Journal of Petrology Advance Access published online on April 28, 2007
Journal of Petrology, doi:10.1093/petrology/egm012
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Origin of CFB Magmatism: Multi-tiered Intracrustal Picrite–Rhyolite Magmatic Plumbing at Spitzkoppe, Western Namibia, during Early Cretaceous Etendeka Magmatism
1Department of Geological Sciences, Durham University, South Road, Durham DH1 3LE, UK
2Division of Geological and Planetary Sciences, California Institute of Technology, MC 170-25, Pasadena, CA 91125, USA
3Department of Geology, McMaster University, 1280 Main Street West, Hamilton, Ontario, Canada L8S 4M1
4British Geological Survey, Keyworth, Nottingham NG12 5GG, UK
5Geological Survey of Namibia, 1 Aviation Road, Private Bag 13297, Windhoek, Namibia
Received February 22, 2006; Revised typescript accepted February 27, 2007
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONEarly Cretaceous tholeiitic picrite-to-rhyolite dykes around Spitzkoppe, western Namibia, are part of the extensive Henties Bay–Outjo swarm, penecontemporaneous with
132 Ma Etendeka lavas 100 km to the NW. Although only intermediate to rhyolitic dykes contain clinopyroxene phenocrysts, the behaviour of Ca, Al and Sc in the dyke suite shows that liquidus clinopyroxene—together with olivine—was a fractionating phase when MgO fell to 9 wt %. Both a plot of CIPW normative di–hy–ol–ne–Q and modelling using (p)MELTS show that a mid-crustal pressure of 0·6 GPa is consistent with this early clinopyroxene saturation. Sr, Nd, Hf and Pb isotope variations all show trends consistent with AFC contamination (assimilation linked to fractional crystallization), involving Pan-African Damara belt continental crust. The geochemical variation, including isenthalpic AFC modelling using (p)MELTS, suggests that the picrites (olivine-rich cumulate suspensions) were interacting with granulite-facies metamorphic lower crust, the intermediate compositions with amphibolite-facies middle crust, and the rhyolitic dykes (and a few of the basalts) with the Pan-African granites of the upper crust. The calculated densities of the magmas fall systematically from picrite to rhyolite and suggest a magmatic system resembling a stack of sills throughout the crust beneath Spitzkoppe, with the storage and fractionation depth of each magma fraction controlled by its density. Elemental and isotopic features of the 20 wt % MgO picrites (including Os isotopes) suggest that their parental melts probably originated by fusion of mid-ocean ridge basalt (MORB) source convecting mantle, followed by limited reaction with sub-continental lithospheric mantle metasomatized just prior to the formation of the parental magmas. Many of the distinctive features of large-volume picritic–basaltic magmas may not be derived from their ultimate mantle sources, but may instead be the results of complex polybaric fractional crystallization and multi-component crustal contamination. KEY WORDS: flood basalts; spitzkoppe; picrite; trace elements; hafnium isotopes; Etendeka
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INTRODUCTION |
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Attempts to understand the genesis and evolution of magmas can be subdivided into: (1) determining the physical and chemical conditions of magma genesis; (2) understanding the processes that have affected the magmas between their genesis and final solidification. Researchers interested in the first process are inclined to see the second one as little more than a nuisance, standing between them and their objectives. For this reason it is common for such students of magmas to focus on samples, techniques or datasets that appear to them to evade the ‘problem’ of post-genesis processes. The flood basalts of large igneous provinces (LIPs) in general, and the early Cretaceous Paraná–Etendeka province in particular, have been and remain favourite battlefields for differing views about the relative importance of the two processes. The many publications that attempt to discern the ultimate sources of these magmas are split between favouring either the sub-continental lithospheric mantle (SCLM) or the underlying convecting mantle, with or without a contribution from either delaminated SCLM or older subducted material. Recent examples of such studies are those by Peate (1997
), Ewart et al. (1998a, 2004a), Gibson et al. (2000, 2005), Marsh et al. (2001), Thompson et al. (2001), Trumbull et al. (2004a) and Tuff et al. (2005). All these studies acknowledge the probability that hot upwelling magmas will react with and absorb relatively fusible continental crust. Some studies attempt to make allowances for such crustal contamination, usually by comparing the predicted chemical effects of a likely crustal contaminant, such as the local granite basement, with geochemical trends observed in the magmatic suite. Most published studies assume a priori that any such crustal contamination takes place by a process of assimilation linked to the fractional crystallization of the melt (AFC), such that magmas become progressively more contaminated with falling liquidus temperature. Hence, they routinely reverse this reasoning and propose that the best shortcut to identifying magmas that have undergone little or no reaction with continental crust is to focus only on those with relatively high MgO contents, and hence high liquidus temperatures.
This is difficult to do in the Paraná–Etendeka province. For example, the compendium of data by Peate (1997) shows almost no Paraná lava analyses with >7 wt % MgO. Likewise, Marsh et al. (2001) and Ewart et al. (2004a) showed that relatively few Etendeka lavas—excluding the basal ferropicrites (Gibson et al., 2000)—have >9 wt % MgO. Even this higher MgO content falls far short of the lowest value calculated for melts to be in equilibrium with a peridotite mantle. Primary magmas do not fall below 13–14 wt % (Thompson et al., 2005) and rise to values above 20 wt % MgO. Therefore all arguments that the elemental and isotopic variations in ‘more Mg-rich’ Paraná–Etendeka lavas are inherited from their mantle sources and unaffected by subsequent polybaric fractional crystallization and concomitant crustal contamination (e.g. Erlank et al., 1984, and many subsequently) are, to some extent, acts of faith.
This study is concerned with a picrite–rhyolite suite of dykes exposed within the Etendeka igneous province, south of the lavas. They are approximately contemporaneous with the Etendeka extrusive rocks and often taken to be their hypabyssal equivalents (e.g. Marsh et al., 1997; Trumbull et al, 2004a). All but one of these dykes shows clear geochemical evidence, detailed below, of reaction between the melts and the local continental crust. The only possible exception is one dyke with 20 wt % MgO. The AFC reactions involved at least three distinctive crustal rock-types, at upper-, middle- and lower-crustal depths. Clinopyroxene precipitated, deep in the crust, when MgO in these fractionating melts reached 9 wt % MgO. The combined effects of this process and the crustal contamination strongly affected many inter-element ratios in the magmas (including some routinely used to classify Paraná–Etendeka magma types), together with isotopes of Sr, Nd, Hf and Pb. This confirms other recent studies, which showed that, even when dealing with Mg-rich basalts and picrites (e.g. Kent et al., 2002; Yaxley et al., 2004; Saal et al., 2005; Harlou et al., 2006; Zhang et al., 2006), post-genesis interactions between the upwelling melts and their surroundings are common.
