselev,
1987), whereas the second phase of activity took place in
the Pleistocene and Holocene. The majority of the volcanic centres (Fig. 2a) are located in the northwestern
1311
JOURNAL OF PETROLOGY
54.2
VOLUME 46
NUMBER 7
JULY 2005
Pleistocene–Holocene lavas
20 km
Miocene–Pleistocene lavas
Palaeozoic granites
Cenozoic sediments
Maly Amalat
Ho
fault (tick on down-thrown side)
ygo
54.0
t
river
volcanic centre
Antasey
Ekzar
iver
Kandidushka
volcano
Vitim
r
53.8
tuff pit
riv
er
Yaksha
volcanoes
latitude (°)
53.4
Vi
tim
53.6
la
Ata
ng
N
a
(a)
54.2
Hoygot
54.0
4613
4569
4563
Antasey
Ekzar
3690
4404
4403
53.8
Burulzay
4059
4770
3043
4772
tuff pit
Muliha
Kol
ic
4490
3046
53.6
4363
4431
Centralni
4454
h ika
n
4426
4633
3889
4659
3833
4132
Atalanga
Bortovoy
4102
53.4
(b)
112
112.5
113
113.5
114
longitude (°)
Fig. 2. (a) Simplified geological map of the VVF, based on a compilation of data obtained during the course of fieldwork and previously
published maps (e.g. Ionov et al., 1993). The ‘tuff pit’ locality is the location of samples 93VBS 1, 4 and 399. (b) The locations of drill-holes in the
VVF. Drilling regions are bounded by dashed lines. Drill-hole 3690 is shown by an open symbol because its region is not known (although
this map suggests that it lies within the Antasey region). The coordinates for each drill-hole are given in Appendix 1, with the exception of Ekzar
3313, 3315 and 4072, Burulzay 4771 and Atalanga 4135, which are not known. This figure can be viewed in colour on Journal of Petrology
online.
1312
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
part of the volcanic field, in an area parallel to the axial
part of the rift zone.
There are several places in the VVF where mantle
xenoliths have been found within the eruptives; the
most well known is the ‘tuff pit’ of Ionov et al. (1993),
which has also been referred to as the ‘Bereya quarry’ by
Glaser et al. (1999) and the ‘picrobasalt quarry’ by Litasov
et al. (2000).
Cenozoic sediments occur mainly to the SW of the
VVF, close to the Atalanga and Vitim rivers (Fig. 2a).
Mesozoic trachybasalt and andesite lavas are also present
in this area. Granitic bodies, of inferred Palaeozoic age
(Litasov et al., 2000), form the basement to the VVF and
crop out further west towards Lake Baikal (Logatchev &
Zorin, 1992). These are widespread and form a considerable proportion of the underlying crust, extending to
depths of 20 km (Suvorov et al., 2002).
There are very few available age data for lavas from the
VVF. Esin et al. (1995) gave K–Ar ages ranging from 66
to 1065 Ma for four Cenozoic lavas from drill-hole 3313
(Ekzar) and three from drill-hole 4431 (Antasey), as well
as an estimate of 1625 Ma for the basalt hosting mantle
xenoliths at the ‘tuff pit’ locality (Ionov et al., 1993).
SAMPLING
Approximately 200 lava samples were taken from 60 mm
diameter drill-cores, as exposure is very limited because of
dense forest covering the plateau. Many of the drill-holes
reach the granitic basement, with one extending to a
depth of 700 m. Those in the Atalanga region intersect
Mesozoic lavas at a depth of 400 m. The drill-hole locations are given in Appendix A and are shown in Fig. 2b.
ANALYTICAL TECHNIQUES
Analyses of olivine, clinopyroxene, plagioclase and spinel
were made using a CAMECA SX-50 electron microprobe in the Department of Earth Sciences, University
of Cambridge. Energy-dispersive spectrometry (EDS)
was used to analyse SiO2, TiO2, Al2O3, FeO, MnO,
MgO, Na2O and Cr2O3 on carbon-coated, polished
thin sections of the samples. Operating conditions for
EDS were an accelerating voltage of 20 kV, a beam
current of 25 nA, a beam diameter of 1 micron, and a
live counting time of 60 s for each analysis. Calibration
was made with reference to a cobalt standard. On-line
peak stripping and corrections were performed using
Link Analytical ZAF4/FLS software. Wavelengthdispersive spectrometry (WDS) was used to analyse NiO
and CaO in olivine, using the same operating conditions
and Link Analytical software as for EDS. Calibration was
made with reference to pure nickel and wollastonite
standards.
At the University of Durham, 142 samples were
analysed for their whole-rock major and trace element
chemistry, by X-ray fluorescence (XRF) for major elements and, for trace elements, by inductively coupled
plasma mass spectrometry (ICP–MS). XRF analyses
were made using a Philips PW1400 X-ray fluorescence
spectrometer with a PW1500 72 automatic sample changer. For major elements, glass discs made from the
powdered sample fused with lithium tetraborate were
analysed; trace element analyses were carried out on
freshly pressed powder pellets. The international standards AGV-1 and DST-1 were used for calibration (Potts
et al., 1992), and analyses were repeated throughout
the run to monitor analytical precision. Rare earth and
trace element (Pb, Th, U, Nb, Ta, Zr, Hf, Y and Ba)
concentrations were determined using a Perkin-Elmer
SCIEX ELAN 6000 ICP–MS. Samples were prepared
by digestion with HF/HNO3 at the University of
Cambridge, following the method of Jarvis & Jarvis
(1992). Blanks were prepared with each batch of samples,
and analytical accuracy and reproducibility were estimated from measurements of international rock standards
GSP-1, BCR-2 and AGV-1. One standard and one blank
were analysed at several intervals throughout the whole
analytical run, to monitor signal drift and contamination
within the instrument.
Sr, Nd and Hf isotope ratios were determined using
a ThermoFinnigan Neptune Plasma Ionisation Multicollector Mass Spectrometer (PIMMS) at the University
of Durham. Nowell et al. (2003) have described the procedure for analysis of Sr, Nd and Hf on this instrument.
Separation of Sr, Nd and Hf for analysis was achieved
using a two-column procedure (Dowall et al., 2003).
Whole-rock geochemical and Sr–Nd–Hf isotope analyses
of representative samples from the VVF are given in
Tables 1 and 2, respectively (the complete dataset is
available as Electronic Appendices A and B, which may
be downloaded from the Journal of Petrology website at
http://www.petrology.oupjournals.org); the internal
errors on the isotope ratios are reported in Table 3.
The Vitim samples were analysed in two analytical sessions, and reproducibility is therefore given for Sr, Nd
and Hf in each session. Details of normalization values,
mass bias and standards used are given in the legend to
Table 3.
MAGMA TYPES
We have classified the Vitim lavas into sub-alkalic basalts,
alkali basalts, basanites, melanephelinites and nephelinites (Fig. 3), using the IUGS total alkali versus silica (TAS)
system of Le Maitre (2002) and the proposals of Le Bas
(1989) for strongly silica-undersaturated compositions.
Melanephelinites are distinguished from nephelinites on
1313
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
JULY 2005
Table 1: Geochemical data for a representative set of lavas from the Vitim Volcanic Field
Locality:
tuff pit
Antasey
Ekzar
Ekzar
Kolichikan
Sample:
93VBS 41
93VBS 215
93VBS 355
93VBS 365
93VBS 112
93VBS 268
Depth (m):
Surface
171
105
283
212
156
Rock type:
MNEPH
MNEPH
BAS
BAS
MNEPH
MNEPH
SiO2
TiO2
Al2O3
Fe2O3*
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
LOI
Mg-number
43.76
2.58
43.91
2.72
44.48
2.71
45.17
2.69
44.26
2.84
44.83
2.49
10.53
13.85
0.20
12.83
12.84
0.17
13.14
13.20
0.18
13.55
12.40
0.17
12.86
13.70
0.18
12.75
0.18
14.57
9.85
11.19
9.65
11.41
9.80
11.28
9.66
2.55
1.61
0.90
3.24
1.75
0.53
2.96
1.64
0.49
2.88
2.10
0.55
9.21
10.19
3.71
100.40
3.67
98.83
0.34
69.8
65.7
99.99
Gain
65.5
100.45
1.30
99.68
4.42
100.45
0.16
66.7
59.7
65.1
4389
484
395
489
556
60
56
58
50
54
Cr
922
323
347
325
Ni
Pb
Rb
5.74
78.8
567
4.90
29
5.07
50.0
258
3.36
36
4.38
45.9
241
2.30
23
5.38
51.3
239
3.47
28
Sc
15
24
29
16
Sr
1367
683
639
709
Ta
Th
U
V
4.19
6.40
1.27
154
1.04
215
2.85
3.31
0.67
226
3.18
3.75
1.04
218
272
4.87
52.6
121
2.76
44
15
814
3.16
3.44
2.14
484
49
353
5.52
57.6
226
4.76
29
21
891
3.38
4.78
1.22
177
20.6
185
22.5
24.0
25.0
Zn
107
77
77
91
89
88
Zr
205
213
175
232
203
215
Y
26.1
3.17
4.01
10.83
9.06
3.95
2.07
0.66
Co
Nb
13.63
1.97
0.77
Ba
Hf
Bortovoy
La
57.28
Ce
109.27
13.84
26.4
29.82
59.90
28.67
57.34
33.17
66.06
29.56
60.59
45.26
86.10
7.82
32.30
6.62
7.58
56.89
10.97
31.24
6.60
8.69
35.67
7.38
8.13
34.50
7.26
10.83
44.10
3.54
10.06
2.13
6.00
2.18
6.36
2.35
6.83
2.35
6.45
1.26
5.88
0.85
4.51
0.91
4.74
0.96
4.93
0.89
4.42
0.93
1.92
0.81
1.90
0.86
2.04
0.89
2.08
0.74
1.61
Yb
0.24
1.26
0.27
1.49
0.30
1.65
0.31
1.70
0.22
1.18
Lu
0.17
0.23
0.25
0.25
0.17
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
1314
8.67
2.80
7.78
1.09
5.40
0.93
2.13
0.31
1.69
0.25
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Locality:
Ekzar
Ekzar
Hoygot
Antasey
Burul’zay
Antasey
Sample:
93VBS 162
93VBS 356
93VBS 25
93VBS 205
93VBS 55
93VBS 283
Depth (m):
180
109
50
125
80
15
Rock type:
BAS
NEPH
NEPH
NEPH
AB
AB
SiO2
TiO2
Al2O3
Fe2O3*
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
LOI
Mg-number
43.55
3.01
44.68
3.17
44.20
2.68
43.82
2.82
46.90
2.15
49.02
2.44
12.24
14.13
13.90
13.98
13.55
12.43
13.11
13.54
13.17
12.09
16.15
11.53
0.19
11.99
0.16
7.42
0.16
10.08
0.18
9.43
0.16
10.26
0.15
5.07
9.90
3.36
8.47
4.91
8.95
4.51
9.43
4.68
9.67
3.29
9.05
4.15
1.16
0.69
2.77
0.99
2.44
0.83
2.18
1.05
1.61
0.46
1.88
0.53
100.22
2.10
100.45
99.83
100.23
0.67
99.76
99.97
0.16
65.1
Gain
53.9
Gain
64.1
60.5
Gain
65.1
49.2
Ba
422
583
552
490
415
Co
59
46
46
47
52
38
Cr
276
79
237
363
403
55
Hf
Nb
Ni
Pb
Rb
6.04
63.2
209
3.54
20
7.31
83.7
100
4.69
35
5.89
79.3
216
4.87
30
7.05
76.6
147
5.00
37
3.82
36.1
185
2.78
23
438
4.76
39.3
39
2.30
20
Sc
18
8
14
16
23
16
Sr
794
1113
1005
1078
621
695
Th
3.98
4.74
5.07
6.03
4.58
5.59
4.70
5.98
2.14
2.81
2.37
2.60
U
1.30
1.56
1.53
1.72
0.75
0.70
Ta
V
Y
204
24.3
171
26.3
169
23.2
156
27.4
165
21.1
199
24.5
Zn
91
110
83
110
74
78
Zr
258
321
267
309
157
206
La
36.84
73.08
59.67
Pr
9.53
115.60
14.48
Nd
39.40
8.10
58.77
11.37
2.61
7.16
3.58
9.86
1.02
5.12
1.30
6.10
0.87
1.95
0.96
1.98
Yb
0.28
1.48
0.25
1.24
Lu
0.21
0.16
Ce
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
46.20
87.62
47.10
92.75
21.63
43.29
25.51
51.68
10.90
44.24
12.10
49.32
8.63
2.79
9.81
3.08
5.74
24.40
5.47
29.70
6.72
7.78
1.03
8.35
1.15
1.83
5.34
2.28
6.64
5.01
0.82
5.76
0.97
0.78
4.13
0.93
4.84
1.82
0.24
2.13
0.29
0.75
1.80
0.88
2.08
1.32
0.18
1.55
0.22
0.27
1.48
0.31
1.74
0.22
0.26
1315
6.96
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
JULY 2005
Table 1: continued
Locality:
Burul’zay
Muliha
Hoygot
Ekzar
Ekzar
Sample:
93VBS 63
93VBS 43
93VBS 92
93VBS 369
93VBS 370
Kolichikan
93VBS 315
Depth (m):
100
101
6
317
319
128
Rock type:
AB
AB
AB
THOL
THOL
THOL
45.98
2.06
47.99
2.51
50.50
1.99
49.28
2.07
49.10
2.02
50.28
2.17
12.97
12.37
12.99
12.87
14.63
10.72
13.60
12.07
13.61
12.06
13.41
12.31
MnO
0.17
MgO
11.27
10.08
0.17
9.52
0.14
7.62
0.14
10.62
8.52
0.15
7.18
SiO2
TiO2
Al2O3
Fe2O3*
CaO
Na2O
K2O
2.79
1.54
8.74
8.05
0.15
10.33
8.69
2.73
1.54
3.81
1.62
2.79
1.16
2.48
1.14
9.23
3.19
1.09
P2O5
0.44
0.49
0.38
0.36
Total
99.56
2.94
99.46
LOI
99.67
1.09
100.50
2.15
0.35
100.04
3.87
99.38
1.84
Mg-number
66.7
61.9
65.3
66.0
56.2
Gain
61.0
Ba
393
420
338
285
305
Co
54
45
39
50
46
Cr
283
349
182
325
Hf
Nb
Ni
Pb
Rb
3.76
37.6
205
3.56
20
5.10
39.6
197
3.36
29
3.67
28.9
148
2.56
21
3.69
29.4
208
2.15
15
Sc
19
21
17
21
Sr
765
1477
607
486
Ta
Th
U
V
2.13
2.73
0.71
161
2.33
3.07
0.76
170
1.78
2.21
0.53
131
1.85
2.09
0.53
150
273
3.45
27.5
201
4.06
15
17
570
1.72
1.99
0.41
0.36
309
42
219
4.22
25.9
149
2.55
18
20
497
1.50
2.12
0.52
137
20.6
155
Y
20.8
25.1
19.5
21.1
Zn
89
72
70
80
82
87
Zr
137
205
145
151
140
166
La
Ce
23.37
45.92
26.60
54.25
24.6
19.02
38.04
19.47
39.90
18.69
38.30
18.68
38.56
Pr
5.97
7.23
5.13
Nd
25.68
5.58
31.60
7.01
22.52
5.39
5.31
23.32
5.28
5.12
22.41
5.09
23.80
5.85
1.88
5.53
2.33
6.92
1.90
5.53
1.81
5.39
1.76
5.20
2.01
6.22
0.79
4.06
1.00
5.12
0.76
3.98
0.79
4.19
0.75
4.02
0.91
4.84
0.73
1.76
0.91
2.10
0.70
1.64
0.76
1.87
0.73
1.79
0.88
2.09
Yb
0.25
1.39
0.30
1.65
0.24
1.34
0.27
1.53
0.27
1.48
0.31
1.67
Lu
0.21
0.24
0.19
0.23
0.22
0.26
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
1
5.27
Lavas that host mantle xenoliths.
