JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
PAGES 227–253
1998
Etendeka Volcanism of the Goboboseb
Mountains and Messum Igneous Complex,
Namibia. Part II: Voluminous Quartz Latite
Volcanism of the Awahab Magma System
A. EWART1∗, S. C. MILNER2†, R. A. ARMSTRONG2‡ AND A. R. DUNCAN2
1
DEPARTMENT OF EARTH SCIENCES, THE UNIVERSITY OF QUEENSLAND, ST LUCIA, QLD. 4072, AUSTRALIA
2
DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF CAPE TOWN, RONDEBOSCH 7700, SOUTH AFRICA
RECEIVED FEBRUARY 23, 1996; REVISED TYPESCRIPT ACCEPTED JULY 30, 1997
The Goboboseb–Messum volcanic centre is the source of two
voluminous silicic eruptive sequences, the Goboboseb quartz latites
(Units I–III), and the Springbok quartz latite unit (both within
the Awahab Formation). Intrusive equivalents exist as plugs and
a laccolith peripheral to the Messum Complex. The recognition of
correlatives of these quartz latite units in the southeastern Parana´
suggests eruptive volumes of ~2320 km3 (Goboboseb units) and
6340 km3 (Springbok unit). The latter is thought to be a single
eruptive event. Phenocryst assemblages are plagioclase (An51–63),
pyroxene, titanomagnetite and apatite. Pyroxene assemblages range
from augite, to augite + pigeonite, to pigeonite, to pigeonite
± hypersthene, the assemblages changing progressively from the
Goboboseb unit through to the Springbok unit. Although pyroxene
phenocrysts from individual samples are compositionally very uniform
there is a small increase in Fe through the sequence, attributed to
decreasing temperature (± pressure). Thermometry suggests melt
temperatures >1000°C. Many plagioclases contain abundant glass
inclusions of three compositional types, thought to result from active
disequilibrium melting at magma chamber walls. Relatively small,
but systematic, changes in whole-rock composition occur stratigraphically from the lowest Goboboseb unit through to the Springbok
unit and to their correlatives in the Parana´, best shown by SiO2
(67–71%) and FeO∗ (5·4–7·4%) which increase and decrease,
respectively, in the progressively younger eruptive phases. P, Ti, Y,
Zr, Nb and Cu are positively correlated with FeO∗, whereas eSr
and Pb isotope compositions correlate inversely with FeO∗. Crustnormalized spidergram plots indicate strong negative Sr anomalies,
accompanied by significant Eu/Eu∗ anomalies (0·62–0·67). The
quartz latite melts can be interpreted in terms of large-scale
∗Corresponding author. e-mail: [email protected]
† Present address: Geological Survey, Ministry of Mines and Energy,
P.O. Box 2168, Windhoek, Republic of Namibia.
‡ Present address: Research School of Earth Sciences, The Australian
National University, Canberra, A.C.T 0200, Australia.
assimilation–fractional crystallization (AFC)-style processes, involving high degrees of lower- and upper-crustal melting, with
thermal and material input from hybridized LTZ.L-type basaltic
magmas (Part I). Thermal source is inferred to be the Tristan
plume. The crustal end-member is thought to be the mid-Proterozoic
restite source of the Damara granites, although some shallower
crustal input is also likely. Modelling suggests the source may be
similar to A-type granites and charnockites (i.e. relatively REE
and HFSE enriched). Available seismic data suggest a simple
velocity crustal profile, possibly the result of the massive crustal and
uppermost mantle melting that accompanied the evolution of the
Awahab magma system.
KEY WORDS:
Quartz latite; AFC processes; Namibia; crustal fusion;
Tristan plume
INTRODUCTION
Volcanic rocks of the Goboboseb Mountains are a southeasterly remnant of the much larger Parana´–Etendeka
Province. The succession comprises interbedded basalts
and silicic quartz latite units and is closely associated
with the subvolcanic Messum Complex. Part I of this
contribution (Ewart et al., 1998) describes the mafic
volcanism related to the Goboboseb–Messum volcanic
centre, the most significant aspect being the identification
 Oxford University Press 1998
JOURNAL OF PETROLOGY
VOLUME 39
of two basaltic (LTZ.L and LTZ.H) series. The LTZ.H
series is dominated by a Tristan mantle plume component, whereas the LTZ.L series show strong evidence
of modification by AFC processes involving crustal
materials. In Part II, we describe the associated quartz
latite eruptives, the Goboboseb and Springbok Members
of the Awahab Formation [see Part I and Milner et al.
(1995a)]. The eruptive centre for these quartz latite units
is identified as the Messum Complex, based on a regional
sag structure centred on Messum, and the occurrence of
chemically and mineralogically equivalent quartz monzonite plugs and a laccolith immediately peripheral to the
Messum Complex (Milner & Ewart, 1989). As systematic
regional mapping has not located other similarly equivalent intrusions elsewhere, Messum is inferred to be the
only eruptive centre for this particular suite of quartz
latites. Systematic chemical, mineralogical and isotopic
changes observed in successive eruptions of the Goboboseb and Springbok quartz latites characterize the
evolutionary development of a very large silicic magma
system, with an estimated erupted volume in excess of
8500 km3. We have termed this the Awahab magma
system. Particular interest in the Goboboseb and Springbok quartz latites lies in their very large erupted volumes,
the recognition of an eruptive centre, and trends (compositional, mineralogical, temperature) which help define
their petrogenesis.
SILICIC VOLCANIC ROCKS OF THE
´
PARANA –ETENDEKA PROVINCE
Silicic volcanic rocks of the Parana´–Etendeka Province
occur in the southeastern part of the province. They are
distributed on either side of the rifted Atlantic margin
and outcrop over an estimated 170 000 km2. Studies in
the Parana´ refer to these rocks as rhyolite or rhyodacite
(e.g. Garland et al., 1995), whereas in the Etendeka the
term quartz latite (QL) has been used. In the IUGS
TAS classification (Le Maitre, 1989), these rocks plot
continuously across the dacite–rhyolite–trachyte fields,
and their distinctive composition (e.g. high Fe and alkalis
and low Al in relation to their silica contents; see below)
make terms such as rhyolite and dacite inappropriate.
The term quartz latite, which is retained here, has been
extensively used historically in the Etendeka and is also
formally used stratigraphically.
The north–south division of the Parana´–Etendeka
Province is based largely on basalt trace element and
isotope geochemistry, with HTZ (higher Ti and Zr)
basalts characteristic of the northern province. This
regional north–south difference is also seen in the associated latitic and quartz latite volcanic rocks, which
exhibit correspondingly higher Zr and lower Zr compositions (Erlank et al., 1984; Milner et al., 1995a, 1995b,
NUMBER 2
FEBRUARY 1998
unpublished data, 1997). In the Parana´ these groups
comprise the high-Zr type Chapeco Acid Volcanics and
the lower-Zr type Palmas Acid Volcanics (PAV), which
correlate with equivalent sequences recognized in the
northern and southern Etendeka regions, respectively. In
addition to this broad regional correlation between the
Parana´ and the Etendeka, detailed stratigraphic correlations have been made between the quartz latites of
the Awahab and Tafelberg Formations in the southern
Etendeka and the PAV of the southern Parana´ (Milner
et al., 1995b). Understanding of the stratigraphy in this
area has been made possible not only by excellent exposures within the Namib Desert, but also because individual quartz latite units are sufficiently homogeneous
and coherent to allow correlations using geochemical
criteria (e.g. using Fe, Cu, Ti, P; Milner & Duncan, 1987;
Milner et al., 1992, 1995a). The level of homogeneity and
coherency exhibited by the correlated units is inferred
to result from the development of discrete silicic magma
systems and eruptive centres. However, with the exception of the Messum Complex (and associated Awahab
magma system), these centres have not been recognized
in the Parana´–Etendeka Province.
Individual units occur as extensive, flat-lying, low aspect
ratio sheets with average thicknesses ranging from 70 to
250 m. Units correlated between the Etendeka and the
Parana´ crop out over areas exceeding 25 000 km2, with
correspondingly large volumes (Milner et al., 1995a).
Large-scale silicic volcanism is now well recognized as a
distinctive class of eruptions, ranging from Archaean to
Tertiary in age, for which Henry & Wolff (1992) suggested
the term ‘flood rhyolites’.
The most recent dates ( 40Ar–39Ar) for the Awahab
quartz latite units (Renne et al., 1996) are 132·14 ± 0·40
and 131·90 ± 0·58 Ma for the Gobobseb quartz latite
Units I and II, and 132·03 ± 0·40 Ma for the Springbok
quartz latite, the ages overlapping within experimental
errors.
THE GOBOBOSEB AND SPRINGBOK
QUARTZ LATITES
Stratigraphy, distribution and correlation
The stratigraphy of the Awahab Formation in the Goboboseb Mountains region is documented in Part I. In
Namibia the Goboboseb and Springbok quartz latites
can be correlated between two distinct outcrop regions,
the Goboboseb Mountains (southernmost occurrence)
and the southern Etendeka, which lie ~60 km apart (fig.
1 of Part I). Such correlations are based on petrographic,
mineralogical, chemical and stratigraphic criteria (Milner
et al., 1992, 1995b). The lower units constitute the Goboboseb quartz latites (Units I–III), and the upper unit
comprises the Springbok quartz latite. Notwithstanding
the preserved thickness of the Springbok quartz latite
228
EWART et al.
ETENDEKA VOLCANISM, PART II
(>300 m; table 1 of Part I) nowhere in the Etendeka is
the top of this unit preserved. Moreover, this unit is
inferred to represent a single eruptive event, as are
each of the Goboboseb quartz latites (Units I–III). The
distribution of the Goboboseb and Springbok quartz
latite units in the Goboboseb Mountains, illustrated in
fig. 2 of Part I, is based on detailed field mapping
and geochemical correlations using 80 analysed samples.