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SPITZKOPPE DYKE SWARM |
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The dyke swarm exposed in the Spitzkoppe area (Fig. 1) is part of the major Henties Bay–Outjo (HB–O) swarm (Lord et al., 1996
; Trumbull et al., 2004a). The HB–O swarm mostly trends between NE–SW and NNE–SSW, and extends inland for at least 500 km from the continental margin. It forms part of the southern Etendeka Igneous Province (Marsh et al., 2001). Whereas most Etendeka lavas are preserved in a belt adjacent to the Atlantic coast, dyke swarms and intrusive sub-volcanic centres extend inland within the Late Proterozoic to Cambrian Damara orogenic belt, along the southern margin of the Congo craton (Fig. 1, inset). The Cretaceous extension and magmatism of this region is generally considered to be an integral part of the igneous event that produced the Paraná–Etendeka lavas, more deeply eroded than the coastal belt (e.g. Milner & le Roex, 1996; le Roex & Lanyon, 1998; Marsh et al., 1997, 2001; Wigand et al., 2003). Trumbull et al. (2004a, fig. 1) defined the HB–O swarm as 70 km wide in the vicinity of Spitzkoppe. To the NW there is a relatively narrow (25 km) dyke-poor zone before reaching the dense parallel dyke swarm between the Cretaceous Brandberg and Messum intrusive complexes (Thompson et al., 2001). For most of its length the HB–O dyke swarm is emplaced in metamorphic rocks—mica schists, migmatites, calc-silicates and marbles—and granites (sensu lato) of the Pan-African Damara belt but, north of Outjo, it enters the Archaean–Proterozoic Congo Craton (Fig. 1, inset).
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At first sight the possible maximum age of the HB–O dykes is only constrained geologically to be post-Damara, whereas their minimum age is constrained by the radiometric dates of the Cretaceous plutonic complexes that intersect them (Fig. 1, inset). Dykes of the Jurassic Karoo igneous province do not occur in central and NW Namibia. The closest Karoo dykes are at the NW extremity of the 179 Ma (Ar–Ar dates) Okavango giant dyke swarm, which terminates in NE Namibia—about 750 km NE of Spitzkoppe—and trends at a right angle to the HB–O swarm (Le Galle et al., 2002
). Trumbull et al. (2004a) discussed the age of the HB–O dykes in detail and noted that their only published radiometric dates were obtained using the K–Ar method and mostly fall in the range 116–143 Ma. All the HB–O dyke samples chemically analysed to date form a single coherent tholeiitic picrite–rhyolite suite (Marsh et al., 1997; Trumbull et al., 2004a; this paper). The picrite–intermediate compositions also occur in the abundant dykes SW of Brandberg and in the Huab and Horingbaai areas (Fig. 1, inset), where they intersect 133 Ma southern Etendeka basal lavas. Both they and the lavas are transected by the 131 Ma Brandberg intrusive complex (Thompson et al., 2001). Immediately NE of Spitzkoppe the HB–O dykes are transected by the large Erongo central volcanic and intrusive complex (Fig. 1, inset), dated at 133–130 Ma (with errors ±0·8–1·9 Ma) by Wigand et al. (2003) using both Ar–Ar biotite and U–Pb zircon methods. The Spitzkoppe swarm dykes cut Damara basement in the relatively flat surroundings of the prominent peaks formed by the Gross and Klein Spitzkoppe granite stocks (Fig. 1). They vary in width from 1 m to 10 m and their trends range from 10° to 90°, with most falling between 30° and 50°. Two of the picrites (VB19 and 32) occur as dykes between 20 m and >30 m thick; both form prominent topographic ridges.
Frindt et al. (2004a) and Trumbull et al. (2004b) both noted that felsic dykes are more common around Spitzkoppe than elsewhere in the HB–O swarm, although they have the same trends as their mafic neighbours. On a regional scale, the Spitzkoppe granites appear to transect all the dykes and thus post-date them. Nevertheless, Frindt et al. (2004a) have shown that several of the mafic regional dykes penetrate the marginal granite. Within the granite, the mafic dykes have irregular margins and are cut by late-stage granitic veins. One mafic dyke also locally forms a zone of pillow-like enclaves within the granite. The chilled cuspate margins of the enclaves are typical of liquid–liquid magmatic contacts elsewhere and leave no doubt that the basic and acid melts coexisted (Frindt et al. 2004a; S. Frindt, personal communication, 2005). The whole-rock Rb–Sr isochron age of the Gross Spitzkoppe stock is 124·6 Ma ± 1·1 Ma (MSWD = 2·5), with initial 87Sr/86Sr = 0·7134 (Frindt et al., 2004b) but this may be lower than the true emplacement and original cooling age, as a result of subsequent hydrothermal redistribution of Rb (S. Frindt, personal communication, 2005). A clear example of this hydrothermal redistribution of Rb in a pluton is the Beinn an Dubhaich granite, Skye, NW Scotland. Dickin (1980) obtained an Rb–Sr whole-rock isochron age for this pluton of 53·5 ± 0·8 Ma (2SD), whereas its zircon U–Pb age (M. A. Hamilton, D. G. Pearson & R. N. Thompson, unpublished data) is 55·89 ± 0·15 Ma (2SD). It therefore remains uncertain whether the Spitzkoppe dyke swarm was emplaced at 125 Ma or a few million years earlier, such as the 132 Ma date for similar mafic dykes in the Brandberg, Huab and Horingbaai areas to the north and the nearby Erongo complex (Wigand et al., 2003). The weight of evidence regionally clearly favours the 132 Ma date.