*Total iron reported as Fe2O3.
MNEPH, melanephelinite; NEPH, nephelinite; BAS, basanite; AB, alkali basalt; THOL, tholeiitic basalt; n.d., not
determined. Fe2þ/(Fe2þ þ Fe3þ) ¼ 0.9; Mg-number ¼ Mg2þ/(Mg2þ þ Fe2þ) 100. Localities are given in Fig. 2b.
1316
3.69
6.91
5.06
7.80
1317
3.56
6.55
4.03
7.21
Hf/177Hf(i)
15
3.62
7.36
0.512802 08c
0.282975 09e
0.282974 09e
0.704878 11b
0.704868 11b
0.512819 08c
AB
100
93VBS 63
Burul’zay
4.29
7.55
0.512831 09c
0.282980 07e
0.282979 07e
0.704101 10a
0.704085 10a
0.512853 09c
BAS
283
93VBS 365
Ekzar
4.55
8.08
0.512842 13d
0.282995 08f
0.282994 08f
0.704036 13a
0.704028 13a
0.512867 13d
AB
101
93VBS 43
Muliha
4.58
7.27
0.512831 10c
0.282972 07e
0.282971 07e
0.703933 10a
0.703912 10a
0.512868 10c
MNEPH
212
93VBS 112
Kolichikan
4.30
7.02
4.84
8.05
1.98
5.85
0.512717 17d
0.282932 13f
0.282931 13f
4.42
8.03
4.08
7.75
0.512830 11d
0.282986 07f
0.282985 07f
0.704303 12a
0.704291 12a
0.512843 11d
OLTH
317
4.63
8.72
0.512847 07c
0.283013 07e
0.283012 07e
0.704123 15b
0.704111 15b
0.512870 07c
NEPH
50
93VBS 25
Hoygot
4.12
7.60
0.512832 09c
0.282982 06e
0.282981 06e
0.704791 09a
0.704781 09a
0.512845 09c
OLTH
319
93VBS 370
Ekzar
0.512833 15d
0.282993 05f
0.282993 05f
0.703931 15a
0.703919 15a
0.512859 15d
NEPH
109
93VBS 356
Ekzar
93VBS 369
Ekzar
0.512865 06c
0.282994 08e
0.282993 08e
0.704033 09b
0.704023 09b
0.512881 06c
BAS
180
93VBS 162
Ekzar
0.704687 15a
0.704673 15a
0.512736 17d
AB
6
93VBS 92
Hoygot
0.512844 15d
0.282965 07f
0.282964 07f
0.703958 14a
0.703945 14a
0.512853 15d
MNEPH
156
93VBS 268
Bortovoy
4.59
8.17
0.512853 19d
0.282998 08f
0.282997 08f
0.703984 13a
0.703970 13a
0.512870 19d
OLTH
128
93VBS 315
Kolichikan
n.d.
n.d.
4.51
n.d.
0.512835 06c
0.703955 19a
0.703941 19a
0.512864 06c
NEPH
125
93VBS 205
Antasey
(n), measured ratio normalized to the standard value (see Table 3); (i), initial ratio; n.d., not determined. For superscripts, af refer to Table 3. An age of 10 Ma and the
following decay constants were used for calculating epsilon values: RbSr: 1.42 1011; SmNd: 6.54 1012; LuHf: 1.876 1011. Values for the Chondritic
Uniform Reservoir (CHUR) were taken as: 143Nd/144Nd ¼ 0.512638; 147Sm/144Nd ¼ 0.196700; 176Hf/177Hf ¼ 0.282772; 176Lu/177Hf ¼ 0.033200.
eHfi
eNdi
176
0.512799 17d
0.282952 08f
0.282951 08f
Hf/177Hf(n)
0.512820 13c
0.282971 07e
0.282970 07e
Nd/144Nd(i)
176
Sr/86Sr(i)
87
Nd/144Nd(n)
0.704307 13a
0.704292 13a
0.512840 13c
Sr/86Sr(n)
87
143
0.704274 12a
0.704263 12a
0.512816 17d
AB
Rock type:
143
AB
80
Depth (m):
93VBS 283
93VBS 55
Antasey
Burul’zay
4.66
7.96
0.512853 08c
0.282992 08e
0.282991 08e
0.704014 19a
0.704029 19a
0.512872 08c
BAS
105
Sample:
171
93VBS 355
Ekzar
Locality:
eHfi
eNdi
Hf/177Hf(i)
176
0.512793 13c
0.282962 06e
0.282961 06e
Hf/177Hf(n)
0.512870 08c
0.282987 07e
0.282986 07e
Nd/144Nd(i)
176
Sr/86Sr(i)
87
Nd/144Nd(n)
0.704183 14a
0.704175 14a
0.512892 08c
Sr/86Sr(n)
87
143
0.704167 15a
0.704146 15a
0.512822 13c
MNEPH
Rock type:
143
MNEPH
Surface
Depth (m):
93VBS 215
93VBS 4
Sample:
Antasey
tuff pit
Locality:
Table 2: Sr–Nd–Hf isotope data for a representative set of lavas from the Vitim Volcanic Field
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
JULY 2005
Table 3: Standard reproducibility for Sr, Nd and Hf in each of the two analytical sessions
Element
Standard
Session 1
Session 2
Reproducibility
2SE
0.000011
0.000015
0.000006
Sr
NBS 987
Nd
J&M Nd
0.710249a
0.511097c
Hf
JMC 475
0.282157e
Reproducibility
2SE
n
18
0.710252b
0.511107d
0.000007
0.000009
11
13
0.282155f
0.000004
7
n
9
4
The superscripts af are used in Table 2 to show which samples were analysed in each analytical session, and the standard
reproducibilities that apply to these. All isotopic data are reported relative to the accepted standard values of 0.71024 for
NBS 987, 0.28216 for JMC 475 (Nowell et al., 1998) and 0.511110 for J&M Nd (G. M. Nowell, personal communication).
The value of 0.511110 for J&M Nd is equivalent to the accepted value for 143Nd/144Nd in the international La Jolla standard
(0.51186). Nd was analysed as part of a total REE cut, using the analytical methods of Nowell & Parrish (2001) and Nowell
et al. (2003). Sm-doped J&M standards were run to test the Sm correction; the reproducibility reported above includes both
Sm-doped and undoped standards. Sm corrections were applied on all relevant Nd isotopes: 144Nd, 148Nd and 150Nd. Mass
bias on 87Sr/86Sr and 176Hf/177Hf was corrected for using an exponential correction to the accepted 86Sr/88Sr and
179
Hf/177Hf ratios of 0.1194 and 0.7325, respectively. As there is a Sm interference on 144Nd, mass bias on the 143Nd/144Nd
ratio was corrected for using an exponential correction to a 146Nd/145Nd ratio of 2.071943, which is equivalent to the more
commonly used 146Nd/144Nd ratio of 0.7219 (Nowell & Parrish, 2001).
Fig. 3. Total alkalis (Na2O þ K2O) vs SiO2 (Le Maitre, 2002) for all analysed Vitim lavas.
the basis of their normative nepheline contents:
melanephelinites contain less than 20% normative
nepheline, whereas nephelinites contain >20%.
Figure 4 shows variations in SiO2 within the lavas from
the Ekzar 3315, Antasey 4404 and Muliha 4633 drillcores. It is clear that the sub-alkaline (tholeiitic) and alkali
basalts and more strongly alkalic (basanite–nephelinite)
magmas were erupted without any temporal pattern;
highly alkalic lavas both overlie and underlie the
tholeiitic basalts. The only consistent feature is that the
youngest lavas shown in Figs 2 and 4 are all strongly
alkalic.
PETROGRAPHY
Table 4 summarizes the average modal proportions of
the various phenocryst phases in the Vitim lavas.
Melanephelinites, nephelinites, basanites and alkali
basalts all have porphyritic textures with olivine, clinopyroxene plagioclase phenocrysts set in a fine-grained
groundmass consisting of olivine, clinopyroxene, opaque
oxides and plagioclase laths. Olivine phenocrysts are
equant, often euhedral, and are most abundant in the
nephelinites and melanephelinites. They vary from 02 to
30 mm in size, are commonly altered around the rims to
1318
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Ekzar 3315
Antasey 4404
100
Muliha 4633
0
0
20
Depth below surface (m)
40
20
150
60
80
40
200
100
120
60
250
140
80
300
160
180
200
100
350
40
45
50
40
55
Granite basement
below sediments,
at ~ 395 m
SiO2 (wt. %)
45
50
55
Granite basement
below sediments/lavas,
at ~ 200 m
nephelinites
basanites
alkali basalts
tholeiitic basalts
40
SiO2 (wt. %)
45
50
55
Granite basement
below sediments,
at ~ 208 m
Volcanics
Lacustrine sediments
Fig. 4. Variation of SiO2 (wt %) with depth (m) in lavas from three drill-holes from the VVF.
Table 4: Visual estimates of average modal proportions of phenocryst phases found in Vitim lavas
Rock type
Olivine
Clinopyroxene
Plagioclase
FeTi oxides
Nephelinite
210%
2%
25%
2%
Melanephelinite
515%
5%
1015%
2%
Basanite
515%
510%
1020%
25%
Alkali basalt
210%
515%
1025%
2%
Tholeiitic basalt
510%
520%
1530%
1%
brown iddingsite, and often contain inclusions of Crspinel. Some basanites and alkali basalts contain kinkbanded olivine crystals, which may be xenocrysts (see
below). Clinopyroxene phenocrysts are often pinkish in
colour because of Ti enrichment. Large plagioclase
phenocrysts (up to 2 mm in length) are found in the
tholeiitic basalts and some alkali basalts, but are rare or
absent in the other magma types. The tholeiitic basalts
are dominated by laths of plagioclase feldspar, with generally fewer olivine phenocrysts than the other rock types.
Cr-spinel inclusions in olivine are rare in the tholeiitic
basalts.
1319
100.06
85.2
0.18
0.18
100.26
86.0
100.10
74.2
0.29
0.27
100.09
79.2
100.54
77.6
0.29
0.25
100.06
80.8
100.56
84.4
0.21
0.24
100.18
80.0
99.40
81.3
Total
Mg-number
CaO
0.36
0.28
99.97
80.9
0.23
100.50
78.6
MgO
100.03
73.3
—
—
0.18
45.23
0.16
46.10
0.34
37.82
0.26
41.14
40.35
42.35
0.24
42.36
0.28
40.90
NiO
0.19
0.27
MnO
0.41
36.99
0.23
0.26
0.25
41.78
0.19
45.17
0.23
42.18
0.11
0.3
17.92
0.28
14.89
0.22
18.68
0.15
24.08
0.16
19.85
0.24
FeO
—
—
14.02
0.27
13.33
0.24
23.40
0.14
19.30
0.15
—
—
—
—
—
—
0.01
20.73
—
—
—
—
—
17.32
—
—
—
—
Cr2O3
—
—
17.78
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
TiO2
Al2O3
39.83
40.17
38.09
38.96
38.76
39.19
39.86
39.07
38.44
39.05
38.99
SiO2
38.01
93VBS 207
93VBS 152
93VBS 54
93VBS 29
93VBS 12
Sample:
Table 5: Electron microprobe analyses of olivine phenocrysts in Vitim lavas
93VBS 214
93VBS 223
93VBS 283
93VBS 288
93VBS 310
93VBS 367
93VBS 373
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
JULY 2005
MINERAL CHEMISTRY
A comprehensive description of the mineral chemistry of
the Vitim lavas is beyond the aim of this paper. However,
below, we give a brief summary of the chemical compositions of olivine, clinopyroxene and plagioclase feldspar
phenocrysts present in the Cenozoic lavas, and describe
how olivine compositions can be used to understand the
processes operating in crustal magma chambers beneath
the VVF.
We have studied the composition of approximately 600
olivine crystals from the Vitim lavas (Electronic Appendix
C, available at http://www.petrology.oupjournals.org);
Table 5), in order to understand whether they crystallized
in equilibrium with the magmas, or were xenocrysts from
the underlying lithospheric mantle. Olivine phenocrysts
in the Vitim lavas have forsterite contents in the range
Fo67–Fo87 (Fig. 5). A trend towards a higher frequency of
olivine phenocrysts with increasing forsterite content is
observed. There are no phenocrysts with Fo > 87%;
olivines with Fo87–Fo91 are invariably xenocrysts (identified by their anhedral shape, embayed margins, kinkbanding, CaO contents below 01 wt %, and high Fo
content). There is no systematic variation in Fo content
with rock type.
There is a general trend of increasing NiO content of
olivine crystal cores with Fo content. The highest Fo
contents (>Fo87) are observed in olivines from lherzolite
xenoliths from the ‘tuff pit’ locality of Ionov et al. (1993).
These also have the highest NiO contents (>04 wt %).
All other olivines are believed to be phenocrysts (consistent with CaO>018 wt %; Brey et al., 1990); those with
the highest NiO contents are likely to have crystallized
from the most primitive magmas. The compositions
of Vitim olivines fall in the range expected for oceanic
picrites [compare with fig. 6 of Gibson (2002)], but the
Fo contents extend to lower values because the lavas
from the VVF are more evolved (the majority have
MgO < 12 wt %, and are therefore not classified as
picrites).
Analyses of representative clinopyroxene phenocrysts
are given in Table 6 (the full dataset is given in Electronic
Appendix D, available at http://www.petrology.
oupjournals.org). Using the classification scheme of
Morimoto (1988), all of the pyroxenes analysed (from
both lavas and xenoliths from the VVF) are quadrilateral.
Their compositions are restricted on a Wo–En–Fs
triangular plot (not shown), with the clinopyroxene
phenocrysts generally falling within or very close to the
diopside field. Clinopyroxene phenocrysts have Mgnumber [Mg-number ¼ Mg/(Mg þ Fe)] in the range
701–834, and are estimated to have crystallized from
liquids with Mg-number ¼ 350–536, using a partition
coefficient for Fe–Mg partitioning between augite and
melt of 023 (Grove & Bryan, 1983).
1320
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
90
Number of analyses
80
70
phenocrysts
xenocrysts
60
50
40
30
20
10
0
66 68 70 72 74 76 78 80 82 84 86 88 90
% Forsterite
Fig. 5. Forsterite contents of olivines (phenocrysts and xenocrysts) from
lavas from the VVF.
Analyses of representative plagioclase feldspar phenocrysts are given in Table 6 (the full dataset is given in Electronic Appendix E, available at http://www.petrology.
oupjournals.org). Plagioclase feldspars typically have core
compositions that fall in the labradorite field, in the range
An60–An70 [An ¼ atomic Ca/(Ca þ Na) 100].
WHOLE-ROCK CHEMISTRY
Whole-rock geochemical analyses for representative samples from the VVF are given in Table 1; the full dataset
can be found in the study by Garner (2002) and Electronic Appendix A (available at http://www.petrology.
oupjournals.org).