Analyses of an additional 50 samples have aided detailed
correlation of these quartz latite units northward into the
southern Etendeka (fig. 1 of Part I). Most quartz latite
samples collected in the Goboboseb Mountains comprise
texturally featureless devitrified lithologies. Pitchstone
material which occurs in the upper and basal flow zones
is only developed distally from source in the southern
Etendeka, and is only poorly represented in the Goboboseb Mountains sequence. Textural and volcanological aspects have been presented by Milner et al.
(1992).
Whittingham (1991) has sampled the PAV outcropping
in the southeastern corner of the Parana´ Basin, and
recognized seven distinct units (designated PAV Units
A–G). These have been shown by Milner et al. (1995b)
to correlate in part with the Awahab Formation quartz
latites recognized within the southern Etendeka and the
Goboboseb Mountains. The most critical correlations
are those between the basal PAV Unit A with the
Goboboseb quartz latite units, and of PAV Unit B with
the extension of the uppermost part of the Springbok
quartz latite. In pre-rift reconstructions of the southeastern Parana´–Etendeka Province (Milner et al., 1995b),
outcrops of PAV Units A and B occur 340 km and
240 km from Messum, respectively, indicating very large
emplacement distances. The Goboboseb Units I and II
together with PAV Unit A encompass an area >33 000
km2 and, with an estimated average thickness of 70 m,
have an estimated volume of 2320 km3. The Springbok–
PAV Unit B quartz latite encompasses an area of 25 360
km2 and, with an average thickness of 250 m, has an
estimated volume of 6340 km3. The total estimated
volume of the erupted Awahab silicic magmas is therefore
inferred to be of the order of 8660 km3.
Mineralogy
Phenocryst phases comprise plagioclase, titanomagnetite
and pyroxenes. Microphenocrysts of apatite are ubiquitous. Modal data for selected samples show the total
phenocryst contents to range between 1·6 and 11·6% for
the Goboboseb units (most samples between 4·5 and
11%), and between 3·8 and 11% for the Springbok
quartz latite (Table 1). The Goboboseb units show no
well-defined systematic change of phenocryst abundances
with distance from source, whereas the phenocryst contents of the Springbok unit do appear to decrease away
from source. However, of more significance is the overall
uniformity of phenocryst abundance patterns between
the Goboboseb and Springbok units, and equivalent PAV
units, suggesting a well-mixed magma system(s).
Pyroxenes
Goboboseb quartz latite pyroxene phenocrysts consist of
pigeonite with traces of augite, except in Goboboseb
Unit I in the Goboboseb Mts (i.e. proximal to source),
in which augite ± pigeonite occur. The Springbok
unit contains phenocrysts of pigeonite ± hypersthene,
accompanied by groundmass and microphenocryst pigeonite. Hypersthene phenocrysts only occur in the more
distal outcrops (fig. 1 in Part I), where they are rimmed
by pigeonite. Closer to Messum, pigeonite is the only
pyroxene phenocryst phase. Based on the inference that
the upper, and more distal parts of each of the Awahab
quartz latite units represent progressively later erupted
magma, an overall pyroxene evolutionary sequence can
be recognized: augite → augite + pigeonite → pigeonite
→ pigeonite + hypersthene. Although a restricted
change of chemistry does occur within the erupted
magmas (Fig. 1), the sequence of pyroxene assemblages
is considered most plausibly to be temperature controlled.
Application of the Lindsley (1983) pyroxene geothermometer (Table 2) suggests temperatures between
1035 and 1110°C for pigeonite ± augite assemblages
and minimum temperatures of 1000–1030°C for the
hypersthene-bearing assemblages. Apatite saturation
temperatures (Harrison & Watson, 1984) are calculated
between 995 and 1025°C, with a small decrease from
the Goboboseb to Springbok units (the latter
995–1005°C). The data highlight the high magmatic
temperatures of these units [see also Milner et al. (1992)],
and suggest a small decrease of melt temperatures from
the Goboboseb through the Springbok eruptive units.
The inferred shift of pyroxene assemblage with temperature is not entirely in accord with the data of Longhi
& Bertka (1996) for Mg-rich pyroxenes, which suggest
that decreases in both T and P are necessary to produce
the observed sequence of pyroxene assemblages.
Pyroxene compositional data (Fig. 1) are based on
multiple samples from the Goboboseb, southern Etendeka
and Parana´ regions [the last data in part after Whittingham (1991)]. Within the Goboboseb units, augite
compositions in proximal samples are Wo34Fs28En37, extending to Wo34·5Fs30·5En35 in the rare augite phenocrysts
from the southern Etendeka. Pigeonite likewise exhibits
a correlated compositional shift from Wo10·5Fs42En47·5
(Goboboseb Mountains) to Wo10·5Fs45·5En44 (southern Etendeka). Three samples from correlated units in Brazil
range between Wo9·5Fs44En46·5 and Wo11Fs50En39, which
notwithstanding more compositional variability do extend
to slightly more Fe-rich compositions. The uniformity
229
230
—
0·4
4·6
Px pseudomorph
Ti-magnetite
RPhenocrysts
—
—
—
95·5
Pigeonite
Ti-magnetite
Groundmass
93·6
—
—
Plagioclase
—
6·4
0·3
1·7
—
4·4
1572
12·5
Unit II
89·2
—
—
—
10·9
1·0
3·3
—
6·6
2928
5·3
SMG059
Unit II
94·5
—
—
—
5·5
0·3
0·8
—
4·4
1583
12.5
SMG047
Unit II
95·4
—
—
—
4·7
0·6
—
—
1·1
—
3·0
3458
24
SMG030
Unit I
91·4
—
—
—
8·6
1·0
—
—
1·6
—
6·0
4487
80·5
SM-212
Unit I
92·4
—
—
—
7·6
0·6
—
<0·1
1·9
—
5·1
2705
80·5
SM-212A
Unit
A∗
94·6
—
—
—
5·4
0·8
0·9
—
—
—
3·7
3360
4·8
95·6
—
—
—
4·5
0·6
1·2
—
—
—
2·7
1804
15·5
88·8
—
—
—
11·2
1·7
—
—
2·2
—
7·3
2398
42†
96·6
—
—
—
3·4
0·4
—
—
0·7
—
2·3
2883
73†
SF1
A∗
98·4
—
—
—
1·6
0·1
—
—
0·2
—
1·3
2069
73†
SF4
A∗
93·4
—
—
—
6·6
0·7
—
—
1·8
—
4·1
2700
73†
SF6
A∗
´
Southeastern Parana Basin, Brazil
IIISMG058 IIISMG074 B024
Unit
Goboboseb Mts
A∗
88·4
—
—
—
11·6
0·9
—
0·6
2·4
—
7·7
1833
94†
GR1
A∗
92·1
—
—
—
7·9
0·7
—
—
1·6
—
5·6
3325
120†
BRA-26
NUMBER 2
Microphenocrysts
1·1
0·2
Augite
Unit I
SMG048
field
Main Etendeka lava
VOLUME 39
Pigeonite
2·9
—
Hypersthene
1922
Plagioclase
Phenocrysts
Counts:
(km):
Distance from source 5·5
Unit I
SMG060
Unit:
Sample:
Goboboseb Mts
Table 1: Modal analyses (vol. %) of the Awahab quartz latites, and their inferred correlatives in the Parana´ Basin, Brazil, in relation to
sample distance from source
(a) Goboboseb Units I and II
JOURNAL OF PETROLOGY
FEBRUARY 1998
EWART et al.
ETENDEKA VOLCANISM, PART II
Table 1: (b) Springbok quartz latite
Goboboseb Mts
Main Etendeka lava field
Brazil
SMG086
SMG087
SM-220
SM220A
SM-172
SM-168
SM-061
SM058
AT18∗
4·5
80·5
80·5
101
107·5
119
124
37†
3326
2993
3042
2619
4959
2029
1855
2669
7·6
3·9
8·5
6·3
3·7
3·0
2·6
1·0
0·5
0·9
0·3
0·5
0·4
Unit:
Sample:
Unit B
Distance from source 4
(km):
Counts:
1536
Phenocrysts
Plagioclase
7·5
Hypersthene
—
Pigeonite
—
Augite
—
—
2·2
5·5
0·8
—
0·2
—
—
—
—
—
—
2·3
—
1·3
—
—
—
—
—
—
—
—
0·8
—
Px pseudomorphs
2·8
Ti-magnetite
0·7
1·2
0·3
0·7
0·8
0·4
0·5
0·6
—
0·4
RPhenocrysts
11·0
11·0
5·2
7·5
11·6
8·0
5·2
4·9
3·8
Microphenocrysts
Plagioclase
5·1
5·4
9·3
n.d.
5·9
10·2
9·5
9·5
12·2
Pigeonite
n.d.
n.d.
5·0
n.d.
5·5
6·6
6·4
4·9
12·9
Ti-magnetite
n.d.
n.d.
1·7
n.d.
2·6
1·9
1·5
1·0
4·7
Groundmass
83·9
83·7
78·8
92·5
74·4
73·3
77·4
79·7
66·4
∗Data from Whittingham (1991).
†Distance from Brazilian coast.
Apatite present as trace microphenocrysts in all samples.
Table 2: Pyroxene geothermometer (after Lindsley, 1983; T°C, 1 bar) applied to
the Awahab quartz latites
Hypersthene
n
—
1000
3
—
—
1
Unit
Pigeonite
Augite
Springbok QL (SE)
990(M)–1030
Springbok QL (G)
1030(M)–1040(M)
Goboboseb Unit III (G)
1050(M)
—
—
1
Goboboseb Unit II (SE)
1025(M)–1050(M)
—
—
2
Goboboseb Unit II (G)
1060(M)
—
—
1
Goboboseb Unit I (SE)
1110
1060–1090
—
1
Goboboseb Unit I (G)
1035
1100–1110
—
2
Quartz monzonite plug (G)
—
850(M)
—
1
M, minimum temperature; n, number of separate rock samples used; SE, S Etendeka; G, Goboboseb Mountains.
of pigeonite compositions within individual Etendeka
samples is notable; for example, microprobe analyses of
33 separate crystals within one southern Etendeka sample
(Goboboseb Unit I; SM212) gives compositions (expressed as mol % Wo, Fs, En, respectively, with r in
parentheses) of 10·6 (0·40), 44·0 (0·45) and 45·4 (0·61).