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PETROGRAPHY AND PICRITE MINERALOGY |
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The thick picrite dykes are coarse and gabbroic. Their abundant olivines enclose only small chromites, whereas the plagioclase and clinopyroxene are intergrown. Small orthopyroxenes are scattered around the margins of the olivines and there are traces of interstitial biotite. Basalt dykes are finer grained and either aphyric (intergrown olivine, plagioclase and clinopyroxene) or contain small olivine phenocrysts. Intermediate dykes (5·4–1·7 wt % MgO) have sparse plagioclase phenocrysts; some euhedral and others rounded. Early pigeonite has been identified optically in VB8. The groundmass becomes progressively finer grained until it is devitrified glass in the felsic dykes. The dominant phenocrysts in the felsic dykes are quartz and turbid sericitized feldspar. Clinopyroxene phenocrysts are extremely scarce (two out of 32 samples). The picrites are almost devoid of visible low-temperature alteration but reaction of ferromagnesian minerals to chlorite ranges from
10% to >80% in some of the less Mg-rich dykes. Our geochemical tests of the effects of this alteration are described below.
Picrite mineral compositions
Samples from two thick picrite dykes were studied by electron probe microanalysis at Manchester University [see Thompson et al. (2005) for method].
Olivine
Representative olivine analyses are given in Table 1. Figure 2 shows individual microprobe analyses as functions of the Mg-number of the host dyke. Rock Mg-number was calculated with 10% of total Fe as Fe3+ (Thompson & Gibson, 2000). Two values of Kd for the distribution of Fe2+ and Mg between olivine and melt are plotted: 0·30 is appropriate for equilibria between picrites and their olivines at 1 atm, whereas 0·31 applies at 0·5 GPa pressure (Ulmer, 1989). Garcia et al. (1995) and Garcia (1996) showed how some Hawaiian basalts and picrites from Mauna Loa and Mauna Kea are in equilibrium with their olivines, whereas others are best explained as cumulates because their olivines plot to the right of the equilibrium Kd (i.e. the bulk composition is more magnesian than would be appropriate to precipitate the olivines it contains). Using this reasoning, the olivine phenocrysts in both VB18 and 32 are clearly accumulative. The Spitzkoppe picrites therefore appear to be suspensions of olivine phenocrysts in basaltic liquids, rather than originally picritic melts. Olivines from the nearby (100 km to the north) Horingbaai suite are also plotted in Fig. 2 for comparison (Thompson et al., 2001). An obvious difference between the two suites is the lack of very magnesian olivine macrocrysts in the two Spitzkoppe picrites studied. Of course, the possibility of some post-crystallization re-equilibration of these olivines in the thick dykes cannot be ruled out but, if this has not happened, then the VB32 olivines must have been precipitated from a melt with Mg-number 60 and the VB18 olivines from an even more Fe-rich melt.
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Chromite
Euhedral chromites up to a few millimetres in size occur within the olivines in both picrites. Representative analyses are given in Table 1. Cr2O3 varies between about 6 and 26 wt %, and Al2O3 between 2·7 and 6·0 wt % in both samples but a few points show Al2O3 values up to 15 wt %.
Clinopyroxene
This phase is clearly late-crystallizing in the picrites as it tends towards large poikilocrysts, enclosing both olivines and plagioclases. Typical analyses are given in Table 1.
Orthopyroxene
This mineral forms small subhedra, mostly close to the margins of olivines, and larger poikilitic crystals. Two analyses are given in Table 1.
Biotite
Small interstitial crystals of biotite are visible in trace quantities within both picrites, especially VB32, and an analysis is given in Table 1. Without analysis for fluorine, it cannot be certain that this biotite is necessarily hydrous and therefore indicative of a wet magma (see below).
Alkali feldspar
Element maps show that alkali feldspar also occurs interstitially in VB32.
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GEOCHEMISTRY |
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Methods
Analytical methods are described in Electronic Appendix A, available for downloading at http://www.petrology.oxfordjournals.org.
Results
Element and oxide variation
The analyses (Table 2) show wide ranges of MgO (20·1–0·1 wt %) and SiO2 (46·0–68·2 wt %). Additionally, five acid dykes from around the Gross Spitzkoppe granite stock, analysed by Frindt et al. (2004a) are plotted in the following diagrams where appropriate. Samples with >12 wt % MgO are called picrites in this paper (Le Bas, 2000). On a plot of (Na2O + K2O) vs SiO2 (TAS; Le Maitre, 2002) the new analyses plot along the same trend as published data for the HB–O swarm dykes (Trumbull et al., 2004a, fig. 3), through the fields basalt, andesite, dacite and rhyolite but adjacent to the boundary with the basalt–trachyte suite.
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A plot of selected oxides (wt %) and trace elements (ppm) vs MgO (Fig. 3) in the Spitzkoppe picrite–rhyolite dykes shows that both Cr and Ni fall progressively with MgO. Although the variations of SiO2, Cr and Ni with MgO are smooth curves in Fig. 3, the behaviour of the other major oxides and Sc in this diagram is more complex. Trend lines have been drawn (by eye) to illustrate qualitatively the extent to which each plot is described by an inflected trend, such as may occur during fractional crystallization of melts by the addition of a new mineral to the separating phases. Such inflected trend-lines describe the behaviour of CaO, Al2O3 and Sc well (Fig. 3). In the Fe2O3, TiO2 and P2O5 plots there is a handful of points lying far above the trends fitting the other data. Thin sections of the samples showing these relatively high Fe2O3, TiO2 and P2O5 values have visible populations of titanomagnetite and apatite microphenocrysts that may be accumulative.
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, Spitzkoppe picrite–rhyolite dykes (
, Spitzkoppe acid dykes (Frindt et al, 2004a
) are used for all the dykes in the Cr–Ni plot, for clarity. Trend lines are drawn by eye. High- and low-Ti Etendeka lava fields from Marsh et al. (2001The trend inflections for Fe2O3 and TiO2 in Fig. 3 correspond approximately to the appearance of titanomagnetite in thin sections, as MgO falls to
3·5 wt %. Likewise, the inflection of the P2O5 trend at 3·0 wt % MgO marks the appearance of apatite microphenocrysts. The inflection at 6·5 wt % MgO in the Al2O3 plot corresponds to the appearance of plagioclase as a phenocryst. In contrast, the inflections at 9·0 wt % MgO in the CaO and Sc trends do not correspond to the first appearance as phenocrysts of any mineral, such as Ca-rich clinopyroxene, that might remove these two elements together from a melt. The only visible phenocryst in dykes with MgO contents on both sides of this inflection is olivine (sometimes with chromite inclusions), a mineral poor in both Ca and Sc. Clinopyroxene is conspicuously absent as phenocrysts in the Spitzkoppe basic and intermediate dykes.