Weathering and alteration
The majority of samples from the VVF are petrographically fresh, and contain few alteration minerals, e.g.
iddingsite. Loss on ignition (LOI) values range from
014 to 626, with the majority of samples having LOI
values of <3. Some were found to have gained mass
during ignition, as a result of oxidation of FeO to Fe2O3
(see Table 1). This occurs if there are large amounts of
FeO and small amounts of OH present in a sample, such
that the increase as a result of oxidation is greater than
the OH loss. Rb and K are particularly mobile during
alteration. However, there does not appear to be any
direct correlation between LOI and Rb and K contents
in the Vitim lavas, suggesting that alteration processes
have not significantly affected their chemistry.
Major element variation
The Cenozoic lavas of the VVF have SiO2 contents in the
range of 43–53 wt %, and form two series on a total
alkali versus silica diagram: a strongly alkaline series
(melanephelinites, nephelinites and basanites) and a
mildly/sub-alkaline series (alkali and tholeiitic basalts;
Fig. 3). SiO2 contents increase from nephelinites
(43–49 wt %) through to tholeiitic basalts (48–53 wt %;
Fig. 6). Conversely, MgO contents are lowest in the
tholeiitic basalts (33–106 wt %) and highest in the basanites (75–120 wt %). The nephelinites have a narrower
range of MgO contents than the other magma types
(65–101 wt %). Sample 93VBS 4 (a melanephelinite)
has anomalously high MgO (MgO ¼ 146 wt %) and
contains xenocrystic olivine.
CaO and Fe2O3* (total Fe) contents of the lavas range
from 77 to 116 and 97 to 147 wt %, respectively
(Fig. 6). Al2O3 abundances are highest in the tholeiitic
basalts (134–163 wt %), and lowest in the basanites,
nephelinites and melanephelinites (122–153 wt %).
TiO2 contents range from 18 to 33 wt %, with the
highest values in the basanites. Nephelinites have the
highest contents of Na2O (37–54 wt %), K2O
(17–30 wt %) and P2O5 (05–11 wt %) of all the rock
types. The melanephelinites exhibit a small range in all
the oxides, and contents are intermediate between those
of the nephelinites and basanites.
Despite the large number of samples analysed, the
Vitim lavas do not always show clear correlations on
Harker variation diagrams (Fig. 6), and there are no
distinct inflections in the trends of the data suggesting
the onset of crystallization of different phases. To account
for the wide range of major element compositions in the
lavas (Fig. 6), it is likely that there was a spectrum of
parental magma compositions. The systematic decrease
in Na2O, P2O5 and K2O (Fig. 6b, d and h) from nephelinites through to tholeiitic basalts for a range of MgO
values cannot be explained by fractional crystallization
and reflects variations in partial melting processes in the
mantle source (see below).
Trace-element variation
Concentrations of compatible trace elements, such as Ni
and Cr, are highest in the basanites. These range from
106 to 279 and 82 to 394 ppm, respectively. The lowest
abundances are found in the alkali basalts, in which Ni
varies from 39 to 238 ppm and Cr from 55 to 403 ppm.
Although the overall ranges for Ni and Cr are large (Ni ¼
39–279 ppm, and Cr ¼ 59–542 ppm), the majority of
samples have concentrations of Ni of <100 ppm, and
between 100 and 400 ppm Cr. They are therefore not
representative of primitive magma compositions. Both Ni
and Cr contents are positively correlated with MgO for
all of the groups, reflecting olivine and spinel crystallization, respectively (Fig. 7).
Abundances of incompatible trace elements in the
Vitim lavas generally increase from the tholeiitic basalts to
the nephelinites, e.g. Rb (9–52 ppm), Th (141–659 ppm),
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Table 6: Electron microprobe analyses of representative clinopyroxene and plagioclase feldspar phenocrysts in Vitim lavas
Mineral:
Clinopyroxene Clinopyroxene Clinopyroxene Clinopyroxene Clinopyroxene Feldspar
Feldspar
Feldspar
Feldspar
Feldspar
Sample:
93VBS 60
93VBS 207
93VBS 283
93VBS 284
93VBS 287
93VBS 112
93VBS 283
93VBS 284
93VBS 286
93VBS 45
51.87
1.21
48.57
2.07
48.74
1.76
49.35
1.70
51.20
1.27
—
—
1.70
0.21
4.61
0.25
5.17
0.38
4.53
0.31
2.33
0.13
—
—
7.01
0.18
6.98
0.14
7.03
0.14
7.46
0.13
8.15
0.18
CaO
15.38
22.13
13.58
23.17
14.16
22.23
14.11
22.14
15.21
20.94
0.14
12.57
0.20
13.41
0.24
13.80
Na2O
0.37
0.52
0.45
0.44
0.39
4.40
0.17
3.76
0.24
3.81
0.26
99.96
An 75.9
100.19
79.8
100.08
80.0
SiO2
TiO2
Al2O3
Cr2O3
FeO
MnO
MgO
K2O
—
—
100.06
Mg-number 79.6
Total
—
—
99.88
77.6
—
—
—
100.04
78.2
100.16
77.1
52.58
29.64
—
—
76.9
Nb (184–837 ppm), Ta (101–507 ppm), Sr (385–
1248 ppm), Zr (118–337 ppm), Hf (302–738 ppm). Pb
contents are more variable, but there is a general increase
observed from the tholeiitic basalts (118–406 ppm) to
the nephelinites (376–500 ppm). Concentrations of Y
are more constant, ranging from 202–319 ppm in the
tholeiitic basalts to 222–274 ppm in the nephelinites.
MgO contents are plotted against selected incompatible trace elements in Fig. 8 to show the effects of
fractional crystallization on the parental magmas. There
is no distinct trend defined by the Vitim lavas, confirming
that their evolution is more complex than simple fractional crystallization of a single parent magma. The wide
range in abundance of incompatible trace elements for a
given MgO content (i.e. between different groups) may,
however, be explained by variation in the degree of
partial melting or mantle source heterogeneity (Fig. 8;
see below).
On normalized multi-element plots (Fig. 9), all the lavas
from the VVF exhibit smoothly curved concave-upward
profiles that peak at Nb and Ta. The most incompatible trace element enriched lavas (relative to chondrites) are the nephelinites and melanephelinites, which
also have the lowest normalized abundances of the more
compatible trace elements, e.g. Yb. Concentrations of
strongly incompatible trace elements, such as Ba, are
higher in the nephelinites, melanephelinites and
basanites (273–594 ppm) than the alkali and tholeiitic
basalts (201–571 ppm).
Slight K depletions are observed in Fig. 9 for some of
the nephelinites and basanites ([Nb/K]n 12), whereas
the alkali basalts and tholeiitic basalts usually have either
no relative depletion or a slight relative enrichment in K
(e.g. [Nb/K]n ¼ 08 for 93VBS 314). In several samples,
30.59
—
—
0.46
—
—
99.78
51.40
—
—
50.59
—
—
30.83
—
—
0.59
—
—
0.55
—
—
50.32
51.74
—
—
31.01
—
30.18
—
—
0.50
—
0.57
—
—
0.23
—
0.17
13.89
3.78
13.15
3.92
0.25
0.23
99.99
80.2
99.96
78.8
there are slight relative enrichments in P (e.g. 93VBS 356)
and Ti (e.g. 93VBS 223). These are particularly marked
for P in the nephelinites, melanephelinites, basanites and
alkali basalts, and Ti in the alkali and tholeiitic basalts.
Chondrite-normalized rare earth element (REE) patterns form smoothly curved trends from La to Lu, with no
significant relative depletions or enrichments of individual elements (Fig. 10). Eu anomalies are absent, indicating that there has not been any significant plagioclase
fractionation. Light REE abundances are relatively
restricted in their overall range, increasing in the different
rock groups in the following order: tholeiitic basalts, alkali
basalts, basanites, melanephelinites, nephelinites.
The gradual flattening-out of the slope of the REE
pattern from nephelinites through to tholeiitic basalts is
clearly visible in Fig. 10. The highest light REE to heavy
REE ratios (LREE/HREE) are observed in the nephelinites ([La/Lu]n ¼ 203–380), and two samples have particularly enriched LREE abundances (93VBS 356 and
281). Notably high [La/Lu]n ratios (>360) are observed
in these nephelinitic samples. These lava flows can be
traced across the volcanic field, and usually lie at or near
the top of the lava pile. In general, the gradient of the
REE patterns is broadly correlated with SiO2 content
(lavas with lower SiO2 contents tend to have steeper
REE slopes, and hence higher [La/Lu]n and [La/Yb]n
ratios).
Radiogenic isotopes
Sr and Nd isotope systematics
Lavas from the VVF exhibit a small range of initial
Sr and Nd isotopic ratios (87Sr/86Sri ¼ 07039–07049
and 144Nd/143Ndi ¼ 051272–051287; Fig. 11).
1322
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
55
12
10
50
CaO
SiO2
8
45
2
melanephelinites
(a)
basanites
35
6
4
2
16
ol
14
12
ol
0
cpx
18
Al2O3
tholeiitic basalts
Na2O
20
par
tial
me
lting
nephelinites
(e)
0
alkali basalts
8
ol
cpx
4
ol
40
6
(b)
(f)
10
5
4
TiO2
Fe2O3
14
12
2
ol
10
plag
3
ol
1
(c)
8
(g)
0
5
K2 O
P2O5
1
par
tial
me
lting
par
tial
me
lting
6
1.5
4
plag
3
2
0.5
0
0
2
4
6
8
10
1
(d)
ol
12
14
0
MgO
(h)
ol
0
2
4
6
8
10
12
14
MgO
Fig. 6. Variation in major element oxides (wt %) vs MgO (wt %). Vectors illustrate the effects of fractional crystallization on magma composition;
the phase removed at each stage is labelled (ol, olivine; cpx, clinopyroxene; plag, plagioclase feldspar). The effects of increasing degree of partial
melting are indicated by arrows, where relevant.
There does not appear to be any systematic correlation of
isotopic composition with rock type or location within the
volcanic field; for example, 87Sr/86Sri ranges from
07039 to 07041 for nephelinites, 07040 to 07041 for
basanites, 07039 to 07042 for melanephelinites and
07040 to 07049 for alkali basalts, compared with
07040–07048 for tholeiitic basalts. 143Nd/144Ndi ranges
are as follows: 051283–051285 for nephelinites,
051279–051287 for melanephelinites, 051283–
051287 for basanites, 051272–051284 for alkali basalts
and 051283–051285 for tholeiitic basalts. All of the lavas
have more enriched Sr and Nd isotopic compositions
than the peridotite mantle xenoliths from the VVF
(Ionov et al., 1995).
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300
melanephelinites
nephelinites
basanites
Ni (ppm)
250
VOLUME 46
NUMBER 7
JULY 2005
ol
alkali basalts
alkali basalts
tholeiitic basalts
200
150
tholeiitic basalts
100
ol +
cpx
basanites
50
(a)
nephelinites
0
Cr (ppm)
500
increasing pressure
of crystallization
basanites
400
300
tholeiitic basalts
200
+
ol
100
sp
alkali basalts
(b)
nephelinites
0
0
2
8
6
MgO (wt. %)
4
10
12
14
Fig. 7. Concentrations of (a) Ni (ppm) and (b) Cr (ppm), plotted against MgO (wt %) for lavas from the VVF. Vectors illustrate the effects of
fractional crystallization on magma composition (phases removed are labelled as ol, olivine; cpx, clinopyroxene; sp, Cr-rich spinel).
Nd and Hf isotope systematics
On a diagram of eNd versus eHf (Fig. 12), the Vitim lavas
define a positive correlation and plot within the field for
OIB and close to Bulk Earth. The highest eHf values are
observed in the nephelinites (87) and the lowest in the
alkali basalts (59). The ranges in eHf overlap between the
different groups, and are as follows: 80–87 for the nephelinites; 69–78 for melanephelinites; 76–81 for basanites; 59–81 for alkali basalts; 76–82 for tholeiitic
basalts. As in Fig. 11, sample 93VBS 92 (an alkali basalt)
plots away from the other samples, at lower eNd and
slightly lower eHf values. All of the lavas have lower eNd
and eHf values than peridotite xenoliths from the VVF
(D. A. Ionov, personal communication, 2003).
CRUSTAL PROCESSES
Before we can model the mantle melting processes that
gave rise to the parental magmas of the Vitim lavas, it is
important to constrain their melt source characteristics.
In order to do this, it is first necessary to establish the
extent to which magma compositions have been affected
by crustal processes such as fractional crystallization and
crustal assimilation.
Parental magma composition
The forsterite contents of olivines from the Vitim lavas
(see Electronic Appendix C for data) were used to estimate the Mg-number and the MgO content of the magma
with which they were in equilibrium. In order to calculate
the Mg-number of the melt, it is first necessary to estimate
the Fe2O3/FeO ratio. The oxygen fugacity ( f O2) was
calculated using Mg2þ and Fe2þ exchange between coexisting olivine and spinel (Sack & Ghiorso, 1991). We estimated Dlog f O2 values for some Vitim lavas from electron
microprobe analyses of olivine phenocrysts and the spinels contained within them (see Electronic Appendix F,
available at http://www.petrology.oupjournals.org, for
1324
JOHNSON et al.
60
7
50
6
5
2
10
1
0
0
100
alkali basalts
tholeiitic basalts
80
La
ol
80
melanephelinites
nephelinites
basanites
100
ol
60
60
40
40
20
0
1600
1400
1200
1000
800
600
400
200
0
0
350
ing
de
ar
gr
tia
ee
lm
elt
ing
20
250
Zr
inc
re
as
300
of
p
Sr
3
20
120
ol
4
in
of crea
pa si
rti n g
al d
m eg
e lt r e
in g e
Th
ol
30
in
d cre
pa egr asi
rti ee n g
al o
m f
el
tin
g
Rb
40
Nb
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
200
150
100
ol
ol
50
0
0
5
10
0
15
5
10
15
MgO
MgO
Fig. 8. Selected incompatible element abundances (in ppm) plotted against wt % MgO for lavas from the VVF. Vectors illustrate the effects of
fractional crystallization of olivine (ol), and increasing degree of partial melting.
data). Dlogf O2 ranges from 38 to 84 units below the
fayalite–magnetite–quartz (FMQ) buffer. Using the
empirical expression of Kilinc et al. (1983) and the revised
constants of Holloway et al. (1992), we estimated that the
Fe2O3/FeO ratio of these magmas is close to 008.
Assuming that the olivine–spinel pair with the highest
calculated Df O2 has undergone the least subsolidus
re-equilibration, and therefore gives a minimum value
for Fe2O3/FeO (Gibson, 2002), the ratio used in the
following calculations is 01.
Equilibrium values for olivine phenocrysts for a range
of whole-rock Mg-numbers are shown in Fig. 13. Many
samples contain olivines whose Fo contents are too low to
be in equilibrium with the whole rock (e.g. 93VBS 204
and 288). This suggests that they have been incorporated
into a magma with lower Mg-number [Mg-number ¼
Mg/(Mg þ Fe)], thus increasing the whole-rock Mgnumber, or that the most forsteritic olivines in these
samples were not analysed. Figure 13 shows that the
most mafic VVF lava containing equilibrium olivine is
93VBS 367 (an alkali basalt). Olivine phenocrysts in this
lava have moderate forsterite contents (up to Fo86) and
we have calculated that these were in equilibrium with
ol-liq
= 03,
magma with a Mg-number of 649 (using KD Fe-Mg
from Roeder & Emslie, 1970). Nephelinite 93VBS 281
contains slightly less forsteritic olivine phenocrysts (Fo85)
that are in equilibrium with the whole-rock Mg-number
(630). Using the variation in whole-rock Mg-number and
MgO contents for the VVF lavas (Fig. 14), parental
magma compositions of 108 and 99 wt % MgO were
estimated for the alkali basalt and nephelinite lavas,
respectively. It is unlikely that a magma with <11 wt %
MgO is representative of a primary magma, as worldwide
primary magmas from intra-plate settings have comparatively higher MgO contents, e.g. MgO ¼ 175 wt %
for Mauna Loa, Hawaii (Garcia et al., 1995); MgO ¼
20–21 wt % for West Greenland (Larsen & Pedersen,
2000).