Zoning is undetectable. Therefore, the small shift towards
Fe-enriched compositions with increasing distance from
the source is considered a significant primary magmatic
characteristic.
The Springbok quartz latite pigeonite phenocrysts
range predominantly from Wo10Fe43·5En46·5 (laccolith
phase of Messum), through Wo10Fs45·5En44·5 (Goboboseb
Mountains), to Wo9Fs45·5Fs45·5 (southern Etendeka). Variation is very slight. Pigeonite within the hypersthenebearing lithologies exhibits a greater variation of
Fe–Mg, and lower Wo, attributed to its occurrence
as microphenocrysts and reaction rims. Hypersthene
phenocrysts range between Wo3·5Fs41En55·5 and Wo3·5
Fs36En60·5.
231
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Fig. 1. Pyroxene phenocryst compositions within the Goboboseb and Springbok quartz latite units, quartz monzonite plug, and their inferred
correlatives in the Parana´, Brazil. Parana´ data mostly after Whittingham (1991). Opx, orthopyroxene; pig, pigeonite; Hbl, hornblende. In (c),
the data are distinguished between the Goboboseb Mts and the southern Etendeka main lava field localities. Numbers in parentheses refer to
sample numbers from which compositions were determined.
Garland et al. (1995) inferred that the pyroxenes within
the Parana´ Palmas rhyolites (generally equivalent to the
southern Etendeka quartz latites) were not in equilibrium
with their groundmass (liquidus) compositions, based on
comparisons with the experimental pyroxene Fe–Mg
partitioning data of Grove & Bryan (1983). However,
comparison of pyroxene K DFe–Mg values calculated for the
Awahab quartz latite units with the values for pyroxenes
in the more mafic latites from the northern Etendeka do
not provide compelling support for such disequilibrium.
Augite and pigeonite K DFe–Mg values for northern Etendeka
latites (unpublished data) are 0·2 and 0·24, respectively,
which compare with augite values of 0·23 from Grove
& Bryan (1983). Goboboseb quartz latite pyroxene K D
values are 0·19 and 0·23, respectively, whereas for pigeonite in the Springbok unit, K D is calculated to be
0·28. This latter value is higher, but this may correlate
with the absence of coexisting augite.
Plagioclase
Phenocrysts (excluding rims) range between An53 and
An63 (mol), with Or solid solution between 2·4 and 3·4%,
and the most calcic compositions occurring in the cores
of the largest phenocrysts. Phenocryst rims are in the
range An47–59, and microphenocrysts between An43 and
An51. Histograms of phenocryst compositions show dominant compositions of approximately An54–An59 for the
Springbok quartz latite and An53–An60 for the Goboboseb
quartz latite units, with FeO∗ (total Fe as FeO) concentrations ranging between 0·7 and 1·1% (Fig. 2). In
the Goboboseb quartz latites, plagioclase shows a subtle
increase in the frequency of more calcic compositions
and a slightly higher median FeO∗ value.
Three petrographic features of plagioclase are notable;
the lack of breaking and fragmentation (Milner et al.,
1992); the common occurrence of quench textures within
the groundmass microlites; and the common occurrence
of coarse sieve textures within phenocrysts. Some sievetextured phenocrysts are overgrown by thin continuous,
inclusion-free euhedral plagioclase rims, whereas in other
crystals such overgrowths are absent (Fig. 3b and c). The
compositions within the clear (i.e. inclusion-free) areas
lie within the normal range (i.e. An51–An58). However,
the optical appearance of plagioclase adjacent to the
glass inclusions (or ‘blebs’) becomes more turbid, and
microprobe analyses (Table 3) reveal this material to be
poorly stoichiometric, with compositions between An50
and An54, but abnormally high FeO∗ (0·9–1·3%) and Or
(5·3–6·8%).
232
EWART et al.
ETENDEKA VOLCANISM, PART II
Fig. 2. Histograms of % An (mol) and FeO∗ compositions of plagioclase phenocrysts within the Goboboseb and Springbok quartz latite units.
Table 3: Microprobe analyses of plagioclase within sieved phenocrysts,
Goboboseb quartz latite (sample SM212)
Average of clear plagioclase between
Average of turbid plagioclase adjacent
inclusions (n=14)
to inclusions (n=4)
SiO2
55·03 (0·56)
57·31 (1·53)
Al2O3
27·16 (0·43)
25·34 (0·76)
FeO∗
0·83 (0·06)
1·21 (0·12)
MgO
0·09 (0·01)
0·13 (0·01)
CaO
11·18 (0·38)
10·30 (0·55)
Na2O
4·77 (0·06)
4·61 (0·21)
K 2O
0·48 (0·06)
R
0·88 (0·21)
99·54
99·78
Structural formulae (O = 8)
Si
2·504
Al
1·457
2·595
1·352
Fe2+
0·032
0·046
Mg
0·006
0·009
Ca
0·545
0·500
Na
0·421
0·405
K
0·028
0·051
R cations
4·993
4·958
An (mol %)
54·8
52·3
Ab
42·4
42·4
Or
2·8
5·3
Numbers in parentheses are standard deviations.
233
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Fig. 3. (a) Photomicrograph of acicular quartz (white) in matrix of quartz monzonite plug. Also present are opaque magnetite, hornblende and
turbid feldspar. Field of view 1·3 mm. (b) and (c), photomicrographs of sieved textures within plagioclase phenocrysts of the quartz latites. Field
of view 3·3 mm.
234
EWART et al.
ETENDEKA VOLCANISM, PART II
Table 4: Microprobe determined compositional ranges of inclusions within sieved
plagioclase phenocrysts; Goboboseb quartz latite (sample SM212) (values in weight per
cent)
SiO2
TiO2
Al2O3
FeO∗
MnO
Type I
Type II
Type III
Low-K–high-Si type
High-K–high-Si type
High-Fe type
72·6–76·8
71·3–75·4
66·3–66·4
0·25–0·61
10·4–12·5
0·8–5·6
n.d.
0·19–0·96
0·5–1·8
9·4–12·7
6·0–11·1
0·6–4·3
19·0–24·2
n.d.
0·15–0·28
MgO
0·1–0·89
0·1–0·2
0·57–1·55
CaO
0·31–2·5
0·14–3·3
1·64–2·73
Na2O
1·4–4·3
0·2–1·9
0·3–1·6
K 2O
1·4–4·6
6·1–10·4
0·5–1·1
F
n.d.
(2·8)1
n.d.
n.d., not detected.
1
Only detected in one inclusion.
Inclusions within the plagioclase phenocrysts are optically and compositionally very variable. Most appear
glassy, some proving extremely unstable under an electron probe beam. Compositions fall broadly into three
groupings as documented in Table 4. The only explanation offered to account for the diverse inclusion
chemistry is that they represent disequilibrium partial
melt droplets, generated at the marginal zones of the
magma body, which because of sluggish mixing with the
main magma mass, either act as nuclei for plagioclase
precipitation, or become incorporated into rapidly growing plagioclase crystals. The poorly stoichiometric plagioclase may therefore represent the initial nucleation of
a plagioclase phase onto the surfaces of melt ‘droplets’.
If the above interpretation is valid, then the compositional
diversity of the inclusions will reflect the mineralogy of
the immediately surrounding country rocks undergoing
incongruent melting. Thus, the high-Fe inclusions may
represent original breakdown of magnetite-bearing assemblages, the high-K types the breakdown of biotitebearing assemblages, and the low-K and high-Si types
the breakdown of feldspar–quartz dominated assemblages. One implication of this interpretation is that
active melting was occurring at the magma chamber
walls, without the development of a chilled envelope or
zone of side-wall crystallization to act as a barrier to
magma–country rock interaction.
Titanomagnetite
Magnetite forms ubiquitous subhedral–euhedral crystals,
0·1–0·5 mm in diameter, which occur as isolated phenocrysts or inclusions in pyroxenes. Primary (unexsolved)
compositions are best preserved in pitchstones, and are
only found distally from source. Compositional ranges
(%) for the Goboboseb units (four samples) and the
Springbok unit (two samples) are: TiO2 18·4–20·1 and
17·6–21·2; Al2O3 2·4–2·7 and 2·0–2·3; MgO 0·71–1·4
and 0·85–1·4; and MnO 0·33–0·48 and 0·43–0·65, respectively. Apart from small shifts towards higher MnO
and lower Al2O3 in the Springbok magnetites, no significant differences are apparent.
Quartz monzonite plugs
These are interpreted as the equivalent intrusive facies
of Goboboseb quartz latite Units I and II (Milner et al.,
1992). The plugs consist of phenocrysts of labradorite
(An50–An55; up to 5 modal %) and magnetite (<1%), set
in a coarse-grained groundmass of intergrown alkali
feldspar and quartz, with minor magnetite, apatite, scarce
brown hornblende (Fig. 1), and plagioclase microphenocrysts. Pseudomorphs of chlorite ± epidote after
pyroxene phenocrysts are present, with unaltered augite
only rarely preserved.
Three features of the mineralogy are significant:
(1) Phenocryst mineralogy is comparable with that of
the quartz latites.
(2) Quartz morphology is unusual, consisting of acicular
crystals, typically intergrown into lattice-like or skeletal
aggregates (Fig. 3a) which locally coalesce into interstitial
quartz patches. The quartz and intergrown alkali feldspars maintain optical continuity over phenocryst-sized
areas. These textures are consistent with high-temperature quenching, and the form of the quartz suggests
235
JOURNAL OF PETROLOGY
VOLUME 39
that it has inverted from b-tridymite, although microprobe analyses reveal only pure SiO2.