The two fields drawn on the TiO2–MgO plot (Fig. 3) encompass >1000 analyses of lavas and hypabyssal intrusions from throughout the Etendeka Igneous Province (Marsh et al., 2001), showing how they divide into low-Ti and high-Ti magmatic suites. Between these fields a group of 12 analyses has been omitted for clarity. These are the Tafelkop basalts of Marsh et al. (2001), known by others as ‘LTZ.H’ (Ewart et al., 1998a) or ferropicrites–basalts (Gibson et al., 2000). It should be noted that, although the two fields in Fig. 3 and their subdivisions were described as mafic by Marsh et al. (2001), and most of the subdivisions are formally named ‘basalts’, the analyses defining the two fields include compositions with MgO as low as 3 wt %. The Spitzkoppe dykes clearly form part of the Etendeka low-Ti suite.
The array of patterns produced by rare earth element (REE) variation in the Spitzkoppe dykes (Fig. 4) is qualitatively concordant with the data on Fig. 3. The progressive increase in all the REE, except Eu, from picrites to rhyolites is consistent with a fractional crystallization process creating this magma suite. Apart from aphyric basalts VB20 (no Eu anomaly) and VB24 (small positive Eu anomaly), all the samples have negative Eu anomalies. These are approximately constant throughout dykes that lack plagioclase phenocrysts and increase in plagioclase-phyric intermediate and acid dykes (see Fig. 14 below).
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Isotopes
The radiogenic isotopic ratios of the Spitzkoppe dykes show large variations (Tables 3 and 4) and these rule out the possibility that their range of elemental compositions could be due to closed-system fractional crystallization of a basic magma. The isotopic ratios are corrected to an emplacement age of 132 Ma, for reasons explained above.
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Tests for effects of low-temperature alteration and very small-scale sample isotopic heterogeneity. The isotope data for dykes in Table 3 can be used to test the effects of low-temperature hydrothermal alteration because samples analysed at McMaster were all first leached in acid whereas those analysed at Durham were not. Eight dyke samples were analysed for Pb isotopes in both laboratories. Excluding the rhyolite sample VB30, the R2 value for the regression between the two sets of 206Pb/204Pb data (the ratio with the largest range of values) is 0·999. Rhyolite VB30 falls slightly off this regression (R2 drops to 0·994) and this may be due to the turbid hydrothermal alteration of its predominant alkali feldspar.
There are small significant differences between 87Sr/86Sr for the McMaster and Durham analyses of picrites VB18 and 32. Although these could be alteration related, they are not both in the same direction. There is also a significant difference between the results of the two laboratories for 143Nd/144Nd in VB18, discussed in detail in Electronic Appendix A. It is generally accepted that the latter are unlikely to be related to the slight alteration of this sample. These data suggest small-scale isotopic variability within the picrites—as might be expected from the crustal assimilation model to be developed below. Although lava samples are usually isotopically homogeneous at the hand-specimen level, it is becoming clear that this homogeneity is illusory, in that glass inclusions within the olivine phenocrysts of some crustally contaminated continental picrites are extremely heterogeneous in both elements and isotopes (e.g. Yaxley et al., 2004; Harlou et al., 2006). Such inhomogeneity should logically eventually be expected to occur between aliquots from a powdered individual olivine-phyric hand specimen, and samples VB18 and 32 appear confirm this point. Clearly, this matter requires more research.
Sr, Nd and Hf isotopes. On a plot of 87Sr/86Sr vs 
Nd (Fig. 5) the picrite–intermediate dykes form a steep continuous array. The picrites have Nd values just above and just below zero, within the field of ocean island basalts (OIB), whereas Nd values for the intermediate and rhyolite dykes fall in the range from –5 to –10. New analyses of crustal rocks comprise: Damara granite VB12, which crops out within the dyke swarm [Frindt et al. (2004b) analysed another Damara granite sample, SF17, 1·5 km SE of VB12, Fig. 1]; Damara granite (VB38) and associated aplite (VB39) and pegmatite alkali feldspar (VB37), which crop in the Horingbaai mafic dyke swarm, south of the Brandberg intrusive complex (Thompson et al., 2001); Damara phyllite (VB41) and metasedimentary schist (VB42) between the Brandberg and Huab areas (Fig. 1). Some of these crustal rocks tend to enlarge the Sr–Nd isotopic fields for Damara granites and metasediments plotted by Frindt et al. (2004b). A mixing curve between picrite dyke VB32 (highest Nd) and the local Damara granite (VB12/SF17) does not pass through the dyke data. The field for pre-Damara basement in the Spitzkoppe area is speculative and follows the suggestion of Frindt et al. (2004b) and others that the Damara belt may be underlain by the same Archaean–Proterozoic gneisses as form the Congo–Angola craton in Kaokoland, NW Namibia (Seth et al., 1998). If this is so, such a deep-crust basement would form a suitable crustal end-member for a mixing curve passing through the picrite and basalt dyke data in Fig. 5. Tangible isotopic evidence for a deep-crust basement of this composition is provided by rhyolites from the Cretaceous Paresis sub-volcanic complex, NE of Spitzkoppe (Fig. 1, inset), which have Sr–Nd isotopic ratios (Fig. 5) very similar to those of the Kaokoland gneisses (Mingram et al., 2000). In this paper we use Paresis rhyolite PA19 (Mingram et al., 2000) as a geochemical proxy for partial melts of acid–intermediate lower crust in the Damara region. This is obviously less satisfactory than having lower crust samples available.
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Table 3 and Fig. 6 show the first Hf-isotope data published for both Etendeka igneous rocks and Damara belt crustal rocks. The dyke data show covariation of
Nd and Hf, such that the picrites and basalts fall in the OIB field whereas the intermediate and rhyolite dykes have similar Nd–Hf isotope ratios to the local Damara belt exposed crust. The MgO content of each Spitzkoppe dyke is marked in Fig. 6, so that the overall correlation between MgO and Nd–Hf isotopic ratios in the dykes can be seen.