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1000
Nephelinites &
Melanephelinites
Basanites
rock/chondrite (except K, Rb and P)
100
10
93VBS 399
93VBS 401
93VBS 356
93VBS 205
93VBS 95 (melanephelinite)
93VBS 403
93VBS 94
93VBS 211
93VBS 37
1000
Alkali Basalts
Tholeiitic Basalts
100
10
93VBS 283
93VBS 266
93VBS 12
93VBS 223
93VBS 315
93VBS 316
93VBS 370
93VBS 288
1
PbBaRbTh U K NbTa LaCeSrNd PSmZr Hf Ti Tb Y TmYbLu Pb BaRbTh U K NbTa LaCeSrNd PSmZr Hf Ti Tb Y TmYbLu
elements
elements
Fig. 9. Normalized multi-element plots for several samples from each of the rock groups: nephelinites, melanephelinites, basanites, alkali basalts,
tholeiitic basalts. K, Rb and P are normalized to primitive mantle values; all other elements are normalized to chondritic values. Normalizing
factors are from Thompson (1982), except for U (McDonough & Sun, 1995) and Lu (Wood, 1979).
Ni and Cr partition coefficients can also be used to
suggest whether parental magmas are primary. Primary
magmas are generally expected to have Ni > 400–
500 ppm and Cr > 1000 ppm, together with Mg-number
>70 (Wilson, 1989). Hart & Davis (1978) showed that
the Ni partition coefficient is dependent on the MgO
content of the melt (their fig. 5). These are related by
the equation
ol -liq
KD Ni ¼ ½12413=wt % MgO 0897:
Using this equation, the partition coefficient most
appropriate for a parental melt containing 108 wt %
olliq
MgO is 106. A KD Ni value of 106 gives an estimated
Ni content of 260 ppm for the melt in equilibrium
with an olivine of Fo86 (the most forsteritic olivine in
93VBS 367). The calculated Cr content of the melt is
800 ppm.
Polybaric fractional crystallization
The pressures and depths of fractional crystallization may
be estimated using CIPW normative compositions. We
have plotted the CIPW norms of the Vitim lavas on a
Ne–Ol–Di–Hy–Qz projection and compared them with
cotectics for basaltic liquids in equilibrium with olivine,
plagioclase and clinopyroxene at different depths within
the crust (e.g. Thompson, 1983; Thompson et al., 2001;
Fig. 15). The large majority of lavas from the VVF are
strongly silica-undersaturated (nepheline normative), but
a few are hypersthene normative and quartz normative.
All of the samples lie between the 1 atm and 9 kbar
cotectics, indicating that the Vitim magmas fractionated
over a wide range of pressures within the crust (Fig. 15).
The nephelinites appear to have undergone fractional
crystallization at the lowest pressures (1 atm), whereas
some of the tholeiitic basalts plot between the 1 atm and
9 kbar cotectics and have undergone fractionation
1326
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
rock/chondrite
1000
Nephelinites &
Melanephelinites
Basanites
100
10
93VBS 401
93VBS 281
93VBS 399
93VBS 356
93VBS 204
93VBS 95
93VBS 1
93VBS 365
93VBS 268
(melanephelinite)
93VBS 212
93VBS 39
93VBS 84
1000
rock/chondrite
Alkali Basalts
Tholeiitic Basalts
100
10
93VBS 21
93VBS 287
93VBS 58
93VBS 62
93VBS 267
93VBS 92
93VBS 370
93VBS 288
93VBS 113
93VBS 45
93VBS 314
1
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
REEs
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
REEs
Fig. 10. Chondrite-normalized REE abundances in nephelinites, melanephelinites, basanites, alkali basalts and tholeiitic basalts from the VVF.
Several samples have been plotted to illustrate variations in concentrations within individual groups. Normalizing factors are from McDonough &
Sun (1995).
0.5135
0.5133
143
Nd/144Ndi
0.5131
DMM
alkali basalts
basanites
nephelinites
Bartoy mantle
xenoliths
Vitim mantle
xenoliths
0.5129
93VBS 370
HIMU
tholeiitic basalts
melanephelinites
upper crust
93VBS 63
93VBS 92
0.5127
EMII
0.5125
EMI
0.5123
Tariat crustal
xenoliths
0.5121
0.7010
0.7020
0.7030
0.7040
0.7050
87
0.7060
0.7070
0.7080
0.7090
0.7100
86
Sr/ Sri
Fig. 11. 87Sr/86Sri vs 143Nd/144Ndi for a subset of lavas from the VVF. The samples are grouped according to rock type. Also shown are the fields
for Tariat crustal xenoliths (Barry et al., 2003), Vitim and Bartoy peridotite mantle xenoliths (Ionov et al., 1995), and the present-day mantle
components DMM, HIMU, EMI and EMII (Hart et al., 1992). Sample 93VBS 92 is an alkali basalt.
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25
Vitim peridotite xenoliths
20
15
10
OIB
εHf 5
MORB
93VBS 92
Bulk Earth
0
melanephelinites
-5
basanites
alkali basalts
-10
Contamination
by upper crust
nephelinites
tholeiitic basalts
2SE
-15
-15
-10
-5
0
5
εNd
10
15
20
Fig. 12. eNd and eHf values for a subset of lavas from the VVF. The samples are grouped according to rock type. The field for Vitim peridotite
mantle xenoliths is from D.A. Ionov (personal communication, 2003). OIB and MORB fields are from Nowell et al. (1998). Sample 93VBS 92 is
an alkali basalt.
95
90
KD = 0.23
KD = 0.25
KD = 0.27
KD = 0.21
KD = 0.30
KD = 0.33
152
397
215
55
281
223
54
204
310
70
373
112
288
75
367
80
207
Fo content %
85
22
65
55
57
59
61
63
65
67
69
whole-rock Mg-number
Fig. 13. Olivine phenocryst compositions compared with predicted equilibrium forsterite (Fo) contents (mol % Fo) as a function of whole-rock
Mg-number. Sample numbers (all 93VBS . . . ) are indicated adjacent to the phenocryst compositions. The grey shaded field indicates the region of
olivine Fo contents in equilibrium with the whole-rock, according to the partition coefficient data of Roeder & Emslie (1970). Dashed lines
indicate the expected equilibrium Fo contents for different values of the partition coefficient.
1328
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
80
Whole-rock Mg-number
70
60
50
40
30
Parental magma
for nephelinites
20
10
Parental magma
for alkali basalts
y = –0·1363x2 + 4·9414x + 27·475
2
R = 0·9095
0
0
1
2
3
4
5
6
7
8
9
10 11 12 13 14 15 16
Whole-rock MgO (wt.%)
Fig. 14. Whole-rock MgO (wt %) vs whole-rock Mg-number for lavas from the VVF. Data were fitted with a polynomial curve with the equation
shown. Short-dashed and long-dashed lines indicate the estimated parental magma compositions (Mg-numbers of 649 and 630, corresponding to
108 (alkali basalts) and 99 (nephelinites) wt % MgO, respectively; see text); the shaded fields show the range of whole-rock analyses that could
represent each of these, taking into account the slight scatter in the Vitim data.
Diopside
Quartz
Nepheline
1a
9 ( 1.5) kb
Olivine
Wt. %
tm
nephelinites
melanephelinites
basanites
alkali basalts
tholeiitic basalts
Hypersthene
Fig. 15. CIPW normative compositions for the lavas from the VVF. Norms were calculated assuming Fe2O3/FeO ¼ 01 (see text). The 1 atm and
9 (15) kbar cotectics are from Thompson et al. (2001). Arrows point in the direction of decreasing temperature.
throughout the crust. A recent seismic study by Suvorov
et al. (2002) suggests that the base of the Moho beneath
the VVF is 35 km, and fractional crystallization of
tholeiitic magmas at 9 kbar during the formation of the
VVF may have caused underplating of the crust.
Crustal contamination
In the subset of samples analysed for their Sr–Nd–Hf
isotopic ratios, three samples (93VBS 63, 92 and 370)
have significantly higher 87Sr/86Sri ratios than the
rest of the group (Fig. 11). This displacement to
higher 87Sr/86Sri ratios may be caused by hydrothermal
alteration and/or crustal contamination, as both the
upper and lower crusts are known to have 87Sr/86Sri >
07045 (Taylor & McClennan, 1985). These lavas have
low U/Pb ratios (<02), which may also indicate contamination (Thompson et al., 2001). The majority of samples,
however, have similar U/Pb ratios to oceanic basalts
(02–04), eNd values of 3–5 and 87Sr/86Sri < 07045
(Fig. 11), which suggest that crustal contamination has
not played a significant role in their petrogenesis. This is
confirmed by combined variations in Hf and Nd isotopic
ratios. Most continental upper crust has present-day eNd
< 0 and eHf < 0, and plots in the lower left quadrant of
the eNd vs eHf diagram in Fig. 12 (Vervoort et al., 1999).
1329
JOURNAL OF PETROLOGY
VOLUME 46
Figure 12 shows that, with one exception, the Vitim data
are tightly clustered and show no sign of a trend towards
crustal compositions. We therefore believe that the parental magmas of the Vitim lavas have not undergone
significant interaction with the continental crust and
that their compositions may be used to assess the nature
of the underlying mantle.
NUMBER 7
X/Al 2O3 (where X = Ba or Sr)
100
MANTLE SOURCE
CHARACTERISTICS
Evidence for mantle source heterogeneity
JULY 2005
nephelinites
90
melanephelinites
80
basanites
70
tholeiitic basalts
Sr
alkali basalts
60
50
Ba
40
30
20
Sr
10
Ba
phlogopite
in source
0
0
The variations in abundance of major, trace and REE in
the Vitim lavas, and their Sr–Nd–Hf isotope systematics,
indicate that the mantle source regions beneath the VVF
may be heterogeneous in composition. Tangible evidence
for lithospheric mantle heterogeneity comes from the
wide variety of xenoliths that are found in the VVF,
including garnet- and spinel-bearing peridotites, as well
as pyroxenites (Ionov et al., 1993; Ashchepkov et al., 1994;
Litasov et al., 2000; Ionov, 2004). These have been collected from a limited number of places, such as the
165 Ma ‘tuff pit’ locality (Esin et al., 1995) (Fig. 2). The
mineralogy of the mantle xenoliths includes olivine,
clinopyroxene, orthopyroxene, garnet, spinel, amphibole,
phlogopite and ilmenite (Ionov et al., 1993; Litasov et al.,
2000).
Mineralogy of the mantle source
We can potentially identify mineral phases that might be
residual in the mantle source region of the Vitim magmas
by examining the behaviour of incompatible trace elements during melting, using published partition coefficient data for accessory minerals (e.g. McKenzie &
O’Nions, 1991; LaTourrette et al., 1995; Green et al.,
2000). Ratios between middle REE (MREE) and HREE
may be used to assess the presence or absence of garnet
during mantle melting. This is because the garnet–melt
partition coefficients are higher for the HREE than the
MREE (McKenzie & O’Nions, 1991). Chondritenormalized Gd/Lu ratios are high in the Vitim samples
(generally between 25 and 45) and therefore indicate
that their parental melts were generated within the garnet
stability field. Most nephelinitic samples have [Gd/Lu]n
ratios in excess of 3, which suggests that their parental
melts were formed at either: (1) greater pressures or (2)
smaller degrees of partial melting than the other Vitim
magmas. Sample 93VBS 4 has [Gd/Lu]n ¼ 7, which is
much higher than in all the other lavas. This was collected from the ‘tuff pit’ locality (Fig. 2), and contains
numerous xenocrysts derived from disaggregated crustal
and mantle xenoliths.
1
2
3
4
5
6
7
8
Th (ppm)
increasing degree of melting
Fig. 16. Variation of Ba/Al2O3 and Sr/Al2O3 (ppm/wt %) vs Th in
the Vitim lavas. Thorium (Th) behaves as an almost totally incompatible element in these lavas, and is, therefore, used as an index of the
degree of melting. Filled symbols represent Ba/Al2O3 and open symbols Sr/Al2O3. Samples that lie within the shaded region are those that
are likely to have been produced by melting in the presence of residual
phlogopite.
Amphibole, phlogopite and ilmenite are metasomatic
phases and their presence in the Vitim mantle xenoliths
provides important information about the nature of the
subcontinental lithospheric mantle. Available partition
coefficient data indicate that Ba is compatible in phlogopite, but only moderately incompatible in amphibole (e.g.
Sp€ath et al., 2001). Neither phlogopite nor amphibole is
capable of fractionating Sr. Figure 16 shows Ba normalized to Al2O3, plotted against the incompatible element,
Th. We chose to normalize Ba to Al2O3 because, during
melting, the ratio of an incompatible element to Al2O3
will decrease systematically with increasing melting as a
result of Al2O3 being buffered by residual garnet. If an
element is behaving compatibly when two completely
incompatible elements are plotted against each other,
this will show as a deviation from the expected linear
trend (Hoernle & Schmincke, 1993). The Vitim magmas
plotted in Fig. 16 form a clear inflection in the trend,
corresponding to the change from basanitic to alkali
basaltic magmas. In contrast, Sr behaves as a strongly
incompatible element in all magma types. A residual
phase capable of fractionating Ba but not Sr (e.g. phlogopite) must therefore be present in the source at lower
degrees of melting. Amphibole may also be residual, but
its effect will be masked by phlogopite.
Figure 17 shows the varying compatibility of elements
during melting to produce the Vitim lavas. If an element
X is totally incompatible during melting, values for X/
Al2O3 will lie on a straight line of positive slope, passing
through the origin (Hoernle & Schmincke, 1993). Deviation from this trend shows that element X is probably
hosted by a residual accessory mineral during melting.
1330
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
4.0
3.5
0.20
Rb/Al2O3
K/Al2O3
0.16
3.0
2.5
0.12
2.0
1.5
0.08
1.0
0.04
0.5
(b)
(a)
0.0
0.25
0.00
TiO2/Al2O3
0.35
Ta/Al2O3
0.30
0.20
0.25
0.15
0.20
0.10
0.15
0.05
0.00
0.00
0.10
alkali basalt
nephelinite
melanephelinite
tholeiitic basalt
basanite
(d)
(c)
0.05
0.00
0.02
0.04 0.06
P2O5/Al2O3
0.08
0.00
0.02
0.04
0.06
P2O5/Al2O3
0.08
0.10
increasing degree of melting
Fig. 17. Variation in X/Al2O3 versus P2O5/Al2O3 which is used as a proxy for the degree of melting, where X ¼ Rb (a); K (b); TiO2 (c); and Ta
(d). Where garnet is residual in the source, the P2O5/Al2O3 ratio will not be significantly affected by fractionation, and can be used as an index of
the degree of partial melting (Hoernle & Schminke, 1993). Dashed lines show the general trends of samples.