(3) Microprobe analyses of the hornblendes reveal
them to be fluor-hornblendes with intermediate Fe–Mg
ratios and low Cl contents. Although the present data
give no direct estimates of F melt fugacities, they do suggest
that F was a significant volatile component in what are
believed to have been strongly volatile-undersaturated
melts. No amphibole or biotite have been identified in
any extrusive quartz latite unit of the Etendeka.
Quartz latite chemistry
On a regional scale the Awahab quartz latite magmas
(Tables 5 and 6) are relatively homogeneous chemically
(and mineralogically), with SiO2 ranging between 67
and 69% (Etendeka), extending to 71% when the PAV
equivalents are included. It is the very restricted chemical
variation, especially within the individual quartz latite
units, that allows geochemistry to be successfully used
for correlation purposes. In detail, however, systematic
variations do occur stratigraphically from the lowest
Goboboseb quartz latite (Unit I) through the Springbok
quartz latite unit, to the Parana´ PAV Unit B, which are
best exemplified by FeO∗ and SiO2. The Springbok
quartz latite and PAV Unit B represent the relatively
SiO2-rich and Fe-depleted end of the trend.
Figure 4 illustrates representative plots for SiO2, TiO2,
Rb, Zr and V vs FeO∗, the last providing both a consistent
differentiation index for the quartz latite eruptives, and
a working chemical division between the Goboboseb and
Springbok units within the Etendeka (although noting
the overlap when the PAV data are included). P2O5,
TiO2, Y, Zr, Nb and Cu are positively correlated with
FeO∗, with similar but less well-defined behaviour
observed for light rare earth elements (LREE), Zn, Sc
and Mn. U, Pb, Ba and Rb are negatively correlated
with FeO∗ [noting the differences between devitrified
and glassy lithologies; see Table 6 and Milner & Duncan
(1987)]. V exhibits different, sub-parallel abundance patterns between the two main QL units (Fig. 4e), although
generally correlating with increasing FeO∗. These abundance patterns result in consistent element ratio changes,
notably Th/U, Ti/V, Ce/Pb, K/Rb, K/U and P/Ce
(all decreasing), and Rb/Ba, Ce/Yb, Rb/Zr, Zr/Y, Sr/
Nd and Rb/Sr (all increasing) from the Goboboseb
through to the Springbok quartz latite units, and extended
by PAV unit B. Nb/Ta, Zr/Hf and Ce/Pb ratios are in
the range of 10, 41–44 and 3–5, respectively. Figure 5a
compares a range of element abundances in selected
samples between the Goboboseb and Springbok units
(Etendeka only), normalized to estimated mean upper
crust (Taylor & McLennan, 1985). Apart from emphasizing the overall similarities of the Awahab quartz
NUMBER 2
FEBRUARY 1998
latites, the spidergram shows a general correspondence
to the inferred upper-crustal composition, with the heavy
rare earth elements (HREE), Ti and Y becoming relatively more enriched relative to the alkalis and LREE.
The most conspicuous feature, however, is the strong
negative Sr anomaly [also shown if the data are normalized to the lower crust and bulk continental crust
estimates of Taylor & McLennan (1985)].
Chondrite-normalized REE patterns (Fig. 5b) are similar for both the Goboboseb and Springbok quartz latite
units, showing LREE enrichment with La/Yb ratios
varying from 10·2 to 10·9 for the Goboboseb units,
increasing to 11·2–11·4 for the Springbok unit. Negative
Eu anomalies are present, with Eu/Eu∗ ratios lying
between 0·62 and 0·67, there being no systematic differences between the two units. The Eu and Sr anomalies
together suggest significant feldspar fractionation accompanying the evolution of the QL magmas.
ISOTOPE GEOCHEMISTRY
The Springbok and Goboboseb quartz latites define the
most ‘evolved’ or ‘crustal’ end of the Sr–Nd–Pb isotope
data arrays for the Goboboseb Mountains succession (figs
12–15 in Part I; Table 7), having relatively radiogenic
Pb and Sr compositions which overlap those of the more
evolved LTZ.L lavas. In Pb isotopic space, the quartz
latite compositions plot beyond the termination of the
geochron, consistent either with an older, high U/Pb
and Th/U source, or a source that has been mixed or
rejuvenated by mantle-derived magmas (see Rudnick &
Goldstein, 1990). The quartz latite melts are, nevertheless,
more radiogenic than model compositions calculated for
average lower crust by the latter workers, and are closer
to their total crustal model compositions.
Initial Sr and Pb isotopic compositions of the Goboboseb and Springbok quartz latites correlate inversely
with FeO∗, with the initial Sr isotope ratios of PAV Unit
B consistent with, and extending the Etendeka trends
(Fig. 6). These isotopic shifts suggest that the major and
trace element variations described above are not simply
the products of melting or fractional crystallization, but
must reflect mixing or AFC processes.
Although no new O-isotope data are presented here,
we summarize the available data. Harris et al. (1990)
reported plagioclase and pyroxene phenocryst d18O values of 10·1–10·9‰ and 9·9–10·0‰ for the Springbok
quartz latite and PAV Unit A, respectively, and a d18O
value of 9·5‰ has been recorded for pyroxene from the
Goboboseb quartz latite Unit I by Harris (1995). Although
the data are few they do show a negative correlation
with FeO∗. Assuming mineral–melt fractionation factors
of –0·2 and –0·7 for plagioclase and pyroxene, respectively, d18O magma values for the Awahab quartz
236
EWART et al.
ETENDEKA VOLCANISM, PART II
Table 5: Major and trace element data for selected quartz latite samples from the Goboboseb
Mountains, the main Etendeka lava field, and Messum Complex core breccias; all data recalculated
on anhydrous basis, with RFe as FeO
Goboboseb Mountains
Main Etendeka lava field
Goboboseb quartz latites
Unit:
Unit I
Unit II
Unit III
Unit I
Unit II
Unit II
Sample:
SMG060
SMG059
SMG058
SM212
KLS036∗
SM215
68·18
SiO2 (wt %)
66·99
67·07
67·62
67·87
67·78
TiO2
1·08
1·06
1·01
1·04
1·01
1·02
Al2O3
12·81
12·84
12·87
13·07
12·87
12·98
FeO∗
7·11
7·23
6·85
6·90
6·78
6·71
MnO
0·12
0·12
0·13
0·12
0·12
0·12
MgO
1·13
1·10
1·09
1·00
1·11
1·03
CaO
3·38
3·38
3·00
3·39
3·28
3·36
Na2O
3·01
2·96
3·06
4·02
2·72
4·35
K 2O
4·05
3·94
4·08
2·28
4·03
1·95
P 2 O5
0·32
0·30
0·30
0·32
0·31
0·31
Trace elements (ppm)
Rb
160
152
161
166
140
160
Ba
706
688
698
735
680
739
Sr
159
155
142
156
165
156
Th
13·7
13·7
13·9
13·5
15·9
U
3·6
3·6
4·1
3·8
4·3
Zr
315
303
302
301
295
14·9
—
298
Hf
7·2
7·1
7·0
7·1
7·0
Nb
23·3
22·4
22·7
25·4
21·8
22·9
—
—
Ta
2·1
2·15
2·2
Cr
4·5
3·2
4·7
V
47
59
Sc
21
Ni
n.d.
Co
2·6
—
4·0
—
1·9
45
40
40
42
21
20
18
19
18
2
n.d.
—
—
—
15
15
13
15
12
14
Pb
24
23
22
27
29
25
Zn
93
79
96
91
84
93
Cu
22
23
18
20
17
20
Y
51·6
49·9
50·1
48·8
46·8
44·2
La
47·6
46·2
46·2
44·7
40·4
47
Ce
101
100
98·9
98·7
Pr
—
—
—
—
12·2
—
Nd
48·1
47·8
46·7
47·3
47·1
57
Sm
9·1
9·0
8·7
8·9
9·0
—
Eu
2·0
1·9
1·9
1·9
1·7
—
Gd
8·5
8·5
8·45
8·8
8·3
—
Tb
Dy
Ho
1·5
—
1·7
1·6
—
1·8
1·5
—
1·85
1·5
—
1·9
Er
—
—
—
—
Tm
—
—
—
—
Yb
4·5
4·4
4·2
4·4
Lu
0·63
0·62
0·61
0·61
237
104
99
1·3
—
7·6
—
1·6
—
4·2
—
0·55
—
4·0
—
—
—
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Table 5: continued
Goboboseb Mountains
Main Etendeka lava field
Messum Complex
Springbok quartz latites
Sample:
Core breccias
SMG087
SMG098
KLS051∗
SiO2 (wt %)
SM168
SM230
SMG094D
SMG094F
68·17
67·89
68·90
68·39
68·04
79·36
TiO2
0·95
0·93
0·96
0·93
0·97
0·22
63·45
0·62
Al2O3
13·04
13·07
13·11
13·36
13·38
11·67
16·75
FeO∗
6·82
6·26
5·97
6·11
6·37
2·41
6·35
MnO
0·12
0·11
0·11
0·10
0·12
0·03
0·10
MgO
0·92
1·27
1·12
0·88
1·57
0·23
0·57
CaO
3·03
2·90
3·27
3·20
2·63
0·38
1·85
Na2O
2·70
3·03
3·59
3·14
2·67
4·13
5·67
K 2O
3·96
4·26
2·69
3·60
4·16
4·56
4·52
P 2 O5
0·29
0·28
0·28
0·28
0·28
0·01
0·13
Trace elements (ppm)
Rb
157
171
165
176
167
207
152
Ba
688
677
690
673
682
322
1381
Sr
153
156
162
167
173
35·7
Th
14·2
14·1
14·1
14·0
13·7
22·6
U
4·4
4·9
4·7
4·5
4·0
4·3
Zr
292
275
275
284
Hf
6·8
6·6
6·7
6·4
Nb
21·5
21·2
21·8
23·9
Ta
2·2
2·3
Cr
6·6
6·5
—
2·5
5·3
3·9
273
—
15·6
24·1
79
—
2·1
V
45
53
43
49
64
Sc
19
18
19
17
20
Ni
5·9
2·8
—
2·2
602
3·0
5·9
159
17·3
3·7
804
16·3
152
7·7
4·3
3·7
3
3
1·0
9
1·2
—
—
Co
14·2
14
10
13
14
1·0
Pb
23
24
30
22
23
8·2
Zn
83
87
81
83
82
Cu
21
8
16
17
17
Y
47·5
45·1
42·7
45·3
43·1
La
45·5
45·5
40·2
44·7
45
Ce
95·6
97·5
95·6
94·5
94
38
8·6
72
7
12
130
104
81·5
167
76·5
161
Pr
—
—
11·0
—
—
—
—
Nd
44·1
45·9
41·9
44·7
47
87·5
75·5
14·5
Sm
8·2
8·4
8·8
8·3
—
17·9
Eu
1·8
1·8
1·6
1·7
—
1·4
3·6
Gd
8·0
7·9
7·3
8·3
—
19·1
15·7
1·4
1·2
1·4
—
Tb
Dy
Ho
1·5
—
1·8
—
1·6
Er
—
—
Tm
—
—
Yb
4·1
4·0
Lu
0·57
0·5
6·7
—
1·4
1·6
—
—
3·7
—
—
0·49
—
3·5
—
3·2
—
4·3
—
2·6
—
3·4
—
—
—
—
4·0
—
11·2
10·4
0·56
—
1·5
1·5
Major elements, Rb, Ba, Sr, Zr, Nb, Cr, V, Sc, Ni, Co, Zn, Cu, Y, and Pb (in part) by XRF (Department of Earth Sciences,
University of Cape Town). Remaining trace elements either by INAA (B. W. Chappell, Australian National University),
following Chappell & Hergt (1989), or those marked by asterisk, after Duncan et al. (1984) based on spark-source mass
spectrography.