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Pb isotopes. When it became clear that Pb isotope variation in the Spitzkoppe dyke swarm is complex, we analysed sample sets at both Durham and McMaster, so as to maximize the data and also to provide interlaboratory cross-checks (Table 3). At first sight, Pb isotopes in the Spitzkoppe dykes (Table 3) behave in much the same way as their Sr–Nd–Hf isotopes, in that the picrites and Mg-rich basalts form a linear array, with a good correlation between Pb isotopes and MgO in the 20–9 wt % range (Fig. 7). However, the rhyolite dyke VB30, and also dyke SF11d of Frindt et al. (2004b
) show a major difference. Instead of continuing the trend of the mafic dykes to lower 206Pb/204Pb ratios, they have relatively high 206Pb/204Pb and 208Pb/204Pb ratios and thus, in terms of Pb isotopes, appear to be unrelated to the picrites and basalts. The Mg-poor basalts and intermediate dykes (MgO 8–3 wt %) divide between these Pb isotopic groups; the majority have relatively low 206Pb/204Pb, combined with slightly higher 208Pb/204Pb than the Mg-rich basalts (Fig. 7), but two Mg-poor basalts, VB14 and 16, cluster with the rhyolite dykes in terms of their Pb isotopes. The field for Damara S-type granites (Frindt et al., 2004b) must be enlarged slightly to incorporate our new Pb-isotope data for these rock-types (Table 3 and Fig. 7) and the Spitzkoppe rhyolite dykes plot within the enlarged fields. They also lie close to the Gross Spitzkoppe granite field (Frindt et al., 2004b). The field for southern Etendeka LTZ lavas and dykes incorporates data for so-called continental flood basalt (CFB)-like dykes in the nearby Brandberg area (Thompson et al., 2001; Fig. 1, inset). This field has been redefined in Fig. 7 by reanalysis of key samples (Table 3). After such standardization, it becomes clear in Fig. 7 that the elongation of this field is collinear with Pb isotopic variation in the Spitzkoppe picrites and Mg-rich basalts. The most Mg-rich Spitzkoppe picrites have 206Pb/204Pb >19·0, which is higher than values of this ratio in both Tristan da Cunha and Walvis Ridge lavas, and also in previously studied Damaraland Cretaceous picrite–basalt dykes (Thompson et al., 2001; Gibson et al., 2005).
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.Os isotopes. Osmium isotopes were determined on separated olivine phenocrysts from picrites VB18 and 32, to see whether this early crystallizing phase might have trapped oxides or sulphides before the magmas were exposed to crustal contamination. The separates were made by hand-picking, with fragments containing opaque inclusions left in the separate, rather than removed. Using an emplacement age of 132 Ma, the initial 187Os/188Os ratios of the picrite olivines (Table 4) are 0·126248 (VB32) and 0·124262 (VB18). The significance of these results will be discussed below in the section ‘Focus on the picrites’.
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DISCUSSION OF DYKE GEOCHEMISTRY |
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Open-system fractional crystallization but at what depth?
The element and oxide data summarized in Figs 3 and 4 make a strong qualitative case for fractional crystallization as the mechanism that produced the wide range of rock-types in the Spitzkoppe dykes, and the isotopic data suggest that this was an open-system (AFC) process. The composition of the parental magma is difficult to determine. All the most Mg-rich dykes in the suite are rich in olivine phenocrysts. Electron probe microanalyses of these in two of the picrites (Fig. 2) has shown that this olivine is not unusually forsteritic and therefore the high MgO of the rocks is clearly due to olivine accumulation. The most direct way to identify a likely parental composition is therefore to focus on the aphyric dyke with the highest MgO content. This is VB26 (Table 2), with 11·1% MgO.
Judging from Fig. 3, such a parental melt would have accumulated olivine phenocrysts to form the picrite suite and then evolved to the basalts, intermediate compositions and rhyolites by fractional crystallization of the following assemblages: olivine alone; ol + cpx; ol + cpx + plag; plag + cpx (ol disappears from thin sections at 5·0–5·5 wt % MgO); plag + cpx + Fe–Ti oxide; plag + cpx + oxide + apatite. The main phenocrysts observed in the rhyolites are alkali feldspar and quartz. This chemistry-based story of fractional crystallization is complicated by two items of evidence: (1) all the basalts that appear, from geochemical considerations, to have fractionated Ca-rich clinopyroxene lack phenocrysts of this mineral; (2) Sr, Nd, Hf and Pb isotope variation in the suite all imply that the fractional crystallization was open-system, with concomitant dissolution of crust in the melts.
The relative stabilities of olivine and Ca-rich clinopyroxene crystallizing from a basaltic magma are very sensitive to pressure. The most direct way to show this is by using experimental melting studies of rocks to determine the crystallization behaviour of a basic–intermediate magmatic suite as a function of pressure. After noting the ‘cryptic clinopyroxene fractionation’ phenomenon during 1 atm phase-equilibria studies of basic and associated lavas (Thompson, 1972), Thompson (1974) studied the Palaeocene alkali olivine basalts and hawaiites of Skye, NW Scotland, using polybaric experimental phase equilibria and showed that they had undergone extensive ‘cryptic’ fractionation of Ca-rich clinopyroxene (aluminous sub-calcic augite in that case) at pressures around 0·9 GPa, despite the lavas being almost devoid of augite phenocrysts in thin section. The effect of pressure on the relative stabilities of liquidus olivine and Ca-rich clinopyroxene in basaltic melts and their fractionation residua can be seen clearly in a plot of CIPW-normative ol–di–hy–ne-Q (Fig. 8). Although this projection ignores plagioclase, it is useful for illustrating the crystallization behaviour of olivine tholeiite suites such as the Spitzkoppe dykes. Parts of several cotectics are marked in Fig. 8a. They are all for the equilibrium assemblage (ol + plag + cpx + basaltic melt). The 1 atm curve was compiled by Thompson et al. (2001) from several sources. The 0·9 GPa curve was constructed by Thompson (1982) from the results of experimental melting P–T studies on olivine tholeiites, alkali basalts and hawaiites that precipitated ol, plag and cpx cotectically on their liquidi at 0·9 GPa. The 0·2 and 0·8 GPa curves are from melting experiments on mid-ocean ridge basalt (MORB) by Grove et al. (1992), analysing progressive residual liquids. The 0·7 GPa curve is for experimental fractional crystallization of a tholeiite (Villiger et al., 2007). A single point, marked M in Fig. 8a, is available for a Columbia River Plateau basalt (Martindale, from Helz, 1980) that shows experimental cotectic liquidus precipitation of ol + plag + cpx at 0·35 GPa, a mid-crust pressure.