On a plot of P2O5/Al2O3 vs Rb/Al2O3 (Fig. 17a), Rb
behaves incompatibly in the alkali and tholeiitic basalts,
but there is a deviation from this trend for the nephelinites and several basanites (at P2O5/Al2O3 ¼ 004). Values
of Rb/Al2O3 are lower than expected for a given degree
of melting, which suggests that Rb is residing in a mineral
phase within the mantle source of the strongly alkaline
magmas. Partition coefficients for amphibole (pargasite)
and phlogopite show that Rb is compatible in phlogopite
(KD > 1) but incompatible in amphibole (KD < 1;
LaTourrette et al., 1995); the compatible behaviour of
Rb in the smaller-degree melts is therefore probably a
result of residual phlogopite. K (Fig. 17b) shows a similar
distribution to Rb, and it is also more compatible in
phlogopite than in amphibole. TiO2 (Fig. 17c) also
behaves similarly to Rb and K in the Vitim nephelinitic
and basanitic magmas, suggesting that a Ti-bearing residual phase is present in the mantle source at small melt
fractions. Ta appears to have been fractionated during
the melting that produced the nephelinitic magmas (at
P2O5/Al2O3 > 006; Fig. 17d).
In summary, there is strong evidence that phlogopite
was a residual mantle phase during the melting process
that produced the Vitim nephelinitic to basanitic magmas. The inferred presence of this phase is consistent with
the occurrence of phlogopite in mantle xenoliths from the
VVF (Ionov et al., 1993; Litasov et al., 2000), but it also
has implications for the location of the Vitim magma
sources. Experimental studies have shown that the stability of phlogopite is dependent on volatile content (OH
and F ) as well as temperature and pressure (Foley et al.,
1986; Sato et al., 1997). Ionov et al. (1993), Litasov et al.
(2000) and Ionov (2002) used mineral equilibria studies
to obtain temperature and pressure estimates for the
equilibration of phlogopite-bearing garnet lherzolite
xenoliths from the VVF. None was found to have equilibrated at T > 1200 C and P > 30 kbar (Fig. 18). This
suggests that the minimum depth of the mechanical
boundary layer (MBL) at the beginning of magmatism
(165 Ma; Esin et al., 1995) was 100 km. For mantle of
normal potential temperature (1300 C), this MBL thickness corresponds to a thermal boundary layer (TBL)
1331
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NUMBER 7
JULY 2005
o
Temperature ( C)
500
1000
1500
50
THOLEIITIC &
ALKALI
BASALTS
Spinel
Garnet
150
4
6
US
SOLID
8
o
Tp = 1450 C
o
Tp = 1300 C
aring pe
ridotite
Garnet & garnet-spinel
lherzolite xenoliths
OTITE
PERID
DRY
Phlogopite-be
200
250
2
L
MB
NEPHELINITES,
MELANEPHELINITES
& BASANITES
0.3%H2O + 2.5%
CO2 peridotite
Depth (km)
100
2000
Pressure (GPa)
0
Fig. 18. Melt generation pressure and temperature estimates for different magma types in the VVF, from forward major element and REE
inversion models (see text). The dry peridotite solidus is from McKenzie & Bickle (1988) and the spinel–garnet transition is from Klemme &
O’Neill (2000). The solidi for phlogopite-bearing peridotite and 03% H2O þ 25% CO2 peridotite are from Sato et al. (1997) and Wallace &
Green (1988), respectively. Mantle adiabats for Tp of 1300 C and 1450 C, and the thickness of the TBL for a MBL thickness of 100 km are from
McKenzie & Bickle (1988). The arrow indicates the change in thickness of the MBL (suggested by the modelling results) as extension progresses.
The field for garnet and garnet–spinel lherzolite xenoliths contains data from Ionov et al. (1993), Ashchepkov et al. (1994) and Litasov et al. (2000).
thickness of 35 km (McKenzie & Bickle, 1988).
Figure 18 shows that at these depths, phlogopite is stable
in both the MBL and TBL but not the convecting
mantle. The presence of residual phlogopite in the
mantle source of the nephelinites and basanites indicates
that they were generated at T < 1300 C. [The definitions that we have used for the lithosphere are from
McKenzie & Bickle (1988) and White (1988). These
state that: (1) the TBL is a transition zone that is neither
totally rigid nor vigorously convecting, and separates
the rigid plate (crust and subcontinental lithospheric
mantle) from the underlying convecting mantle; (2) the
MBL is the outer layer of the Earth, which responds
elastically to a depth where the temperature reaches
550–600 C. Beneath this, the MBL behaves plastically
to long-term loads. This definition of the MBL differs
from those of Jordan et al. (1989) and Anderson (1994),
who used the term ‘MBL’ as equivalent to only elastic
thickness of the lithosphere, which approximately corresponds to the 600 C isotherm.]
MAGMA GENESIS
Here, we use independent geochemical modelling techniques to assess the composition of the contributing melt
source regions, and to suggest the depth and extent of
partial melting, beneath the VVF.
1332
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Alkali and tholeiitic basalts
We have used a forward major element modelling
method, similar to that developed by Langmuir et al.
(1992) for MORB, but with modifications to make it
more appropriate for lavas from the VVF. The resulting
model calculates major element compositions (FeO and
Na2O) for polybaric and isobaric melting paths for
plausible mantle sources, and compares these with the
major element abundances in the VVF rocks. Model
parameters are given in Table B1, and melting equations
in Table B2 (both in Appendix B). From this point
onwards, the modified model will be referred to as the
‘hybrid’ model, as it also incorporates some of the ideas
from the melting models of Kostopoulos & James (1992).
The original Langmuir et al. (1992) model was based on
Fe–Mg partitioning, assuming that Fe is partitioned only
into residual olivine (it was used for modelling melt generation at shallow depths where garnet is not a residual
phase). However, during melting beneath the Vitim
region, MREE/HREE ratios indicate that garnet is a
significant residual phase (see above), and the hybrid
model is designed to incorporate this. It also uses melting
proportions (from Kostopoulos & James, 1992) to recalculate the residue composition at each stage of melting.
The output from the hybrid melting model applies only
to primary melts and the Vitim samples must be extrapolated to primary compositions in order to make a
quantitative comparison. For this fractionation correction, we used the method of Turner & Hawkesworth
(1995), in which data are corrected by fitting least squares
linear regression lines. Figure 6 shows that VVF magma
compositions converge towards a common value at
12 wt % MgO. This value is therefore likely to represent the MgO content of a primary magma. Only samples
with MgO > 10 wt % are projected along the regression
vectors, in order to minimize the effects of fractional
crystallization.
Initially, we assumed an anhydrous fertile peridotite
source (KLB-1) which has moderate abundances of
FeO* (859 wt %) and Na2O (030 wt %) (Takahashi,
1986). Model curves for both batch melting and accumulated fractional melting are shown in Fig. 19. In accumulated fractional melting, several melt increments are
collected together in a common reservoir, after isolation
from the melt source. The results of our polybaric melting
calculations are shown in Fig. 19a. The Na and Fe contents of the Vitim tholeiitic basalt melts can be modelled
by 7% decompression melting of KLB-1 between 35
and 33 kbar, i.e. from 115 to 110 km, assuming an
increase of 1 kbar pressure for every 33 km depth
increase in the mantle. Similarly, the Na and Fe contents
of the alkali basalts indicate 5% partial melting
between 115 and 100 km. At 35 kbar, the KLB-1 solidus
is at 1560 C; this corresponds to a mantle potential
temperature of 1490 C, assuming a mantle adiabat with
gradient of 06 C/km (McKenzie & Bickle, 1988).
An alternative approach to modelling the generation of
the Vitim alkali and tholeiitic basalts is the method of
REE inversion. We have followed the scheme described
by McKenzie & O’Nions (1991), and subsequently
modified by White et al. (1992), in which partial melt
distributions were obtained using whole-rock REE concentrations. It assumes that the compositions of all melts
and residues are governed by the depth and degree of
partial melting. The alkali and tholeiitic basalts were
modelled by single-stage melting of an asthenospheric
mantle source, which has eNd ¼ 45 (similar to the eNd
values of the Vitim basalts, Table 2). This is equivalent to
a mixture of 55% Primitive Mantle (PM) and 45%
Depleted Mantle (DM; McKenzie & O’Nions, 1991). A
good fit to the REE (Fig. 20a and c) and a reasonable fit
to other incompatible elements (Fig. 20b and d) was
obtained. After correction for fractionation, the inversion
modelling predicts 5 and 3% melting to generate the
tholeiitic and alkali basalts, respectively, with melting
occurring initially within the garnet stability field (between
85 and 105 km; Fig. 20e). The predicted melt distribution closely follows the decompression-melting curve
for a mantle potential temperature of 1450 C (Fig. 20e).
Nephelinites, melanephelinites and
basanites
The results of experimental studies suggest that silicaundersaturated melts, such as nephelinites, melanephelinites and basanites, may be generated by: (1) highpressure, small-degree melting of fertile peridotite
(Takahashi & Kushiro, 1983; Zhang & Herzberg,
1994), (2) melting of fertile peridotite in the presence of
CO2 or H2O (Hirose, 1997; Kawamoto & Holloway,
1997), (3) melting of an amphibole or phlogopite-bearing
peridotite (Wallace & Green, 1988; Mengel & Green,
1989; Thibault et al., 1992), or (4) partial melting of
garnet pyroxenite (Yaxley & Green, 1998; Hirschmann
et al., 2003; Kogiso et al., 2003). Figure 19a suggests that
polybaric melting of a fertile peridotite source (such as
KLB-1) cannot generate the high Na and Fe contents of
the VVF nephelinites, melanephelinites and basanites,
even at the smallest degrees of partial melting. In addition, both high-pressure melts of fertile peridotite and
those generated in the presence of volatiles (from fertile
peridotite) have considerably higher MgO contents [and
Ca(Ca þ Mg) ratios] and lower Al2O3 contents than the
VVF silica-undersaturated magmas. These discrepancies
in MgO and Al2O3 cannot be attributed to post-melt
generation processes, such as fractional crystallization,
and we therefore believe that melts derived from partial
melting of fertile peridotite are unlikely to have made
1333
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
4.5
3.5
Na2O
3.0
Po=35 kbar
To=1560oC
Tp=1490oC
Po=30 kbar
To=1500oC
o
Tp=1440 C
melanephelinites
nephelinites
basanites
alkali basalts
tholeiitic basalts
4.0
JULY 2005
1%
Po=40 kbar
o
To=1613 C
Tp=1534oC
2.5
6%
Po=45 kbar
To=1657oC
Tp=1568oC
2.0
1.5
12%
1.0
18%
0.5
(a) Polybaric melting of
fertile peridotite (KLB-1)
0.0
3
4
5
6
7
8
9
10
11
12
13
14
FeO
4.5
melanephelinites
nephelinites
basanites
alkali basalts
tholeiitic basalts
4.0
3.5
30 kbar
35 kbar
1%
Na2O
3.0
40 kbar
2.5
45 kbar
6%
2.0
12%
1.5
18%
24%
1.0
0.5
(b) Isobaric melting of fertile peridotite (KLB-1)
0.0
7
8
9
10
11
12
13
FeO
10
9
1%
8
25 kbar
7
Na2O
20 kbar
6%
6
12%
5
30 kbar
18%
4
50 kbar
30%
3
2
1
melanephelinites
nephelinites
basanites
(c) Isobaric melting of pyroxenite (MIX1G)
0
8
9
10
11
12
FeO
1334
13
14
15
16
17
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
significant contributions to the silica-undersaturated
magmas.
Pyroxenite xenoliths from the VVF were described by
Litasov et al. (2000). We consider that pyroxenites within
the lithosphere could be the dominant melt source region
for the silica-undersaturated magmas. We have explored
this possibility by plotting the fractionation-corrected
compositions of the nephelinites, melanephelinites and
basanites from the VVF together with those of experimental melts of garnet pyroxenite on a SiO2 versus FeO
diagram (Fig. 21). It can be seen from this that the VVF
silica-undersaturated melts and garnet pyroxenite melts
have very similar compositions.
We have used the whole-rock composition of garnet
pyroxenite, MIX1G (Hirschmann et al., 2003) in our
hybrid melting models. This has FeO* and Na2O contents of 780 and 140 wt %, respectively. The polybaric
melting model that we discussed above, for the tholeiitic
and alkali basalts, assumes that melting occurred by adiabatic decompression of the convecting mantle. However,
this mechanism may not be strictly appropriate for generation of the nephelinitic, melanephelinitic and basanitic
magmas if their contributing melts were generated from
metasomatic veins in the lithospheric mantle (as we have
inferred above on the basis of residual phlogopite).
Because a vein consisting of different mineral phases
may be melted in discrete increments at the same pressure, an isobaric melting model is probably more
suitable for generation of these silica-undersaturated
melts beneath the VVF.
We tested isobaric melting using the hybrid model,
with both KLB-1 and MIX1G as source compositions.
Isobaric melting of KLB-1 (Fig. 19b) cannot reproduce
the high Fe (or Na) contents of some of the Vitim melts.
A closer fit to these Vitim data is achieved, however, by
isobaric melting of MIX1G (Fig. 19c). The melanephelinites and basanites predominantly fall between the
MIX1G melting curves at 20 and 25 kbar, and at 18–
30% melting. The fractionation-corrected nephelinites
have higher Fe contents and fall along the 25 kbar melting curve, at similar or lower degrees of melting. From
this, we conclude that the nephelinites were formed by
partial melting at slightly higher pressures than the basanites and melanephelinites. This is consistent with the
conclusions of Hirschmann et al. (2003), who showed that
the most Fe-rich silica-undersaturated magmas were
formed by melting garnet pyroxenite at higher pressures
and/or lower temperatures than magmas with lower Fe
contents.
Despite the close correlation between most of the
major element abundances in the VVF silicaundersaturated melts and those of experimental partial
melts of garnet pyroxenite (Hirschmann et al., 2003), we
note that the former have much higher contents of
K2O, TiO2 and P2O5. We have shown above that both
phlogopite and garnet were present as residual phases
in the melt source region of the silica-undersaturated
VVF melts and propose that they were generated
from a phlogopite-bearing garnet pyroxenite source.
Hirschmann et al. (2003) showed that, at 25 kbar, the
garnet pyroxenite solidus is between 1375 and 1400 C.
However, this initial melting temperature would be
reduced and the pressure increased if the amount of
K2O was higher, as a result of of the presence of phlogopite
(Tsuruta & Takahashi, 1998; Wang & Takahashi, 2000).
DISCUSSION
Implications for the sub-Vitim
lithospheric mantle
The results of our geochemical modelling of the tholeiitic
and alkali basalts from the VVF indicate that adiabatic
decompression melting (up to 7%) of the convecting
mantle occurred between 85 and 105 km (from inversion modelling), or 100 and 115 km (from major
element modelling). We have assumed that the top of
the melting column was controlled by the thickness of
the overlying rigid MBL.
Oligocene to Recent lithospheric extension associated
with the Baikal Rift appears to have reactivated Mesozoic
rift structures beneath the region and may have caused
Fig. 19. Variation of FeO and Na2O during (a) polybaric batch and accumulated fractional melting of a fertile peridotite source, KLB-1
(Takahashi, 1986); (b) isobaric batch and accumulated fractional melting of KLB-1; and (c) isobaric batch and accumulated fractional melting of
garnet pyroxenite, MIX1G (Hirschmann et al., 2003). Dashed and continuous melting curves represent accumulated fractional melting and batch
melting, respectively. Where melting curves for accumulated and batch melting are indistinguishable, the latter are not shown. Values of FeO and
Na2O for whole-rock compositions have been extrapolated to primary melt compositions by normalizing to 12 wt % MgO, using regression
analysis (see text). Only Vitim lavas with >10 wt % MgO (un-normalized) are shown. In (a), each point along the model curve represents a
decrease in pressure and corresponding increase in the degree of melting, which is assumed to be 12% per kilobar decrease in pressure (Langmuir
et al., 1992). Po and To represent the pressure and temperature, respectively, at which the ascending mantle first intersects the solidus. Mantle
temperatures corresponding to these pressures were estimated from the data of Hirose & Kushiro (1993). Tp is the mantle potential temperature,
calculated using To and a mantle adiabat with gradient of 06 C/km (McKenzie & Bickle, 1988). Melting equations are from Rollinson (1993).