238
4·34
0·32
K 2O
P 2O5
239
25·9
84·0
18·8
48·8
46·9
99·3
54·7
32
Pb
Zn
Cu
Y
La
Ce
Nd
n
33
54·0
99·2
47·2
47·6
18·2
94·5
27·4
14·8
4·1
50·0
20·6
23·6
303
6·0
15·2
219
771
171
0·32
3·20
1·06
0·12
6·87
12·94
44
50·4
94·0
47·3
44·6
17·2
81·2
25·2
13·1
5·1
50·0
19·2
23·0
279
5·4
15·9
155
696
185
0·29
3·88
2·93
2·88
1·18
0·10
6·26
13·22
0·95
68·31
QL
Springbok
All data
8
55
98
47
48
17·9
84
24
15·4
5·1
46
21
25
299
6·0
13·8
149
712
175
0·31
4·56
2·61
2·67
1·01
0·10
6·81
13·03
1·03
67·86
5
52
101
46
48
18·7
91
26
14·1
4·0
43
18·7
24
302
5·2
14·4
164
726
186
0·31
2·53
3·71
3·50
1·01
0·12
6·83
13·00
1·03
67·96
25
50
93
47
44
17·9
81
25
13·3
5·3
48
19·8
23
277
5·6
15·7
145
678
174
0·29
4·47
2·64
2·63
1·27
0·11
6·30
13·11
0·95
68·24
Devitrified
Devitrified
Pitchstone
Springbok QL
Goboboseb QL units
Southern Etendeka only
∗Taylor & McLennan (1985). †Chappell & White (1992). ‡Kilpatrick & Ellis (1992).
6·7
15·4
44·7
V
Co
20·2
Sc
Cr
26·2
Nb
300
3·8
U
Zr
14·3
Th
169
2·72
Na2O
Sr
2·62
CaO
758
1·13
MgO
184
0·10
MnO
Ba
3·87
6·83
FeO∗
Rb
2·93
12·97
Al2O3
1·05
1·04
67·64
67·93
TiO2
Unit II
SiO2
Unit I
All data Goboboseb units
12
50
95
46
44
17·0
84
25
12·6
5·2
53
17·7
23
282
5·1
15·9
182
701
202
0·29
2·45
3·45
3·57
1·04
0·11
6·22
13·28
0·95
68·65
Pitchstone
—
—
64
30
22
25
71
20
10
35
60
11
25
190
2·8
10·7
350
550
112
—
3·4
3·9
4·2
2·2
0·08
4·5
15·2
0·5
66·0
crust∗
continental
upper
Average
Table 6: Averaged compositions of Goboboseb and Springbok quartz latites and comparative data
—
—
33
16
20
75
80
8
29
185
230
30
11
100
0·91
3·5
260
250
32
—
1·1
3·1
7·4
5·3
0·18
9·1
15·9
0·9
57·3
crust∗
continental
Bulk
1074
—
66
31
31
9
48
19
10
20
57
13
11
150
5
20
235
519
164
—
3·48
3·16
3·07
1·38
0·07
3·12
14·21
0·41
69·50
Australia)†
granite (SE
Belt I-type
Lachlan Fold
Average
—
—
95
41
37
19
75
24
—
11
90
—
10
305
1
4·5
186
1140
153
0·35
4·08
2·51
3·67
1·44
0·09
6·47
13·85
1·16
66·37
Charnockite‡
Ardery
EWART et al.
ETENDEKA VOLCANISM, PART II
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Fig. 4. Selected variation diagrams for the Goboboseb and Springbok quartz latite units, and their inferred Parana´ correlatives.
latites are estimated to range between 10·2 and 11·6‰,
with an average value of 10·8‰ (including PAV correlatives; Harris, 1995).
The plot of Sm/Nd–Rb/Sr ratios (Fig. 7a) shows these
to be negatively correlated in the quartz latite units and
LTZ.L basalts, and is therefore similar to the equivalent
eNd vs eSr plot (fig. 12, Part I). This highlights the general
correspondence between the Rb–Sr and Sm/Nd isotope
and parent/daughter trace element ratios; a relationship
which was noted by Hawkesworth et al. (1984) for the
240
regional Etendeka basaltic lavas, and interpreted as providing evidence for source inherited geochemical and
isotopic characteristics from ‘enriched’ subcontinental
lithospheric mantle (SCLM). Although Hawkesworth et
al. (1984) suggested the correlations were unlikely to
reflect crustal contamination processes, the two mixing
curves shown in Fig. 7, between Springbok QL and EMORB, and Springbok QL and Lo.L. crust end-members (see Part I) suggest that crustal mixing could produce
the observed parent/daughter inverse ratio correlations,
EWART et al.
ETENDEKA VOLCANISM, PART II
Fig. 5. (a) Upper crust normalized (Taylor & McLennan, 1985) spidergram, and (b) chondrite-normalized (Sun & McDonough, 1989) REE
plots for the Goboboseb and Springbok quartz latite units. REE plot based on instrumental neutron activation analysis (INAA) data only.
and as shown in fig. 17 of Part I, can also, in principle,
explain the eNd–eSr isotopic arrays.
Isotopic data for rhyolite and latite clasts from the
Messum core breccias (fig. 12, Part I) show these to be
different from the LTZ.L and quartz latite magmas in
terms of eNd and initial 87Sr/86Sr. Although eNd exhibits
closer isotopic affinities to the LTZ.H basalts, the higher
initial 87Sr/86Sr of the rhyolite is distinctive. Their Pb
isotope compositions are more radiogenic than the
LTZ.H lavas, and for the rhyolite, more radiogenic than
the quartz latite units. The data indicate a petrogenesis
for these early volcanic phases within the Messum Complex that is independent of the quartz latite and basaltic
LTZ.L magmas.
PETROGENESIS
Introduction
Milner (1988) advocated a model for the southern Etendeka quartz latite magmas involving partial melting,
within mid- to lower-crustal levels (30–35 km), of a
241
87
Sr/86Sr
0·725288±10
SMG058 (Unit III)
0·72320±8
0·725020±10
KLS036 (Unit II)∗
SM215 (Unit II)
242
0·72563±8
0·727018±10
0·725704±10
0·723847±10
KLS034∗
SM168
SM222F
SM230
165
167
172
176
167
165
0·723708±10
0·711318±10
SMG094F
152
206·8
161
159
35·7
158
173
163
167
139
153
162
156
153
156
165
156
142
0·717449
0·706013
0·708843
0·718043
0·718483
0·719855
0·721137
0·718968
0·719276
0·720877
0·720888
0·719918
0·719307
0·718479
0·718810
0·718977
0·718309
183·1
20·7
60·9
191·5
197·7
217·1
235·4
204·6
209·0
231·7
231·9
218·1
209·5
197·7
202·4
204·8
195·3
Nd/144Nd
0·512655±10
0·512555±10
0·512085±10
0·512133±10
—
0·512108±10
—
—
0·512190±10
0·512095±10
0·512115±10
0·512189±10
0·512190±20
0·512181±10
0·512089±10
0·512125±10
0·512115±10
143
14·5
17·9
9·16
—
—
8·31
—
—
8·83
8·43
8·22
—
8·98
8·93
8·65
9·00
9·09
ppm Sm
75·5
87·5
48·6
—
—
44·7
—
—
41·9
45·9
44·1
—
47·1
47·3
46·9
47·8
48·1
ppm Nd
0·512553
0·512446
0·511984
—
—
0·512009
—
—
0·512077
0·511997
0·512015
—
0·512088
0·512080
0·511990
0·512024
0·512014
1·83
−0·26
−9·26
19·139±5
19·825±5
19·003±5
—
19·137±5
—
19·189±5
−8·79
−8·16
—
—
19·180±18
—
—
−7·44
19·080±5
19·116±5
—
−7·3
−8·65
19·000±18
−9·01
—
−7·23
19·017±5
−9·14
−7·39
18·975±5
18·969±5
Pb/204Pb
−8·69
206
−8·48
( 143Nd/144Nd) eNd
Analyst: R. A. Armstrong, except those samples marked by asterisk, which are from Bristow et al. (1984) and Hawkesworth et al. (1984).