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All of these cotectics show consistent migration away from the Cpx apex of Fig. 8a with rising pressure. This means that a fractionating olivine tholeiite magmatic system could undergo cotectic separation of olivine and clinopyroxene in the deep crust, but any remaining clinopyroxene phenocrysts might redissolve during the subsequent uprise of the melts, so that chemical study of the suite would indicate ‘cryptic’ clinopyroxene fractionation. Thompson (1987
) showed how the olivine tholeiite lavas of Kilauea, Hawaii, behave on this diagram just as would be expected for a magmatic suite fractionating within the upper crust, and therefore controlled by the 1 atm cotectic. Likewise, the 132 Ma Horingbaai dyke swarms, immediately NW of Spitzkoppe (Thompson et al., 2001) conform closely on this diagram (Fig. 8b) to the expected behaviour of olivine tholeiite magmas fractionating within the upper crust. In contrast, the Spitzkoppe dykes behave very differently (Fig. 8a). Picrites and basalts with >9 wt % MgO form a fan-shaped array, as would be appropriate for compositions controlled by accumulation or removal of olivine. The trend for dykes with lower MgO contents is scattered. Although MgO generally falls as the Spitzkoppe dyke compositions migrate towards the Qz apex of the diagram, this relationship is not a close one. The non-picritic dyke analyses give the overall impression in Fig. 8a that they lie scattered along a cotectic (or bundle of cotectics) at similar depths to, or perhaps a little deeper than, the one at 0·35 GPa (14 km depth) marked by the Martindale basalt liquidus phase equilibria (Helz, 1980). This places the likely site or sites for the evolution of the Spitzkoppe dykes within the lower crust. Furthermore, they appear to have evolved at depths less than the 30–35 km of the Moho in this area (Green, 1983; Miller, 1983; Gladczenco et al., 1997; Bauer et al., 2000) because the compositions do not trend towards alkalic residua, as has been shown experimentally to occur at 0·8–0·9 GPa (Thompson, 1982; Grove et al., 1992). Although some of the natural cotectics for Spitzkoppe dykes seem to be further from the Cpx apex of Fig. 8a than predicted by the closed-system fractional crystallization experimental results of Villiger et al. (2007), modelling using pMELTS indicates that this phenomenon appears to be caused by AFC processes involving granitic (sensu lato) crust (Fig. 8b and see below). The lavas of the Snake River Plain, Idaho, behave in the same way (Thompson et al., 1983, fig. 1b).
When the radiogenic isotope data are viewed in this light, the trends of the Spitzkoppe dyke analyses in Figs 5–7
can be interpreted with more confidence. Relatively low 143Nd/144Nd (negative Nd), 87Sr/86Sr and 206Pb/204Pb are all characteristic of the high-grade metamorphic rocks that are thought to form the lower half of most continental crust (Moorbath et al., 1975; Dickin, 1981; Scherer et al., 1997; Rudnick & Gao, 2003; Dickin, 2005). The trends for picrite and some of the basalt data in Figs 5 and 7 are therefore clearly compatible with AFC processes at lower-crust depths in the Spitzkoppe dyke suite, as will be discussed further below.
Rhyolite melt evolution in the upper crust?
When their elemental compositions alone are considered, the rhyolite dykes in the Spitzkoppe swarm appear to be part of the same magmatic suite as their picrite–intermediate companions (Figs 3 and 4). The plot of Nd–Hf isotope ratios shows the same relationship (Fig. 6) but the situation becomes subtly different when Sr isotopes are considered. The rhyolite dykes [Table 3 and Frindt et al. (2004b)] have considerably higher 87Sr/86Sr ratios (Fig. 5) than can be fitted by any plausible mixing curve passing through the data for the picritic–basalt dykes. Instead, the rhyolite dykes have 87Sr/86Sr ratios in the same range as the metasediments and extensive S-type granites of the Damara belt (McDermott et al., 1989, 1996; McDermott & Hawkesworth, 1990; Jung et al., 2003), the regional rock-types into which they are emplaced. Lead isotopes (Fig. 7) show a subtly different relationship. Both the rhyolitic and two of the accompanying Mg-poor basaltic dykes lie within the compositional fields of the Damara S-type granites and metasediments. Figures 5 and 7 also show how relatively high 87Sr/86Sr and 208Pb/204Pb separate the Spitzkoppe rhyolite dyke compositions from those of the granites in the penecontemporaneous Gross Spitzkoppe stock, around which they are emplaced. The simplest way to interpret these isotopic relationships is to deduce that the picritic–magnesian basalt magmas of the Spitzkoppe swarm evolved within the lower or middle crust and the rhyolite melts developed by extreme AFC at upper-crust levels. Nevertheless, localized wholesale upper-crustal melting cannot be ruled out. The Mg-poor and intermediate dyke magmas followed a complicated series of individual paths through the sub-Spitzkoppe continental crust.
Modelling the fractional crystallization
Open-system fractional crystallization in the Spitzkoppe magma system can be confirmed by means of modelling calculations using MELTS, pMELTS and Adiabat_1ph (Smith & Asimow, 2005), as detailed in the Appendix. Apart from traces of interstitial biotite (Table 1) in the thickest picrite dykes—with the volatiles probably inherited from traces of hydrous minerals in the SCLM (see below)—the mineralogy of all the Spitzkoppe dykes is anhydrous. The only exception is a trace of interstitial blue–green amphibole in rhyolite VB30. It therefore seems a reasonable approximation to treat the Spitzkoppe magma system as anhydrous until further research clarifies its water content.
Method. All the picrite dykes are olivine-phyric and limited microprobe study suggests that the olivine is accumulative (Fig. 2). Attempts to model fractional crystallization using the original version of MELTS, however, predicted extensive liquidus orthopyroxene. It is known (Table 1) that at least picrites VB18 and 32 contain small amounts of interstitial orthopyroxene; their bulk compositions are thus not far from saturation with this phase. The original MELTS algorithm overestimated the relative stability of orthopyroxene compared with olivine (Hirschmann et al., 1998) and it is perhaps unsurprising that this systematic bias is enough to stabilize the ‘wrong’ phase. Although it is intended for calculations at >1 GPa, and for peridotite compositions only, pMELTS sometimes captures the behaviour of MgO-rich basaltic melts (sensu lato) more closely than MELTS (e.g. Smith et al., 2003). Hence, we repeated the calculations using pMELTS, which gave olivine as the liquidus phase. We stress that because (with the exception of the 1 GPa result) the calculated trends lie outside the calibrated range of the pMELTS model, the quantitative results should be taken with some caution.