Partition coefficient data for KLB-1 are from (1) for Fe: Herzberg & Zhang (1996) and Gibson et al. (2000), and (2) for Na: Langmuir et al. (1992),
Herzberg & Zhang (1996) and Taura et al. (1998). Melting proportions and mineral assemblages for garnet and spinel lherzolite are from
Kostopoulos & James (1992) and Smith et al. (1993). In (c), partition coefficient data and solidus temperatures and pressures for MIX1G are from
Hirschmann et al. (2003) and Kogiso et al. (2003). A modal mineralogy of 55% clinopyroxene and 45% garnet was assumed for MIX1G (M.M.
Hirschmann, personal communication, 2004).
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JOURNAL OF PETROLOGY
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Fig. 20. (a–d) Best-fit single-stage melt distributions (for decompression melting) for tholeiitic basalts (a and b) and alkali basalts (c and d) from the
VVF, based on a mantle source with eNd ¼ 45, which is equivalent to a mixture of 55% Primitive Mantle (PM) and 45% Depleted Mantle (DM)
(McKenzie & O’Nions, 1991). (e) Variation in the melt fraction with depth for tholeiitic and alkali basalts. The garnet–spinel stability field is at
85–90 km and is from Klemme & O’Neill (2000). Dotted lines representing melt-distribution curves for isentropic decompression paths for 1450
and 1550 C are taken from White & McKenzie (1995). The dashed and continuous lines representing the melt distribution are corrected for
fractionation (the inversion model output gives values of 122 and 178% for fractionation in tholeiitic and alkali basalts, respectively). On all plots,
circles are mean sample element ratios (normalized to DM) and error bars are for the standard deviation plus the estimated error in source
compositions.
localized thinning of the MBL to 85 km. Our findings
are consistent with geophysical estimates, which suggest
that the base of the MBL is at 100 km depth (Burov
et al., 1994) beneath the whole of the Vitim region.
Furthermore, no mantle xenoliths from the VVF appear
to have equilibrated at depths below 80–100 km (Ionov
et al., 1993; Litasov et al., 2000; Fig. 18). Model age
estimates from the available Nd and Os isotopic data
1336
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Fig. 21. SiO2 vs FeO (wt %) for the silica-undersaturated magmas
(melanephelinites, nephelinites and basanites) from Vitim, in comparison with data from experimental melts of fertile peridotite (KLB-1)
and garnet pyroxenite (MIX1G). Data are corrected to 12 wt % MgO,
and only Vitim lavas with >10 wt % MgO are shown, in order to
minimize fractionation effects. FeO ¼ 09 FeO*. The data for the
experimental melt fields are as follows: KLB-1 (Takahashi, 1986);
KLB-1 þ CO2 (Hirose, 1997); KLB-1 þ H2O (Kawamoto &
Holloway, 1997); MIX1G (Hirschmann et al., 2003; Kogiso et al.,
2003). The numbers shown refer to the pressures (in GPa) of the
melting experiments.
(Ionov et al., 1993; Pearson et al., 1998, 2003; D. A. Ionov,
personal communication, 2003) suggest that the
peridotite xenoliths analysed to date only sample the
MBL of the lithospheric mantle.
We have shown above that the nephelinitic and
basanitic melts from the VVF were formed by partial
melting of a phlogopite-bearing garnet pyroxenite source.
Such material can reside in the convecting mantle as
‘streaks’ (Gibson, 2002) or in the TBL and/or MBL as
veins. Our estimated depths of garnet pyroxenite melting
(83–66 km; 25–20 kbar) are less than those for the top of
the melting column in the convecting mantle, and this
confirms our earlier suggestion (based on phlogopite stability) that the basanites and nephelinites were generated
in the base of the MBL and/or the TBL.
Litasov et al. (2000) proposed a model in which melt,
derived from pyroxenite veins at depth within the subVitim MBL, trickles upwards and subsequently crystallizes at lower pressures. These veins are re-melted at a
later stage of rifting, mobilized, and then erupted at the
Earth’s surface. In order to melt metasomatic veins in the
MBL, heat must be transferred by conduction from the
asthenospheric mantle or by advection from asthenospheric melts (McKenzie, 1989). At high mantle potential
temperatures, metasomatized zones at the base of the
MBL would melt in less than 10 Myr by heat conduction
(Roberts, 2002), and heat advected by rising melts would
cause immediate melting. Both processes would lead to
the occurrence of small volumes of silica-undersaturated
magmas interbedded with Miocene to Recent
tholeiitic and alkali basalts in the VVF. Nevertheless, we
note that the silica-undersaturated magmas in the
VVF have different Sr, Nd and Hf isotopic ratios
from those of entrained peridotite mantle xenoliths (Figs 11 and 12). This suggests that either garnet
pyroxenite veins in the sub-Vitim MBL have different
Sr, Nd and Hf isotopic ratios from the peridotite
xenoliths or that the magmas were derived from garnet
pyroxenites in the TBL.
During periods of tectonomagmatic quiescence,
continental lithospheres will develop an underlying TBL
by conductive cooling (McKenzie & Bickle, 1988). Thus,
the sub-Vitim lithosphere would have undergone conductive thickening during the interval between the earlier
phase of Mesozoic magmatism (Rasskazov, 1994) and the
Miocene. Although the TBL is likely to undergo convective overturn on a time-scale of >10 Myr (McKenzie &
O’Nions, 1995), its uppermost few kilometres will characteristically remain stable for several million years at
temperatures (by definition) lower than those of the
underlying convecting mantle. This thin zone at the top
of the TBL is an ideal place for batches of very-smallfraction incipient melts to solidify as they leak from
the asthenosphere (McKenzie, 1989; Wilson et al.,
1995). Thompson et al. (2005) have suggested that substantial amounts of such melts can accumulate within
less than 20 Myr.
Experiments conducted by Yaxley & Green (1998)
have shown how veined mantle in the TBL might melt.
Initial fusion of the veins produces Mg-poor liquids that
react rapidly with the surrounding peridotite, enriching it
in garnet and clinopyroxene. When this reaction zone in
turn begins to melt with rising temperature, the liquids
produced are picritic nephelinites and basanites with
higher Fe contents than anhydrous peridotite melts.
This process might explain the relative Fe enrichment
of the strongly alkalic magmas observed in the VVF.
Because all these processes would have taken place
beneath Vitim within 10 Myr or less, the basanites and
nephelinites would be expected to retain their OIB-like
Sr–Nd–Hf isotopic ratios.
Implications for the cause of
mantle melting
We have estimated the potential temperature (Tp) of the
convecting mantle beneath the VVF in our forward
major element models (1480 C; Fig. 19a) and REE inversion models (1450 C; Fig. 20e). A mantle potential temperature of 1450 C is considerably hotter than
ambient mantle, which has Tp 1300 C [Thompson
& Gibson (2000), based on the calculations for the entropy
of melting by Kojitani & Akaogi, (1995)]. It is therefore
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JOURNAL OF PETROLOGY
VOLUME 46
necessary to suggest a source for the excess heat during
the Cenozoic. Three possibilities were put forward by
Barry et al. (2003) in relation to the petrogenesis of
contemporaneous Mongolian basalts: (1) the Asian continent may have been acting as a thermal blanket,
causing the upper mantle to warm up; (2) a large-scale
deep mantle plume beneath Asia allowed hot asthenospheric material to reach shallow depths by feeding it
into ‘thinspots’ on the base of the lithosphere (Thompson
& Gibson, 1991); (3) a small-scale plume was active
beneath the Baikal region during the early stages of
magmatism but, after this, only the cooling head
remained. The teleseismic tomographic data of Petit
et al. (1998) strongly support hypothesis (3). They located
a relatively narrow (100–200 km diameter) mantle plume
(seismically slow) rising from at least 600 km depth
beneath the Siberian Craton and Baikal Rift axis.
Mantle plumes of similar dimensions have been located
beneath parts of the Central European rift system
(Granet et al., 1995; Ritter et al., 2001).
Artemieva & Mooney (2001) suggested that the
regional base of the TBL beneath Baikal is between
110 and 125 km. The VVF is located above relatively
thin crust (35 km thick), which extends in a zone
( 200 km wide) NE from Lake Baikal (Suvorov et al.,
2002). This may correspond to localized thinning of the
underlying lithospheric mantle. The VVF is also located
above a Mesozoic rift system and reactivation of this may
explain the preferential location of volcanism at Vitim,
rather than in the axial zone of the Baikal Rift. In addition, there are many faults in the VVF (Fig. 2) that may
have aided uprise of magmas in this part of the rift zone.
Figure 22 illustrates our main conclusions and their
implications. The concept of mantle plume upwelling
and outflow beneath the BRZ is taken from fig. 9 of
Petit et al. (1998). Mantle xenolith studies show that the
geothermal gradient changed during the later Cenozoic
beneath the VVF (Ionov, 2002). The Miocene geotherm
is 100 C colder at a given pressure than the Pleistocene
geotherm, and this indicates that heating of the lower
lithosphere occurred during the late Cenozoic. Our study
suggests that this was caused by both lithospheric
thinning and a mantle plume, rather than lithospheric
extension alone.
CONCLUSIONS
Extension-related magmatism in the VVF occurred
during the Cenozoic, near the boundary between the
eastern margin of the Siberian craton and the Sayan–
Baikal fold belt. The VVF consists of 5000 km3 of
melanephelinite, nephelinite, basanite, alkali basalt and
tholeiitic basalt lavas. The nephelinites generally occur
towards the top of the lava pile and represent a relatively
NUMBER 7
JULY 2005
small volume of the overall volcanic succession. A comparison between the CIPW normative compositions of
the VVF magmas and experimental studies of basalts
suggests that their parental magmas had undergone
polybaric fractional crystallization in the sub-Vitim
crust prior to eruption.
All of the magmas have similar 87Sr/86Sri (0704–
0705), 143Nd/144Ndi (05127–05129) and 176Hf/177Hfi
ratios (02829–02830). This suggests that their parental
melts were derived from the convecting mantle and/or
the recently enriched base of the lithospheric mantle.
Major and trace element abundances and Sr, Nd and
Hf-isotope systematics, combined with geochemical
modelling, suggest that the source for the melanephelinitic, nephelinitic and basanitic magmas is predominantly
the sub-Vitim lithospheric mantle. The estimated
composition of the primary silica-undersaturated melts
corresponds to those generated in partial melting experiments of garnet pyroxenite between 20 and 25 kbar. Ba/
Sr ratios combined with relative depletions in K on
normalized multi-element plots suggest that phlogopite
was a residual phase in the melanephelinite, nephelinite
and basanite mantle source. We envisage that melting
occurred at the base of the MBL and/or top of the TBL.
The precise nature of this mantle source awaits the isotopic study of pyroxenite xenoliths from the VVF. The
results of our geochemical modelling agree well with
geophysical estimates for the thickness of the MBL
( 100 km; Burov et al., 1994) and the TBL
(110–125 km; Artemieva & Mooney, 2001) beneath the
eastern flank of the Baikal Rift. In support of this, no
mantle xenoliths from the VVF have been found to come
from depths greater than 100 km (Ionov et al., 1993;
Litasov et al., 2000).
Both forward major element and REE inversion
models indicate that the VVF alkali and tholeiitic basalts
are the product of larger degrees of adiabatic decompression melting (up to 7%) of a fertile peridotite source at
between 115 and 85 km depth. The high mantle potential
temperatures ( 1450 C) that we calculated suggest that
melting occurred in the convecting mantle. We find it
difficult to explain the presence of such anomalously hot
mantle beneath the BRZ without invoking a mantle
plume, and this concept is supported by the teleseismic
tomographic data of Petit et al. (1998).
The location of melt generation beneath the VVF may
have been influenced by the relatively thin underlying
lithosphere, caused by reactivation of Mesozoic rift structures. A deep fault or faults may have aided the uprise of
magmas beneath Vitim. The general lack of volcanism
beneath the Baikal Rift could be due to the thicker crust
than at Vitim; some magma may have underplated the
axial zone of the rift at Moho depths [as Petit et al. (1998)
suggested, based on seismic evidence]. The thick sediment infill in the rift axis (Logatchev & Zorin, 1987)
1338
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Fig. 22. Schematic diagram, summarizing the petrogenesis of the lavas from the VVF. The diagram is not drawn to scale in the horizontal
dimension. The ellipses in the lithospheric mantle represent metasomatic veins, which are thought to contribute to the petrogenesis of the silicaundersaturated melts. The stippled fill for three of the volcanoes highlights those that are silica-undersaturated in composition.
might also have prevented magma eruption. The widespread contemporaneous magmatic activity, high heat
flow, elevated geotherms and uplifted topography in
northern Mongolia and the Baikal region are consistent
with the presence of a mantle plume and active rifting
inferred by our study.
ACKNOWLEDGEMENTS
S. V. Rasskazov, I. V. Ashchepkov and A. V. Ivanov
organized and assisted our fieldwork and sample
collection in Siberia. We thank William and Mary
Knowles for their hospitality in Moscow, and Ivan
Mahotkin for his invaluable help with extracting and
exporting our samples from Russia. We are also grateful
to Ron Hardy, Chris Ottley and Stephen Reed for help
with geochemical analyses, and to Dan McKenzie and
Paula Smith for their assistance with inversion modelling.
Dmitri Ionov and Graham Pearson gave helpful discussions and access to unpublished data. We thank Paula
Smith for her perceptive comments on an earlier version
of the manuscript. Reviews by Andy Saunders, Robert
Trumbull and an anonymous reviewer, together with
the editorial comments of Marjorie Wilson, have
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JOURNAL OF PETROLOGY
VOLUME 46
substantially improved the manuscript. The Royal
Society generously funded SAG and RNT for fieldwork
in Siberia. This work was supported by NERC studentship GT04/98/46/ES to JSJ, and the Department of
Earth Sciences (University of Cambridge), Cambridge
Philosophical Society and Clare College, Cambridge.
This is Department of Earth Sciences, University of
Cambridge contribution no. 8071.
SUPPLEMENTARY DATA
Supplementary data for this paper are available on
Journal of Petrology online.
REFERENCES
Anderson, D. (1994). Sublithospheric mantle as the source of
continental flood basalts: The case against the continental
lithosphere and plume head reservoirs. Earth and Planetary Science
Letters 123, 269–280.
Artemieva, I. M. & Mooney, W. D. (2001). Thermal thickness and
evolution of Precambrian lithosphere: a global study. Journal of
Geophysical Research 106, 16387–16414.
Ashchepkov, I. V., Litasov, Y. D. & Dobretsov, N. L. (1994).
Pyroxenites and composite garnet peridotite xenoliths from
picrite–basalt. Vitim Plateau (Transbaikal): implication for thermobarometry and mantle reconstruction. Proceedings of the Vth International
Kimberlite Conference, Rio de Janeiro, Brazil 1/A, 455–466.
Balijinnyam, I., Bayasgalan, A., Borisov, B. A., Cisternas, A.,
Dem’yanovich, M. G., Ganbaatar, L., Kochetkov, V. M.,
Kurushin, R. A., Molnar, P., Philip, H. & Vashchilov, Y. Y.