0·740388±10
SMG094D
Messum Complex—core breccias
SMG090B
159
155
eSr
Pb/204Pb
15·651±5
15·682±5
15·685±5
15·685±5
—
15·736±5
—
—
15·770±15
15·658±6
15·694±5
—
15·750±15
—
15·687±5
15·671±5
15·688±5
207
Pb/204Pb
39·180±5
39·564±5
38·908±5
38·852±5
—
39·033±5
—
—
39·080±40
38·763±5
38·914±5
—
39·040±40
—
38·892±5
38·859±5
38·938±5
208
NUMBER 2
Quartz monzonite plug—Goboboseb Mts
0·726530±8
0·72528±7
KLS051∗
171
157
160
140
166
161
152
159
ppm Rb ppm Sr ( 87Sr/86Sr)0
VOLUME 39
KLS033∗
Main Etendeka lava field
0·725628±20
0·726992±10
SMG087
SMG098
Goboboseb Mts
Springbok quartz latite
0·724750±10
SM212 (Unit I)
Main Etendeka lava field
0·723028±20
0·723778±10
SMG060 (Unit 1)
SMG059 (Unit II)
Goboboseb Mts
Goboboseb quartz latites
Sample
Table 7: Sr, Nd, and Pb isotopic analyses of Goboboseb quartz latite, and Messum core volcanic breccia phases; initial ratios and epsilon
values corrected to 132 Ma
JOURNAL OF PETROLOGY
FEBRUARY 1998
EWART et al.
ETENDEKA VOLCANISM, PART II
Fig. 6. Initial Sr isotopic and 206Pb/204Pb compositions, vs FeO∗, for the Goboboseb and Springbok quartz latite units, and their inferred
correlatives from the Parana´, Brazil (Sr only; after Whittingham, 1991).
source of mafic to intermediate composition, the latter
approximating the average bulk crustal composition
(Taylor & McLennan, 1985). Initiation of melting was
inferred to result from underplating of basaltic magma
within the lower crust, with the crustal source identified
as plausibly the ~2·0 Ga pre-Damara basement. Harris
et al. (1990) and Harris (1995) supported this conclusion,
on the basis of O and Sr isotope data, and suggested
that the southern Etendeka quartz latites could have
been derived from the same source as the Damara
granites, or from the restite remaining after Damara
granite genesis.
For the equivalent PAV of the Parana´, Bellieni et al.
(1984) proposed a derivation from basaltic precursors
through low-pressure crystal fractionation processes accompanied by crustal contamination. Bellieni et al. (1986)
and Piccirillo et al. (1988, p. 199) subsequently suggested
two alternatives. The first is the melting of lower-crustal
mafic to intermediate granulites (initial Sr isotopic
composition ~0·714), with subsequent evolution of the
243
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 2
FEBRUARY 1998
Fig. 7. Sm/Nd vs Rb/Sr ratios of: (a) the Awahab mafic and quartz latite volcanic phases, and Messum core breccias, with two mixing curves
based on end-members labelled; (b) the generalized compositional fields of the Awahab volcanic phases [from (a) above; 1, LTZ.H basalts; 2,
LTZ.L mafic lavas; 3, quartz latites], compared with various model crustal compositions (see Part 1; data after Taylor & McLennan, 1985;
Zartman & Haines, 1988; Rudnick & Presper, 1990); SCLM (McDonough, 1990); averaged MORB composition (Sun & McDonough, 1989);
Parana´ granitoid and basement gneiss phases (May, 1990); generalized fields of granulites, charnockites and khondalites (after Hubbard &
Whitley, 1979; Field et al., 1980; Weaver, 1980; Demaiffe & Hertoogen, 1981; McCulloch & Black, 1984; Sheraton et al., 1985; Condie et al.,
1992); A-type granitoids (after Collins et al., 1982; Eby, 1990); and felsic I-type granites (Champion & Chappell, 1992).
Palmas silicic melts occurring through AFC processes
involving radiogenic crust and relatively low values of
assimilation/crystallization ratios (r ~0·2). The second
alternative proposed 10–20% melting of an underplated
mafic source corresponding to the associated low- and
high-Ti basalts. These workers noted that major and
trace element modelling alone is consistent with fractional
crystallization of the PAV from low-Ti basalts. Conversely, for the same Parana´ silicic units, Whittingham
(1991) advocated a high degree of crustal partial melting
of a source rock with a composition similar to that of
average upper crust. In the most recent work, Garland
et al. (1995) concluded that the PAV are derived predominantly from a low-Ti basaltic parent, rather than
crustal basement, with a genetic link via open system
fractional crystallization, and further inferred that they
evolved in stable shallow-level magma chambers. Clearly,
a significant divergence of opinion exists as to the origin
244
EWART et al.
ETENDEKA VOLCANISM, PART II
of the PAV (and by correlation, the Etendeka Awahab
quartz latite units).
General aspects
The following petrological aspects are relevant to quartz
latite melt petrogenesis:
(1) The very large estimated volume of the Awahab
magma system. Put in the perspective of crustal thickness,
a magma sphere of an equivalent volume to the Springbok
unit alone (>6500 km2) would have a diameter of 23 km,
increasing to 25·5 km for both the Goboboseb and
Springbok quartz latite volumes. These volumes indicate
minimum estimates of magma chamber dimensions, as
the chambers are unlikely to have been totally evacuated
(e.g. Smith, 1979).
(2) Systematic chemical variations within the Awahab
quartz latites, shown for example by SiO2, which varies
between 67 and 71%, and FeO∗, which varies from 5·3
to 7·5% (including the Parana´ correlative units). With
progressive eruption of the system, the melts become
more geochemically evolved (i.e. higher SiO2, lower Fe),
with the most evolved compositions occurring in PAV
Unit B. The Springbok–PAV Unit B quartz latite is
inferred to represent the latest erupted magma, and
significantly, is characterized by the most radiogenic Sr
and Pb isotopic compositions. The bulk chemical changes
are correlated with systematic changes of pyroxene phenocryst assemblages and compositions. Notwithstanding
these variations, the magma system apparently was not
strongly stratified chemically, suggesting strong convective mixing.
(3) High crystallization temperatures (1000–1100°C)
are inferred from pyroxene geothermometry, consistent
with the absence of any hydrous mineral phases. Application of the ‘MELTS’ magma modelling procedure
(Ghiorso et al., 1994) gives anhydrous liquidus temperatures (°C) for the averaged Goboboseb Unit I and
Springbok quartz latite compositions of 1084 and 1064
(1 bar); 1107 and 1089 (2 kbar); and 1155 and 1149 (5
kbar), respectively. The two lower-pressure estimates are
consistent with pyroxene thermometry, and suggest a
small temperature drop during the eruption of the Awahab quartz latite system, also consistent with apatite
thermometry.
(4) The correlation of bulk chemical changes with the
systematic changes of pyroxene phenocryst compositions,
and the inferred high crystallization temperatures imply
low volatile contents. If this is correct, then the relative
constancy of phenocryst abundances suggests that the
vertical magmatic temperature gradient approximately
paralleled the phenocryst saturation curves, which will
have positive dT/dP slopes under these conditions. Based
on relevant experimental data for granitic melts (e.g.
Brown, 1970; Thompson, 1988; Johannes & Holtz, 1991),
inferred temperature gradients are of the order of
3·5–4·5°C/km. Although the pyroxenes appear to be in
equilibrium with the magmas, the compositions of the
plagioclase phenocrysts (An50–An60) are more problematic. In terms of the relatively high normative An
(~10%) of the quartz latite melts, the data of Nekvasil
(1988) suggest that the plagioclase compositions are more
consistent with water-unsaturated crystallization. Nevertheless, their relatively calcic compositions are not readily
reconciled with the melt chemistry; for example, application of the MELTS modelling procedures (Ghiorso
et al., 1994) has failed to reproduce observed compositions.
The abundance of melt inclusions, and their wide range
of compositions, further suggests a degree of plagioclase
disequilibrium.
(5) The relatively high average magma d18O value of
the Awahab quartz latites (10·8‰) is consistent with a
mid- to upper-crustal basement source input. Basement
granites of Damara age have d18O values of 6–15‰,
averaging 11·4‰, whereas Damara schists of the Kuiseb
Formation (T NdChur ~1 Ga) and the Zerrissene group
have d18O values of 11·6–14·7‰ (average 13·2‰) and
11·2–19·1‰ (average 14·3‰), respectively (Haack et al.,
1982; Harris, 1994). If the Awahab quartz latites were
derived by pure assimilation (simple mixing) of these
basement lithologies by a basaltic precursor with a mantle
d18O value (5·7‰), the amount of assimilation would be
in the range 60–90%, using average values for the
Zerrissene schists and the Damara granites, respectively.
(6) Although the large volumes and siliceous compositions of the Awahab quartz latites imply a major
crustal input into the erupted magmas, the quartz latite
compositions are atypical of silicic melts commonly interpreted as originating by crustal anatexis. For example,
comparison with the averaged SE Australian I-type granites (Chappell & White, 1992; table 3) reveals the latter
to be higher in SiO2, Al2O3, Sr, Zn, Cr and Th, but
markedly lower in high field strength elements (HFSE),
LREE, Ba, Sc, Cu, P and especially Fe. S-type granites
are even more divergent. More significantly, experimental
studies do not reproduce compositions comparable with
the Awahab QL melts: for example, the experiments of
Beard & Lofgren (1991) on H2O-saturated and
-undersaturated melting of metamorphosed basalts and
andesites; Beard et al. (1994) on the melting of felsic
charnockites, dioritic gneiss and felsic garnet granulite;
and Rushmer (1991) on partial melting of various amphibolites. Fluid absent melting in pelitic systems (Vielzeuf
& Holloway, 1988) can produce high-degree melts
(>1100°C) of intermediate–silicic compositions, with high
Fe, Ti and alkalis, but also very high Al2O3, the latter
not characteristic of the Awahab quartz latite melts. In
these, as in other reported melting experiments especially
245
JOURNAL OF PETROLOGY
VOLUME 39
from more silicic parent compositions, none of the melts
are similar to the Awahab quartz latite compositions.