The most MgO-rich aphyric dyke sample in this suite is VB26 (11· 09 wt % MgO) although, as discussed in the ‘Focus on the picrites’ section below, the real parental melts at Spitzkoppe may have had as much as 15–20 wt % MgO. The usual way in which algorithms such as (p)MELTS are employed to model fractional crystallization is to select a putative parental composition from the samples in the magmatic suite, such as VB26, and then to use the software to predict how this might undergo fractional crystallization, or AFC if required, under various P–T regimes. This approach means that elemental abundances in the calculated evolving melt are forced to be the observed values and any mismatch between parameterized and actual phase compositions must be taken up by the solid phases only, which sometimes causes the predicted melt to evolve to extreme compositions. In contrast, when (p)MELTS is used for melting calculations, the bulk composition is imposed and any mismatches between actual and predicted phase compositions are distributed over all the phases present. Consequently, rather than using the natural dyke chemical analysis of VB26 as the starting composition for fractional crystallization modelling, it was preferable to use pMELTS to generate a picrite ‘model parental melt’, which (according to pMELTS) could then fractionate olivine alone until it closely resembled VB26.
Extensive trial-and-error mantle fusion calculations with pMELTS produced a suitable ‘model primary melt’. This aggregated fractional melt, with 20 wt % MgO, was generated by isentropic decompression melting under a relatively thin lithospheric lid. For Os-isotope reasons explained below, depleted mantle compositions such as those from McKenzie & O’Nions (1991, 1995) and Workman & Hart (2005) were tried and the best fit was achieved by using the former with the potential temperature (Tp) of the mantle set to 1500°C. The ‘model parental melt’ was derived by subtracting olivine fractionally from the ‘model primary melt’ until MgO reached 15 wt % (the lower limit of the estimated primary magma range, with a calculated equilibrium olivine of Fo90). The actual trace-element abundances and isotopic ratios of VB19 were attached to ‘model VB19’ (Table 2). Of the three picrite dykes with similar MgO contents, VB19 was chosen for its wide range of isotope determinations and because its trace-element abundances were close to the average values of the other two (i.e. VB3 and 18).
It is conceivable that parental magmas were transported laterally as dykes from a location near the axis of the Proto-Atlantic rift, where the lithosphere would have been thinner (e.g. Bauer et al., 2000), but obviously the pMELTS starting model cannot be used to discuss in detail the ultimate origin of the Spitzkoppe picrites. It is simply a device to let the pMELTS algorithm interact with a ‘starting composition’ that it ‘recognizes’ as having been originally in equilibrium with a lherzolite mantle.
AFC
Bowen (1928) concluded that, because upwelling magma is generally porphyritic, most magmas are saturated and carry little or no superheat. Therefore heat for dissolving country rock comes from that released during crystallization of the melt. As his detailed phase-diagram approach demonstrated, however, the reaction progression depends on the complicated interactions of temperature-dependent phase stability fields. Hence, AFC is ideally suited for analysis with a thermodynamically based model, such as MELTS, and we have applied the Adiabat_1ph front-end here also.
For the Spitzkoppe suite it has been necessary to modify the procedure in several ways and new features were added to the public version of Adiabat_1ph in the process; details are given at www.gps.caltech.edu/~asimow/adiabat and in the Appendix. This included an implementation of isenthalpic calculations together with a means to specify the thermal state of the assimilant. The initial crustal temperature was taken as 500°C after considering Cretaceous palaeogeothermal gradients derived from lherzolite mantle xenolith P–T equilibria from Okenyenya (Baumgartner et al., 2000) and near Swakopmund (Whitehead et al., 2002), the two nearest appropriate localities. The starting situation was 100 g of ‘model VB19’ at its liquidus, and fifty 1 g increments of either VB12 or PA19 acid crust (see Fig. 9 legend) were added.
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Reiners et al. (1995
) incorporated the original software of Ghiorso & Sack (1995) into a quantitative model of heat-balanced AFC and showed that the ratio of the rates of assimilation and fractional crystallization, r, could vary dramatically as a typical magma evolved. Naturally, cooling and magma transport rates can strongly affect the r value (e.g. Kuritani et al., 2005) and very complex models of magma chamber processes, such as the EC-E'RA
FC model of Spera & Bohrson (2004), can be developed if all likely factors are included. Here a relatively simple equilibrium model was used and the calculations were undertaken in two stages, which represented extreme cases. First, the magmatism was treated as closed-system fractional crystallization of a cooling dry parental magma. Second, the thermal evolution during AFC was constrained by balancing the heat released during fractional crystallization with that required to melt and incorporate plausible country rock. This latter case is approximately equivalent to the ‘maximal contamination’ end-member of Spera & Bohrson (2004) and assumes no heat loss by conduction.
Results. Using the procedure described above, pMELTS calculated the closed-system fractional crystallization trends shown in Fig. 9. Examples at pressures between 0·5 and 1· 0 GPa are plotted and compared with the Spitzkoppe dyke data. The variation in the calculated MgO–CaO inflexion with pressure is relatively small and thus the inferred pressure is only loosely constrained, especially given our caveat about extrapolating pMELTS. It is encouraging, however, that the bold continuous line corresponding to a pressure of 0·6 GPa, which is in the middle of the range inferred qualitatively from Fig. 8 (0·35–0·8 GPa), falls closest to most of the Spitzkoppe data for CaO and Al2O3. At higher anhydrous pressures, spinel precipitated in the pMELTS models (see Fig. 9) and this phase is not observed in any Spitzkoppe dykes.
For comparison, the model results for adding granitic crust VB12 during AFC of ‘model VB19’ at 0·6 GPa are shown (dashed line) in Fig. 9. Results for adding Paresis rhyolite PA19 (Mingram et al., 2000), the proxy for lower-crust partial melt used in this paper, are almost identical. For most oxides the data fall between the closed-system and open-system curves, suggesting that AFC processes dominated the dyke evolution but that the r value was lowered by conductive heat loss. This is particularly clear in the FeO* plot, which shows that Fe–Ti oxides reached saturation below 7 wt % MgO in the dyke magma and AFC models but not in the closed-system models, in agreement with the experiments of Villiger et al. (2007). With the exception of Al2O3, the fit to all oxides (including P2O5 and TiO2, which are not shown) is good and reinforces our confidence in the ability of pMELTS to describe Mg-rich basalt evolution. We tested the possibility that small amounts of water in the parental melts—such as might be held in the minerals of nominally anhydrous mantle (Hauri et al., 2006)—was the reason for the poor Al2O3 fit by re-running the 0·6 GPa calculations using pHMELTS and H2O contents up to 1 wt %. The misfit for Al2O3 became worse and therefore remains unexplained. This is clearly a limitation to the complete success of the modelling.