(1993). Ruptures of Major Earthquakes and Active Deformation in Mongolia
and its Surroundings. Geological Society of America, Memoir 181, 62 pp.
Barry, T. L. & Kent, R. W. (1998). Cenozoic magmatism in Mongolia
and the origin of central and East Asian basalts. In: Flower, M. J.,
Chung, S. -L., Lo, C. -H. & Lee, T. -Y. (eds) Mantle Dynamics and Plate
Interactions in East Asia. American Geophysical Union, Geodynamics Series 27,
347–364.
Barry, T. L., Saunders, A. D., Kempton, P. D., Windley, B. F.,
Pringle, M. S., Dorjnamjaa, D. & Saandar, S. (2003). Petrogenesis
of Cenozoic basalts from Mongolia: evidence for the role of
asthenospheric versus metasomatized lithospheric mantle sources.
Journal of Petrology 44, 55–91.
Brey, G. P., Kohler, T. & Nickel, K. G. (1990). Geothermobarometry
in 4-phase lherzolites. 1. Experimental results from 10 to 60 kb.
Journal of Petrology 31, 1313–1352.
Burov, E. B., Houdry, F., Diament, M. & Deverchere, J. (1994). A
broken plate beneath the North Baikal Rift Zone revealed by gravity
modelling. Geophysical Research Letters 21, 129–132.
Delvaux, D., Moeys, R., Stapel, G., Petit, C., Levi, K.,
Miroshnichenko, A., Ruzhich, V. & San’kov, V. (1997). Paleostress
reconstructions and geodynamics of the Baikal region, Central Asia,
Part 2. Cenozoic rifting. Tectonophysics 282, 1–38.
Dowall, D. P., Pearson, D. G. & Nowell, G. M. (2003). Chemical preconcentration procedures for high-precision analysis of Hf–Nd–Sr
isotopes in geological materials by plasma ionisation multi-collector
mass spectrometry (PIMMS) techniques. In: Holland, J. G. &
Tanner, S. D. (eds) Plasma Source Mass Spectrometry: Applications and
Emerging Technologies. Cambridge: Royal Society of Chemistry,
pp. 321–337.
NUMBER 7
JULY 2005
Esin, S. V., Ashchepkov, I. V., Ponomarchuk, V. A., Yamamoto, M.,
Travin, A. V. & Kiselyeva, V. Y. (1995). Petrogenesis of Alkaline
Basaltoids from the Vitim Plateau (Baikal Rift Zone). Novosibirsk:
UIGGM, Siberian Branch, Russian Academy of Sciences Press,
pp. 1–58.
Foley, S. F., Taylor, W. R. & Green, D. H. (1986). The effect of
fluorine on phase relationships in the system KAlSiO4–Mg2SiO4–
SiO2 at 28 kbar and the solution mechanism of fluorine in silicate
melts. Contributions to Mineralogy and Petrology 93, 46–55.
Gao, S., Davis, P. M., Liu, H., Slack, P. D., Zorin, Y. A.,
Logatchev, N. A., Kogan, M., Burkholder, P. D. & Meyer, R. P.
(1994). Asymmetric upwarp of the asthenosphere beneath the
Baikal Rift Zone, Siberia. Journal of Geophysical Research 99,
15319–15330.
Garcia, M. O., Hulsebosch, T. P. & Rhodes, J. M. (1995). Olivine-rich
submarine basalts from the southwest rift zone of Mauna Loa
Volcano: implications for magmatic processes and geochemical
evolution. In: Rhodes, J. M. & Lockwood, J. P. (eds) Mauna
Loa Revealed. Geophysical Monograph, American Geophysical Union 92,
219–239.
Garner, J. S. (2002). Magmatism of the Vitim Volcanic Field,
Baikal Rift Zone, Siberia. Ph.D. thesis, University of Cambridge,
231 pp.
Gibson, S. A. (2002). Major element heterogeneity in Archean to
Recent mantle plume starting-heads. Earth and Planetary Science Letters
195, 59–74.
Gibson, S. A., Thompson, R. N., Leat, P. T., Morrison, M. A.,
Hendry, G. L., Dickin, A. P. & Mitchell, J. G. (1993). Ultrapotassic
magmas along the flanks of the Oligo-Miocene Rio Grande Rift,
USA: monitors of the zone of lithospheric mantle extension and
thinning beneath a continental rift. Journal of Petrology 34,
187–228.
Gibson, S. A., Thompson, R. N. & Dickin, A. P. (2000). Ferropicrites:
geochemical evidence for Fe-rich streaks in upwelling mantle
plumes. Earth and Planetary Science Letters 174, 355–374.
Glaser, S. M., Foley, S. F. & G€
unther, D. (1999). Trace element
compositions of minerals in garnet and spinel peridotite xenoliths
from the Vitim volcanic field, Transbaikalia, eastern Siberia. Lithos
685, 1–23.
Granet, M., Wilson, M. & Achauer, U. (1995). Imaging a mantle plume
beneath the French Massif Central. Earth and Planetary Science Letters
136, 281–296.
Green, T. H., Blundy, J. D., Adam, J. & Yaxley, G. M. (2000). SIMS
determination of trace element partition coefficients between garnet,
clinopyroxene and hydrous basaltic liquids at 2–75 GPa and
1080–1200 C. Lithos 53, 165–187.
Grove, T. L. & Bryan, W. B. (1983). Fractionation of pyroxene-phyric
MORB at low pressure: an experimental study. Contributions to
Mineralogy and Petrology 84, 293–309.
Hart, S. R. & Davis, K. E. (1978). Nickel partitioning between
olivine and silicate melt. Earth and Planetary Science Letters 40,
203–219.
Hart, S. R., Hauri, E. H., Oschmann, L. A. & Whitehead, J. A. (1992).
Mantle plumes and entrainment: isotopic evidence. Science 256,
517–520.
Hawkesworth, C. J., Gallagher, K., Kelley, S., Mantovani, M.,
Peate, D. W., Regelous, M. & Rogers, N. W. (1992). Parana
magmatism and the opening of the South Atlantic. In: Storey, B. C.,
Alabaster, T. & Pankhurst, R. J. (eds). Magmatism and the Causes of
Continental Break-up. Geological Society, London, Special Publications 68,
221–240.
Herzberg, C. & Zhang, J. (1996). Melting experiments on
anhydrous peridotite KLB-1: compositions of magmas in the
1340
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
upper mantle and transition zone. Journal of Geophysical Research 101,
8271–8295.
Hirose, K. (1997). Partial melt compositions of carbonated peridotite at
3 GPa and genesis of alkali basalt magmas. Geophysical Research Letters
24, 2837–2840.
Hirose, K. & Kushiro, I. (1993). Partial melting of dry peridotites at
high pressures: determination of compositions of melts segregated
from peridotite using aggregates of diamond. Earth and Planetary
Science Letters 114, 477–489.
Hirschmann, M. M. (2000). Mantle solidus: experimental constraints
and the effects of peridotite composition. Geochemistry, Geophysics,
Geosystems 1.
Hirschmann, M. M., Kogiso, T., Baker, M. B. & Stolper, E. M. (2003).
Alkalic magmas generated by partial melting of garnet pyroxenite.
Geology 31, 481–484.
Hoernle, K. & Schmincke, H.-U. (1993). The role of partial melting in
the 15-Ma geochemical evolution of Gran Canaria: a blob model for
the Canary hotspot. Journal of Petrology 34, 599–626.
Holloway, J. R., Pan, V. & Gudmundsson, G. (1992). High-pressure
fluid-absent melting experiments in the presence of graphite: oxygen
fugacity, ferric/ferrous ratio and dissolved CO2. European Journal of
Mineralogy 4, 105–114.
Ionov, D. (2002). Mantle structure and rifting processes in the Baikal–
Mongolia region: geophysical data and evidence from xenoliths in
volcanic rocks. Tectonophysics 351, 41–60.
Ionov, D. (2004). Chemical variations in peridotite xenoliths from
Vitim, Siberia: inferences for REE and Hf behaviour in the garnetfacies upper mantle. Journal of Petrology 45, 343–367.
Ionov, D. A., Ashchepkov, I. V., Stosch, H.-G., Witt-Eickschen, G. &
Seck, H. A. (1993). Garnet peridotite xenoliths from the Vitim
volcanic field, Baikal region: the nature of the garnet–spinel
peridotite transition zone in the continental mantle. Journal of
Petrology 34, 1141–1175.
Ionov, D. A., O’Reilly, S. Y. & Ashchepkov, I. V. (1995). Feldsparbearing lherzolite xenoliths in alkali basalts from Hamar-Daban,
southern Baikal region, Russia. Contributions to Mineralogy and Petrology
122, 174–190.
Jarvis, I. & Jarvis, K. E. (1992). Plasma spectrometry in the earth
sciences: techniques, applications and future trends. Chemical Geology
95, 1–33.
Jordan, T. H., Lerner-Lam, A. L. & Creager, K. C. (1989). Seismic
imaging of boundary layers and deep mantle convection. In:
Peltier, W. R. (ed.) Mantle Convection: Plate Dynamics and Global
Dynamics (The Fluid Mechanics of Astrophysics and Geophysics, Vol. 4).
New York: Gordon and Breach, pp. 97–201.
Kawamoto, T. & Holloway, J. R. (1997). Melting temperature and
partial melt chemistry of H2O saturated mantle peridotite to 11
gigapascals. Science 276, 240–243.
Kilinc, A., Carmichael, I. S. E., Rivers, M. L. & Sack, R. O. (1983).
The ferric/ferrous ratio of natural silicate liquids equilibrated in air.
Contributions to Mineralogy and Petrology 83, 136–140.
Kiselev, A. I. (1987). Volcanism of the Baikal Rift Zone. Tectonophysics
143, 235–244.
Kiselev, A. I. & Popov, A. M. (1992). Asthenospheric diapir
beneath the Baikal Rift: petrological constraints. Tectonophysics 208,
287–295.
Klemme, S. & O’Neill, H. StC. (2000). The near-solidus transition
from garnet lherzolite to spinel lherzolite. Contributions to Mineralogy
and Petrology 138, 237–248.
Kogiso, T., Hirschmann, M. M. & Frost, D. J. (2003). High-pressure
partial melting of garnet pyroxenite: possible mafic lithologies in the
source of ocean island basalts. Earth and Planetary Science Letters 216,
603–617.
Kojitani, H. & Akaogi, M. (1995). Measurement of heat of fusion of
model basalt in the system Diopsite–Forsterite–Anorthite. Geophysical
Research Letters 22, 2329–2332.
Kostopoulos, D. K. & James, S. D. (1992). Parameterization of the
melting regime of the shallow upper mantle and the effects of
variable lithospheric stretching on mantle modal stratification and
trace-element concentrations in magmas. Journal of Petrology 33,
665–691.
Langmuir, C. H., Klein, E. M. & Plank, T. (1992). Petrological
systematics of mid-ocean ridge basalts: constraints on melt
generation beneath ocean ridges. In: Morgan, J. P., Blackman, D. K.
& Sinton, J. M. (eds). Mantle Flow and Melt Generation at MidOcean Ridges. Geophysical Monograph, American Geophysical Union 71,
183–280.
Larsen, L. M. & Pedersen, A. K. (2000). Processes in high-Mg,
high-T magmas: evidence from olivine, chromite and glass in
Palaeogene picrites from West Greenland. Journal of Petrology 41,
1071–1098.
LaTourrette, T., Hervig, R. L. & Holloway, J. R. (1995). Trace
element partitioning between amphibole, phlogopite, and basanite
melt. Earth and Planetary Science Letters 135, 13–30.
Leat, P. T., Thompson, R. N., Morrison, M. A., Hendry, G. L. &
Dickin, A. P. (1988). Compositionally diverse Miocene–Recent riftrelated magmatism in northwest Colorado: partial melting and
mixing of magmas from three different asthenospheric and
lithospheric mantle sources. Journal of Petrology, Special Lithosphere
Issue, 351–377.
Le Bas, M. J. (1989). Nephelinitic and basanitic rocks. Journal of Petrology
30, 1299–1312.
Le Maitre, R. W. (2002). Igneous Rocks: a Classification and Glossary of Terms:
Recommendations of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks. Cambridge: Cambridge
University Press, 236 pp.
Litasov, K. D., Foley, S. F. & Litasov, Y. D. (2000). Magmatic
modification and metasomatism of the subcontinental mantle
beneath the Vitim volcanic field (East Siberia): evidence from trace
element data on pyroxenite and peridotite xenoliths from Miocene
picrobasalt. Lithos 54, 83–114.
Logatchev, N. A. & Florensov, N. A. (1978). The Baikal system of rift
valleys. Tectonophysics 45, 1–14.
Logatchev, N. A. & Zorin, Y. A. (1987). Evidence and causes of the
two-stage development of the Baikal Rift. Tectonophysics 143,
225–234.
Logatchev, N. A. & Zorin, Y. A. (1992). Baikal Rift Zone: structure and
geodynamics. Tectonophysics 208, 273–286.
McDonough, W. F. & Sun, S.-s. (1995). The composition of the Earth.
Chemical Geology 120, 223–253.
McKenzie, D. (1989). Some remarks on the movement of small
melt fractions in the mantle. Earth and Planetary Science Letters 95,
53–72.
McKenzie, D. & Bickle, M. J. (1988). The volume and composition of
melt generated by extension of the lithosphere. Journal of Petrology 29,
625–679.
McKenzie, D. & O’Nions, R. K. (1991). Partial melt distributions from
inversion of rare earth element concentrations. Journal of Petrology 32,
1021–1091.
McKenzie, D. & O’Nions, R. K. (1995). The source regions of ocean
island basalts. Journal of Petrology 36, 133–159.
Mengel, K. & Green, D. H. (1989) Stability of amphibole and
phlogopite in metasomatized peridotite under water-saturated and
water-undersaturated conditions. In: Ross, J. (ed.) Proceedings of the
4th International Kimberlite Conference. Perth: Geological Society of
Australia, pp. 571–581.
1341
JOURNAL OF PETROLOGY
VOLUME 46
Molnar, P. & Tapponnier, P. (1975). Tectonics of Asia: consequences
and implications of a continental collision. Science 189, 419–426.
Morimoto, N. (1988). Nomenclature of pyroxenes. American Mineralogist
73, 1123–1133.
Nowell, G. M. & Parrish, R. R. (2001). Simultaneous acquisition of
isotope compositions and parent/daughter ratios by non-isotope
dilution solution-mode Plasma Ionisation Multi-collector Mass
Spectrometry (PIMMS). In: Holland, G. & Tanner, S. D. (eds)
Plasma Source Mass Spectrometry: The New Millennium. Royal Society of
Chemistry, Special Publication 267, 298–310.
Nowell, G. M., Kempton, P. D., Noble, S. R., Fitton, J. G.,
Saunders, A. D., Mahoney, J. J. & Taylor, R. N. (1998). High
precision Hf isotope measurements of MORB and OIB by thermal
ionization mass spectrometry: insights into the depleted mantle.
Chemical Geology 149, 211–233.
Nowell, G. M., Pearson, D. G., Ottley, C. J., Schwieters, J. &
Dowall, D. (2003). Long-term performance characteristics of a
Plasma Ionisation Multi-Collector Mass Spectrometer (PIMMS): the
ThermoFinnigan Neptune. In: Holland, J. G. & Tanner, S. D. (eds)
Plasma Source Mass Spectrometry: Applications and Emerging Technologies.
Cambridge: Royal Society of Chemistry, pp. 307–320.
Pearson, D. G., Ionov, D., Carlson, R. W. & Shirey, S. B. (1998).