(7) An alternative petrogenesis involves total crustal
fusion above the focus of the Tristan da Cunha plume.
The bulk compositions of the Awahab quartz latites are
broadly similar to the estimated average upper crust,
although they differ from the estimated bulk continental
crustal composition (Table 3). Data from Clemens &
Vielzeuf (1987) and Rushmer (1991) show that for fluid
absent fusion melt proportions depend mainly on source
composition and higher temperatures will increase melt
production from a given source. However, for a low H2O
pelitic and quartzo-feldspathic source, near-complete fusion will require temperatures in excess of 1100°C (5–10
kbar range), and even higher temperatures are required
for intermediate and mafic sources (Rushmer, 1991).
(8) Mass balance modelling confirms that fractionation
of the Awahab quartz latite phenocryst assemblages alone
cannot buffer the observed bulk magma compositions,
and that the melt compositions must be externally
buffered, as for example, by crustal assimilation and/or
basaltic input facilitated by efficient mixing.
(9) Although much of the above discussion has emphasized the role of continental crust as a major contributor to the Awahab quartz latite magmas, their
chemistries, volumes and high temperatures are here
considered compelling evidence that a significant basaltic
input into the magmas has occurred, from both heat
transfer and chemical viewpoints. The lack of correspondence of the quartz latite melt compositions with
experimental ‘crustal’ melts, and their relatively high
FeO∗ and TiO2 (although not MgO) are taken as additional support for basaltic input. An additional, although indirect line of evidence is provided by
contemporaneous existence of silicic and mafic magmas
within the Messum igneous complex, and the Goboboseb
volcanic sequence. The Etendeka regional correlations
of high- and low-Ti basalts with high- and low-Ti quartz
latites and latites also imply significant basaltic input.
Quartz latite modelling
Details of the end-member compositions, and general
proceedures, follow those given in Part I (e.g. fig. 17 and
table A1).
Goboboseb quartz latite units
Least-squares modelling does not support simple mixing
between the model mafic end-members and either of the
defined silicic end-members (or, indeed, a variety of other
silicic end-members tried). Specifically, the fitted models
attempting to reproduce the Goboboseb QL compositions
are too high in MgO and Al2O3, and low in FeO∗, which
suggests that contemporaneous fractionation of cpx ±
NUMBER 2
FEBRUARY 1998
oliv + pl has occurred. Conventional AFC modelling,
using analysed mineral phases from the LTZ.H basalts,
produces consistently better fits between the mixed Lo.L.
crust + plume mafic ‘parental’ end-member (Part I, table
A1), and the Cascata leucogranite end-member, than
with any other combinations tried. The best fit models
were obtained using the mineral assemblages
cpx + oliv + pl, opx + pl, pl + opx + oliv + cpx,
pl + opx + oliv and pl + oliv + cpx + Ti-mag, with
the following phase compositions: pl = An66–70; oliv =
Fo85–86; cpx = Wo41–43Fe11–12Mg45–48; opx =
Wo3·9Fe16·1Mg80·0.
Table 8 details the models based on two of the
four-phase assemblages: pl + oliv + cpx + mag and
pl + opx + oliv + cpx. Orthopyroxene is included as a
phase in one of the models because its fractionation is
predicted in basaltic magma–silicic crust interactions [e.g.
modelling procedures of Ghiorso et al. (1994)]. Trace
element and isotopic data are tabulated for these specific
models using two sets of K D values, one appropriate to
basaltic magmas and the second to dacitic magmas (Part
I, table A2). Calculated results show variable degrees of
agreement with observed data, but they do highlight
some significant discrepancies which fall into two general
categories: those elements controlled by the assumed
silicic assimilants, and those controlled by the assumed
model mafic parental compositions. Of the former, Sr
and Zr are the most problematic, needing lower and
higher concentrations in the assimilant, respectively. In
the second group, V and especially Cu are discrepant,
requiring significantly lower abundances in the basaltic
end-member component. To evaluate these discrepancies, the ‘ideal’ calculated ranges of the trace
element concentrations in the basaltic (Cu only) and
silicic end-members are shown (Table 8), which provide
the best fits for all trace elements for the selected models.
Comparison of these calculated element abundances of
the assimilant with average crustal abundances (Table 6)
shows general agreement, although suggesting that the
higher REE and HFSE abundances required by the
calculations are closer to those of A-type granites and
some charnockites (references in Fig. 7 caption). The
calculated Sr and Nd isotope data show reasonable
agreement with observed compositions when the dacitic
K D values are used, and the observed negative Eu anomalies are reproduced.
In summary, the AFC calculations provide general
support for a model of quartz latite genesis involving the
interaction of mafic and silicic melts, together with mafic
and silicic crust, as originally advocated by Bellieni et al.
(1984, 1986). The magnitudes of the Awahab magmas,
however, preclude genesis in localized AFC systems,
favouring very large scale, dynamic open systems instead.
The open system equations of Aitcheson & Forrest [1994;
specifically their equations (10) and (11)] have therefore
246
247
9·50
66·57
1·01
12·73
6·69
0·12
0·99
3·43
3·63
4·53
0·30
175
712
149
299
25
5·1
46
21
2·9
24
84
17·9
48
46·1
99·7
47·5
4·35
0·66
183
−7·23
Actual
Goboboseb QL1
Cascata leucogranite1
Pl (An66)
Magnetite
Oliv ( Fo77)
Cpx
9·02
66·60
0·99
12·70
6·65
0·14
0·91
3·30
3·15
4·00
0·26
155
743
297
180
20
7·1
147
19
2·9
22
84
72
31
30·6
62·3
31·7
2·5
0·94
145
−7·89
57·94%
2·27%
24·02%
15·76%
Basaltic
1·09
1·15
0·54
Estimated
154
732
137
175
20
3·5
72
8·5
9·3
21
73
69
27
30·0
60·5
29·2
2·2
0·73
197
−7·94
Estimated
Dacitic
Mixed ‘plume’ + lowest crust1
35
45–85
[20–25]4
56
53–56
115–117
50–56
4·3–4·9
330–340
24–25
240
520–550
Preferred optimum
range for assimilant3
9·45
66·42
1·19
12·58
6·43
0·16
0·85
3·25
3·40
4·32
0·24
170
743
276
179
20
19
193
18
23
24
80
66
31
30·7
61·3
31·8
2·4
0·94
169
−8·11
170
731
131
175
20
3·4
132
8·2
2·9
23
51
63
29
28·1
59·0
30·3
2·1
0·73
226
−8·14
Mixed ‘plume’ + lowest lower crust1
Goboboseb QL1
Cascata leucogranite1
Pl (An70)
58·77%
Opx
30·96%
Oliv ( Fo85)
5·72%
Cpx
4·56%
Basaltic
Dacitic
1·25
1·38
0·30
Estimated
Estimated
2
See Appendix table A1 (Part I) for details of model compositions used.
See Appendix table A2 (Part I) for partition coefficients used.
3
Preferred values calculated for assimilant composition to give best fits to Goboboseb QL composition, for range of K D values.
4
Values calculated for starting composition (source) to give best fits, for a range of K D values (basaltic and dacitic).
1
d18O
SiO2 (wt %)
TiO2
Al2O3
FeO∗
MnO
MgO
CaO
Na2O
K 2O
P 2O5
Rb (ppm)
Ba
Sr
Zr
Nb
Cr
V
Sc
Ni
Pb
Zn
Cu
Y
La
Ce
Nd
Yb
Eu/Eu∗
eSr
eNd
K D values:2
F-value:
r-value:
R 2:
Starting composition :
Target composition:
Assimilant:
Fractionating phases:
Table 8: Least-squares models of derivation of Goboboseb quartz latite by AFC processes
31–33
45–85
[20–25]4
47–50
52–56
110–115
48–52
4·2–4·8
300–310
20
220
525–550
Preferred optimum
range for assimilant3
EWART et al.
ETENDEKA VOLCANISM, PART II
JOURNAL OF PETROLOGY
VOLUME 39
been applied. These give an indication of q (crust/magma
ratio; recharge model), b (rate of replenishment/rate of
assimilation) and r (rate of assimilation/rate of fractional
crystallization). To apply these equations to the Awahab
system, the following end-members are assumed: a parent
of Lo.L. crust composition, a replenishing magma equivalent to the ‘plume’ end-member (Sample SMG127),
and the Cascata granites as silicic crustal (assimilant) endmember (with the ‘preferred’ calculated REE and HFSE
abundances, and dacitic bulk K D values, as listed in
Table 8).
For the opx-free phase assemblage, the calculated trace
element curves converge at approximate values of q =
0·85–0·95; b = 0·15–0·25; r = 1·5. With the opxbearing assemblage, the values are q = 0·9–1·0, b =
0·2–0·6; and r = 1·5 (Fig. 8). Similar calculations for
Sr–Nd isotopes are consistent with these values, except
that values of r converge near 0·95. In these calculations,
Cr, Sc, Ni, Cu and V do not provide good solutions,
and the intersections shown are based only on REE, Zr,
Nb, Sr, Ba, Rb, Zn and Pb.