Lithospheric evolution in circum-cratonic settings: a Re–Os isotope
study of peridotite xenoliths from the Vitim region, Siberia.
Mineralogical Magazine 62A, 1147–1148.
Pearson, D. G., Canil, D. & Shirey, S. B. (2003). Mantle samples
included in volcanic rocks: xenoliths and diamonds. In:
Turekian, K. K. & Holland, H. D. (eds) Treatise of Geochemistry,
Volume 2: The Mantle and Core. Amsterdam: Elsevier.
Petit, C., Koulakov, I. & Deverchere, J. (1998). Velocity structure
around the Baikal Rift Zone from teleseismic and local earthquake
traveltimes and geodynamic implications. Tectonophysics 296,
125–144.
Polyansky, O. P. (2002). Dynamic causes for the opening of the
Baikal Rift Zone: a numerical modelling approach. Tectonophysics 351,
91–117.
Potts, P. J., Tindle, A. G. & Webb, P. C. (1992). Geochemical Reference
Material Compositions. Caithness: Whittles, 320 pp.
Rasskazov, S. V. (1994). Magmatism related to the eastern Siberia rift
system and the geodynamics (Magmatisme associe au rift de Siberie
orientale: implications geodynamiques). Bulletin des Centres de
Recherches Exploration–Production Elf-Aquitaine 18, 437–452.
Ritter, J. R. R., Jordan, M., Christensen, U. R. & Achauer, U. (2001).
A mantle plume below the Eifel volcanic fields, Germany. Earth and
Planetary Science Letters 186, 7–14.
Roberts, M. A. (2002). The geochemical and volcanological evolution of
the Mt. Meru Region, Northern Tanzania. Ph.D. thesis, University
of Cambridge, 266 pp.
Roeder, P. L. & Emslie, R. F. (1970). Olivine–liquid equilibria.
Contributions to Mineralogy and Petrology 29, 275–289.
Rollinson, H. R. (1993). Using Geochemical Data: Evaluation, Presentation,
Interpretation. Harlow: Longman.
Ruppel, C. (1995). Extensional processes in continental lithosphere.
Journal of Geophysical Research 100, 24187–24215.
Sack, R. O. & Ghiorso, M. S. (1991). Chromian spinels as petrogenetic
indicators: thermodynamics and petrological applications. American
Mineralogist 76, 827–847.
Sato, K., Katsura, T. & Ito, E. (1997). Phase relations of natural
phlogopite with and without enstatite up to 8 GPa: implication for
mantle metasomatism. Earth and Planetary Science Letters 146, 511–526.
Smith, D., Griffin, W. L. & Ryan, C. G. (1993). Compositional
evolution of high-temperature sheared lherzolite PHN1611.
Geochimica et Cosmochimica Acta 57, 605–613.
NUMBER 7
JULY 2005
Snyder, J. P. (1987). Map Projections: a Working Manual. US Geological
Survey Professional Paper 1395.
Sp€ath, A., Le Roex, A. P. & Opiyo-Akech, N. (2001). Plume–
lithosphere interaction and the origin of continental rift-related
alkaline volcanism: the Chyulu Hills Volcanic Province, Southern
Kenya. Journal of Petrology 42, 765–787.
Suvorov, V. D., Mishenkina, Z. M., Petrick, G. V., Sheludko, I. F.,
Seleznev, V. S. & Solovyov, V. M. (2002). Structure of the crust in
the Baikal Rift Zone and adjacent areas from Deep Seismic
Sounding data. Tectonophysics 351, 61–74.
Takahashi, E. (1986). Melting of a dry peridotite KLB-1 up to 14 GPa:
implications on the origin of peridotitic upper mantle. Journal of
Geophysical Research 91, 9367–9382.
Takahashi, E. & Kushiro, I. (1983) Melting of a dry peridotite at high
pressures and basalt magma genesis. American Mineralogist 68,
859–879.
Taura, H., Yurimoto, H., Kurita, K. & Sueno, S. (1998). Pressure
dependence on partition coefficients for trace elements between
olivine and the coexisting melts. Physics and Chemistry of Minerals 25,
469–484.
Taylor, S. R. & McClennan, S. M. (1985). The Continental Crust: Its
Composition and Evolution. Oxford: Blackwell, 312 pp.
Thibault, Y., Edgar, A. D. & Lloyd, F. E. (1992). Experimental
investigation of melts from a carbonated phlogopite lherzolite:
implications for metasomatism in the continental lithosphere.
American Mineralogist 77, 784–794.
Thompson, R. N. (1982). Magmatism of the British Tertiary Volcanic
Province. Scottish Journal of Geology 18, 49–107.
Thompson, R. N. (1983). Thermal aspects of the origin of
Hebridean Tertiary acid magmas, II. Experimental melting
behaviour of the granites at 1 kbar PH2O. Mineralogical Magazine 47,
111–121.
Thompson, R. N. & Gibson, S. A. (1991). Subcontinental mantle
plumes, hotspots and pre-existing thinspots. Journal of the Geological
Society, London 148, 973–977.
Thompson, R. N. & Gibson, S. A. (2000). Transient high temperatures
in mantle plume heads inferred from magnesian olivines in
Phanerozoic picrites. Nature 407, 502–506.
Thompson, R. N., Leat, P. T., Morrison, M. A., Hendry, G. L.,
Dickin, A. P. & Gibson, S. A. (1990). Strongly potassic mafic
magmas from lithospheric mantle sources during continental
extension and heating: evidence from Miocene minettes of northwest
Colorado, USA. Earth and Planetary Science Letters 38, 139–153.
Thompson, R. N., Gibson, S. A., Dickin, A. P. & Smith, P. M. (2001).
Early Cretaceous basalt and picrite dykes of the Southern Etendeka
Region, NW Namibia: windows into the role of the Tristan mantle
plume in Parana–Etendeka Magmatism. Journal of Petrology 42,
2049–2081.
Thompson, R. N., Ottley, C. J., Smith, P. M., Pearson, D. G., Dickin,
A.- P., Leat, P. T., Morrison, M. A. & Gibson, S. A. (2005). Source
of the Quaternary alkalic basalts, picrites and basanites of the
Potrillo Volcanic Field, New Mexico, USA—Lithosphere or conrecting mantle? Journal of Petrology (in press).
Tsuruta, K. & Takahashi, E. (1998). Melting study of an alkali basalt
JB-1 up to 125 GPa: behaviour of potassium in the deep mantle.
Physics of the Earth and Planetary Interiors 107, 199–130.
Turner, S. & Hawkesworth, C. (1995). The nature of the subcontinental mantle: constraints from the major-element composition
of continental flood basalts. Chemical Geology 120, 295–314.
Vervoort, J. D., Patchett, P. J., Blichert-Toft, J. & Albarede, F. (1999).
Relationships between Lu–Hf and Sm–Nd isotopic systems in the
global sedimentary system. Earth and Planetary Science Letters 168,
79–99.
1342
JOHNSON et al.
CENOZOIC BAIKAL RIFT-RELATED MAGMATISM
Wallace, M. E. & Green, D. H. (1988). An experimental determination of primary carbonatite magma composition. Nature 335,
343–346.
Wang, W. & Takahashi, E. (2000). Subsolidus and melting experiments
of K-doped peridotite KLB-1 to 27 GPa: its geophysical and
geochemical implications. Journal of Geophysical Research 105,
2855–2868.
White, R. S. (1988). The Earth’s crust and lithosphere. Journal of
Petrology, Special Lithosphere Issue, 1–10.
White, R. S. & McKenzie, D. (1995). Mantle plumes and flood basalts.
Journal of Geophysical Research 100, 17543–17585.
White, R. S., McKenzie, D. & O’Nions, R. K. (1992). Oceanic
crustal thickness from seismic measurements and rare earth
element inversions. Journal of Geophysical Research 97,
19683–19715.
Wilson, M. (1989). Igneous Petrogenesis. London: Chapman & Hall,
466 pp.
Wilson, M., Rosenbaum, J. M. & Dunworth, A. D. (1995). Melilitites:
partial melts of the thermal boundary layer? Contributions to Mineralogy
and Petrology 119, 181–196.
Windley, B. F. & Allen, M. B. (1993). Mongolian plateau: evidence for
a late Cenozoic mantle plume under central Asia. Geology 21,
295–298.
Wood, D. A. (1979). Dynamic partial melting: its application to the
petrogenesis of basalts erupted in Iceland, the Faeroe Islands, the Isle
of Skye (Scotland) and the Troodos Massif (Cyprus). Geochimica et
Cosmochimica Acta 43, 1031–1046.
Yaxley, G. M. & Green, D. H. (1998). Reactions between eclogite
and peridotite: mantle refertilisation by subduction of oceanic
Drilling region
Drill-hole
Antasey
4363
X
Y
60280
308940
crust. Schweizerische Mineralogische und Petrographische Mitteilungen 78,
243–255.
Zhang, Y.-S. (1998). Three dimensional upper mantle structure
beneath east Asia and its tectonic implications. In: Flower, M. J.,
Chung, S.-L., Lo, C.-H. & Lee, T.-Y. (eds) Mantle Dynamics and Plate
Interactions in East Asia. American Geophysical Union, Geodynamic Series 27,
11–24.
Zhang, J. & Herzberg, C. (1994). Melting experiments on anhydrous
peridotite KLB-1 from 50 to 225 GPa. Journal of Geophysical Research
99, 17729–17742.
Zorin, Y. A. (1981). The Baikal Rift: an example of the
intrusion of asthenospheric material into the lithosphere as a cause
of disruption of the lithospheric plates. Tectonophysics 73,
91–104.
APPENDIX A: COORDINATE SYSTEM
USED IN THE VITIM VOLCANIC FIELD
The table below shows the original coordinate data,
the estimated Universal Transverse Mercators (UTMs)
and the corresponding latitude/longitude values for the
drill-holes. The location of the drilling region marked
‘???’ is unknown. A value of þ111 was used for the
Central Meridian, which places the Vitim region in
zone 49 U of the UTM grid (Snyder, 1987). Numbers in
bold type are those that have been changed from the
original data.
UTM N (m)
UTM E (m)
Latitude ( N)
Longitude ( E)
5960280
630894
53.774
53.817
112.986
112.981
Antasey
4403
965073
304458
5965073
630445.8
Antasey
4404
965878
304585
5965878
Antasey
4431
960554
296413
5960554
630458.5
629641.3
53.824
53.776
112.982
112.967
Atalanga
4102
25610
60480
5925610
576048
Atalanga
4132
27065
41970
5927065
574197
53.473
53.486
112.146
112.118
Bortovoy
3833
5940694
6283119
5940694
596283.1
Burulzay
4770
51600
58115
5951600
581150
53.605
53.706
112.455
112.229
Burulzay
4772
49965
58630
5949965
586300
Centralni
3043
5949094
6269812
5949094
53.690
53.681
112.307
112.468
Centralni
3046
5941945
627779
5941945
596981.2
596277.8
Ekzar
4059
57570
26150
5957570
592615
53.617
53.758
112.455
112.405
Hoygot
4563
82810
96575
5982810
Hoygot
4569
84260
303120
5984260
623965.7
623031.2
53.978
53.991
112.890
112.877
623032.8
629189.7
54.005
53.652
53.705
112.956
113.000
53.654
53.686
112.879
113.320
53.684
53.647
113.309
113.319
53.853
112.951
Hoygot
4613
85780
303280
5985780
Kolichikan
3889
5946711
6291897
5946711
Kolichikan
4454
952665
320029
5952665
Kolichikan
4490
946765
241842
5946765
632002.9
624184.2
Muliha
4426
952323
316394
5952323
631639.4
Muliha
4633
50980
24730
5950980
632473
Muliha
4659
946905
328406
5946905
???
3690
5969091
6283321
5969091
632840.6
628332.1
1343
112.877
JOURNAL OF PETROLOGY
VOLUME 46
NUMBER 7
JULY 2005
APPENDIX B: MODEL PARAMETERS AND EQUATIONS FOR
THE ‘HYBRID’ MELTING MODEL
Table B1: Parameters used for the ‘hybrid’ melting model, including starting compositions for KLB-1 and MIX1G
Starting compositions
Na2O
FeO
KLB-1
0.3
1.4
7.73
7.02
MIX1G
Mineral assemblages
ol
opx
cpx
sp
gt
Po 30 kbaru
60.1
58.0
18.9
25.0
13.7
15.0
0
2.0
7 .3
Po < 30 kbary
0
MIX1G
Po (kbar)
T (K)v
cpx
KD Fe
sp
KD Fe
gt
KD Fe
cpx
KD Na
sp
KD Na
gt
KD Na
20
1608
—
—
0.730x
0.600x
0.30y
0.35y
0.01x
1658
0.50w
0.40w
0.92x
25
—
—
0.0050x
0.0050x
30
1743
50
1873
0.575x
0.550x
0.50y
0.80y
0.40w
0.32w
—
—
—
—
0.0085x
0.0300x
—
—
—
—
KLB-1
Po (kbar)
T (K)z
ol
KD Fe
opx
KD Fe
cpx
KD Fe
sp
KD Fe
gt
KD Fe
ol
KD Na
opx
KD Na
cpx
KD Na
sp
KD Na
gt
KD Na
15
1578
20
1643
2.020x
1.351x
1.48x
0.99x
1.7x
1.1x
0.55x
0.55x
0.85x
0.70x
0.035z
0.040z
0.010y
0.010y
0.279y
0.326y
0.07x
0.07x
0.0042x
0.0056x
25
1710
30
1773
1.061x
0.903x
0.78x
0.66x
0.9x
0.7x
0.55x
0.55x
0.65x
0.60x
0.042z
0.040z
0.008y
0.008y
0.371y
0.428y
0.07x
0.07x
0.0070x
0.0085x
35
1833
40
1886
0.805x
0.738x
0.59x
0.54x
0.7x
0.6x
0.55x
0.55x
0.50x
0.45x
0.045z
0.050z
0.008y
0.008y
0.500y
0.590y
0.07x
0.07x
0.0101x
0.0119x
45
1930
0.691x
0.51x
0.6x
0.55x
0.42x
0.055z
0.008y
0.730y
0.07x
0.0138x
Superscripts denote data sources: u, Kostopoulos & James (1992); x, Herzberg & Zhang (1996); v, Kogiso et al. (2003);
y
, Langmuir et al. (1992); w, Hirschmann (2000); z, Taura et al. (1998); y, Hirose & Kushiro (1993); z, McKenzie & Bickle (1988).
Table B2: Equations used for the ‘hybrid’ melting
model (from Rollinson, 1993)
Isobaric melting
Batch melting
Fractional melting
Definitions of terms
CL/C0 ¼ 1/[D0 þ F(1 D0)]
1
CL =C0 ¼ F1 ½1 ð1 FÞD0 Polybaric melting
Batch melting
Residue calculation
Fractional melting
CL/C0 ¼ 1/[D0 þ F(1 D0)]
CS/C0 ¼ DRS/[DRS þ F(1 DRS)]
1
CL =C0 ¼ 1 ½1 ð1 FÞD0 Residue calculation
CS =C0 ¼ ð1 FÞ D0
F
ð 1 1Þ
CL Weight concentration of a trace element in the liquid
L Average weight concentration of a trace element in
C
a mixed melt
C0 Weight concentration of a trace element in the
original unmelted solid
CS Weight concentration of a trace element in the
residual solid after melt extraction
DRS Bulk distribution coefficient of the residual solids
D0 Bulk distribution coefficient of the original solids
F
Weight fraction of melt produced during partial
melting
1344