The results suggest that, within the limits of the assumed
model, the Goboboseb quartz latite melts represent nearly
equal input of crust and mafic magma with rates of
assimilation exceeding rates of crystallization, but with
relatively low rates of replenishment relative to assimilation. The results of both the above approaches to
AFC modelling point to rather high levels of assimilation,
and thus crust/magma ratios in the quartz latite melts
[see also Reiners et al. (1995)]. These results are feasible
in terms of the models presented in Part I (figs 17 and
18), if extensive initial assimilation occurred in, and with,
mafic lower crust (as required by the starting composition
used in Table 8), followed by continuing upward migrating fusion fronts (with increasing rates of fractional
crystallization) into silicic crust, producing the quartz
latite magma systems. As shown in Part I, mixing of
‘plume’ melt with both mafic lower and silicic upper
crust can account for the similarity of eNd between LTZ.L
basalts and quartz latites.
Springbok quartz latite
The shift in isotopic compositions between the Goboboseb and Springbok units is significant, the measured
ranges being, respectively: eSr 183–202 and 198–235;
eNd –7·23 to –9·14 and –7·44 to –9·01; 206Pb/204Pb
18·97–19·02 and 19·08–19·19. Although 143Nd/144Nd
ratios show no systematic change, a small increase of
T DM,Nd ages from 1·3–1·4 (Goboboseb) to 1·4–1·6 Ga
(Springbok) is noted. In terms of the model presented
here, the isotopic changes are interpreted as the continuation of the mixing–AFC processes inferred for the
Goboboseb quartz latites, previously described. Shifts
towards more radiogenic Sr and Pb compositions indicate
NUMBER 2
FEBRUARY 1998
either increasing input of the same silicic crustal endmember component, or the involvement of a new, more
radiogenic crustal end-member component during the
continued melt evolution. Associated trace element
changes include increasing Pb, and generally decreasing
LILE, HFSE and REE from the Goboboseb through to
the Springbok quartz latite units.
A simple mixing model between the Goboboseb quartz
latite and the Cascata leucogranite (Table 9) results in
feasible solutions, although with low calculated eSr and
eNd. Likewise, the foliated Cascata granitoid also produces
good major and trace element fits, but with even lower
calculated eSr and eNd. Nevertheless, these models are
consistent with the interpretation of an upward migrating
melt front, with incorporation of increasing silicic (and
more radiogenic Pb and Sr) crustal end-member(s). To
reproduce the observed higher eSr values, however, requires the crustal end-member to have eSr >500. This
implies incorporation of older sequences of Damara
sediments (Fig. 9) from shallower crustal levels into the
upward migrating, actively developing magma system.
DISCUSSION
Numerical modelling indicates that the quartz latite melts
can result from large-scale AFC-style processes, involving
high-degree lower- and upper-crustal melting with thermal and mass transfer input from plume and hybrid
LTZ.L-type melts (Part I). Evidence that the Awahab
magma system was driven by an unusually large thermal
anomaly is surely provided by the large volumes and
high temperatures of the silicic volcanism. The silicic
crustal end-member is modelled on the Parana´ equivalents of the Damara granites. The early Proterozoic
(or older) basement, or their derived sediments do not
possess the appropriate Sr–Nd isotopic source compositions, and based on available isotopic data, we believe
that the restite source(s) of the Damara granites are
plausible quartz latite crustal melt sources.
Kilpatrick & Ellis (1992) emphasized the mineralogical
and chemical similarities between quartz latite and
charnockite magma types (see Table 6). Although not
described specifically from the Damara fold belt, charnockites and feldspathic granulites, of mid-Proterozoic
metamorphic age, are widespread in southern Namibia
(extending south into the South African Namaqua mobile
belt), and encircle the western zone of the Kalahari–
Kapvaal cratons ( Jackson, 1979). Charnockites represent
deeper crustal lithologies, which are commonly characterized by relatively high HFSE and REE, both inferred
to be geochemical features required by the proposed
Awahab magma source(s). According to Kilpatrick &
Ellis (1992), charnockitic magmas result from partial
fusion of hornblende-free, LILE-enriched, biotite-bearing
granulite sources which were dehydrated by a previous
248
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ETENDEKA VOLCANISM, PART II
Fig. 8. Application of the open system AFC equation of Aitcheson & Forrest (1994) for selected trace elements to model the Goboboseb quartz
latites (see text for details). Two sets of models are shown, involving fractionation of (a) pl + oliv + cpx + mag assemblage, and (b)
pl + oliv + opx + cpx assemblage. The trace element curves converge towards values of q = 0·85–0·95, b = 0·15–0·25 and r = 1·5 (a) and
q = 0·9–1·0; b = 0·2–0·6 and r = 1·3 (b).
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Table 9: Least-squares AFC and mixing models linking Goboboseb and Springbok quartz
latites
Starting composition: Goboboseb QL
Goboboseb QL
Target composition: Springbok QL
Springbok QL
Assimilated material: Namaqualand mafic granulites1
Cascata Leucogranite1
Fractionating phases (wt %):
Pigeonite:
34·61%
Magnetite:
12·17%
Apatite:
Plagioclase (An50):
91·77% (Goboboseb QL)
8·23% (assimilant)
Quartz latite2
Partition coefficients:
F-value:
0·991
r-value:
0·89
R 2:
Actual
SiO2 (wt %)
3·72%
49·51%
Mixing only
0·070
0·2 (major elements)
Estimated
Estimated
67·18
67·11
TiO2
0·93
0·92
0·94
Al2O3
12·99
12·94
12·81
FeO+
6·09
6·04
6·26
MnO
0·10
0·11
0·11
MgO
1·01
1·05
0·93
CaO
3·50
3·50
3·25
Na2O
3·38
3·56
3·61
K 2O
4·47
4·60
4·58
P 2 O5
0·28
0·18
67·23
0·28
Rb (ppm)
174
179
178
Ba
678
707
700
Sr
145
126
149
Zr
277
307
285
Nb
23
25
24
Cr
5·3
11
6·2
V
48
48
44
Sc
19·8
16
19
Ni
2·6
8·7
2·9
Pb
25
25
24
Zn
81
64
80
Cu
18
19
17
Y
44
39
46
La
45·2
42·5
44·7
Ce
95·9
89·1
95·9
Nd
44·9
39·9
45·7
Yb
4·0
3·5
4·1
Eu/Eu∗
0·65
0·77
0·68
eSr
235 (198–235)
158
206
eNd
−9·0 (–7·4 to 9·0)
−7·48
−7·34
9·29
9·71
d18O
10·50
1
See Appendix table A1 (Part I) for details of model compositions used.
See Appendix table A2 (Part I) for partition coefficients used.
2
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ETENDEKA VOLCANISM, PART II
Fig. 9. eNd–eSr plot (calculated at 132 Ma) showing the generalized Awahab compositional fields in relation to model lower crust, Lo.L. crust,
and upper-crust compositions (Rudnick & Presper, 1990; Zartman & Haines, 1988), SCLM 1 and 2 and OPM (oceanic-plume magmas) (Gibson
et al., 1995) and SCLM3 (Zartman & Haines, 1988); the four major Namibian Damara sediment formations (McDermott et al., 1989); and the
Damara equivalent granite phases in Brazil, plus the Parana´ basement Encantada Gneisses (May, 1990).
partial melting event. We suggest this to be a plausible
scenario for the Awahab magmas, involving melting
of Mid-Proterozoic crustal silicic granulite–charnockitic
basement, previously partially dehydrated during Damaran granite magmatism. AFC plots (not shown) suggest
that the Awahab quartz latite melt chemistries were more
probably in equilibrium with a biotite-bearing, rather
than hornblende-bearing source(s), and are consistent
with the high Fe and K chemistries of the quartz latite
melts, and the potassic inclusions within the plagioclase
phenocrysts. Henson & Osanai (1994) confirmed that
fluor-biotite is stable to 1000°C (at 9 kbar) in high-grade
metamorphic rocks of appropriate composition (e.g. hightemperature granulites), including those having undergone prior melt loss. The presence of fluor-amphibole in
the quartz monzonite (albeit in low abundances), and
the water-undersaturated condition of the QL melts have
already been noted.
Seismic reflection profiles across the central Damara
belt (including Messum; Green, 1983), indicate a simple
velocity crustal section, the absence of a high-velocity
lower crust, and relatively low-velocity upper mantle.
Green (1983) interpreted the reduction and restructuring
of the lower crust to result from crustal thinning, partial
melting, and merging with pre-existing upper mantle
during the Damara orogen. As an alternative, we suggest
that the seismic data could equally reflect the lithospheric
thinning, large volume crustal melting, hybridization and
assimilation accompanying the Awahab magmatic event.
The heat source required to produce the Awahab
magmatism was apparently of unusually large magnitude,
and we therefore follow the interpretation of White &
McKenzie (1989) that this source was plume initiated,
and has resulted in a common origin for the magmatism
of the whole Parana´–Rio Grande Rise–Walvis Ridge–
Etendeka systems. Lithospheric stretching over the hotspot, together with the apparent positioning of the plume
beneath the Etendeka where predicted highest temperatures will occur above the core of the new plume
system (White, 1993), resulted in the focussing of lithospheric and crustal heat, and fluid flow through prexisting
crustal lineaments and heterogeneities between cratonic
blocks (Thompson & Gibson, 1991). These are further
believed to have been factors controlling localization of
specific magmatic centres. Furthermore, quartz latite
magmatism seems to occur dominantly on either side
of the Atlantic margin, where crustal and lithospheric
attenuation from rifting is greatest.
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VOLUME 39
Henry et al. (1990) proposed that the Damara basement
developed in a zone of continental collision, with the
potential both for tectonic thickening and interleaving of
lower- and upper-crustal components, also consistent
with the range of T NdChur ages reported for Damaran
leucogranites. The existence of a deep, tectonically imbricated upper-crustal segment, if overlain by more mafic
lower crust, would provide an effective mechanism of
producing very large volume, high-temperature silicic
melts, in which the overlying mafic zone acts as a thermal
lid, thereby inhibiting premature upward leakage of
smaller volume silicic melts.
ACKNOWLEDGEMENTS
We wish to again thank M. Wilson, S. Turner and an
anonymous reviewer for valuable comments and input
to the manuscript.
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