The success of pMELTS in modelling the isotopic consequences of AFC is clear in Fig. 5. The model curves are marked with increments of added crust in grams (see the Appendix). This is close to, but not identical to, added crust marked as percentages. Starting from ‘model VB19’, the AFC curve in Fig. 5 involving Paresis rhyolite PA19 (Mingram et al., 2000) is consistent with the concept that fractionation from parental to less Mg-rich Spitzkoppe picrites also involved AFC interaction with melts from ancient granulite-facies lower crust below Spitzkoppe. The AFC curve involving local Damara belt granite VB12 (Table 2) passes through the Sr–Nd isotopic ratios of Spitzkoppe basaltic and intermediate dykes. Its curvature is very similar to that of the simple mixing curve in Fig. 5 described previously but this is made harder to see by the two curves having different origins—VB32 for the mixing curve and ‘model VB19’ for the AFC curves. The advantage of using the pMELTS algorithm over the simple mixing calculation is that a single model can simultaneously describe the variations in major and trace element abundances (see Fig. 16 and discussion below) as well as isotopic variations.
Magma densities and crustal magmatic plumbing
Figure 10 summarizes how various radiogenic isotopic ratios in the Spitzkoppe dykes behave as functions of MgO. The published isotopic ranges for Damara belt pre-Cretaceous crustal rock-types exposed at the surface are also given for comparison (double-headed grey arrows), together with our new data (Table 3). Where significant sections of the dyke data show good linear trends, these have been fitted by least-squares regressions. This diagram emphasizes how the 206Pb/204Pb data for dyke compositions between 20 and 9 wt % MgO follow trends away from the composition range of the exposed Damara belt crust. The 207Pb/204Pb and 208Pb/204Pb trends are the same but with a lower R2. Although 207Pb/204Pb and 208Pb/204Pb are relatively low in Damara crust samples VB37 (pegmatite alkali feldspar) and VB39 (aplite), it is clear from Fig. 10 that their extremely high 87Sr/86Sr ratios rule these rock-types out as responsible for the Pb isotope trends. Between 9 and 3 wt % MgO the majority of the dykes have similar, or slightly higher, 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb ratios to the basalts with 10 wt % MgO. A minority of this dyke group is isotopically distinctive, with 206Pb/204Pb >19·0, 207Pb/204Pb >15·71 and 208Pb/204Pb >39·2, as in rhyolite VB30. In contrast, the 143Nd/144Nd and 176Hf/177Hf linear trends extend throughout the dyke compositions, from picrite to rhyolite. Finally, the complicated behaviour of 87Sr/86Sr helps to explain the other isotopic data. Between 20 and 10 wt % MgO, the picrites and olivine-basalts show no significant variation in 87Sr/86Sr with falling MgO. Below 10 wt % MgO, 87Sr/86Sr in the dykes rises irregularly to values similar to those in Damara belt upper-crustal rock-types.
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If, as we argued above, the fall in MgO between picrites and rhyolites in the Spitzkoppe dykes is due to AFC fractional crystallization, then the progressive changes in isotopic ratios through the suite must logically be linked to the same process, AFC, with the contaminant varying progressively as MgO falls. At first sight this sounds impossible because the microprobe studies (Fig. 2) show that the picrites are cumulates. How could the process of crystal accumulation have anything to do with changing radiogenic isotopic ratios? A possible link can be found in the densities of the magmas. In Fig. 11 all the dyke compositions, including the picrites, have been treated as liquids of equivalent compositions. This simplification avoids considering the picrites as liquid–solid suspensions. It is difficult to do otherwise because of the problems associated with estimating what proportion of the margins of their olivine phenocrysts crystallized during their uprise and emplacement. The effect of this simplification is to underestimate the pre-emplacement densities of the picrites, relative to the phenocryst-poor more evolved magmas. No crystal sorting during and after dyke emplacement is also assumed, and this is another considerable simplification.
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The melt densities for each dyke in Fig. 11 were calculated from their partial molar volumes, taking no account of pressure and reducing melt temperature in steps from 1250°C in the picrites to 1000°C in the rhyolites. The resulting melt densities fall progressively from picrites to rhyolites. Therefore, in a sense, Fig. 11 may be seen to show systematic changes in the radiogenic isotopic ratios of the Spitzkoppe dykes with the densities of the magmas. This might occur if each dyke tapped from various depths a fractionating magmatic system in which the density of the melts varied vertically. Such a magmatic system would be stable in a crust with similar vertical density variations to the magmas, so that the relative buoyancy of all the melts was similar. Figure 11 shows that this situation could be achieved if, for instance, the sub-Spitzkoppe crust resembles that of the well-studied northern Scotland region, where Proterozoic metasediments overlie an Archaean middle and lower crust, within which the metamorphic grade varies downwards from amphibolite facies to granulite facies.
There is no seismic study of the crust immediately below Spitzkoppe but this region falls between the two detailed seismic transects of Bauer et al. (2000), about 70 km NW of the eastern end of Transect 2. All the onshore part of this transect shows a laterally uniform crustal seismic structure, with P velocities increasing progressively from 6·1 km/s at 6 km to 6·7 km/s at 33 km depth. The calculated equivalent rock densities (Christensen & Mooney, 1995) are shown in Fig. 11. These form a similar array to the MOIST data (Fig. 11) but at slightly higher densities. Nevertheless, the base-of-crust P velocity is below the values corresponding to mafic igneous rocks. In contrast, about 100 km offshore from NW Namibia the upper crust becomes strongly attenuated and, below 11 km depth, P velocity rises from 7·2 to >7·6 km/s and contains seaward-dipping seismic reflectors (Bauer et al., 2000). Cretaceous palaeotemperatures calculated from mineral equilibria in peridotite xenoliths from localities 100 km NE and SW of Spitzkoppe, respectively, are 1000°C for xenoliths derived from the mantle immediately underlying the Moho (Baumgartner et al., 2000; Whitehead et al., 2002). Therefore is very likely that the lower crustal Cretaceous palaeotemperatures below Spitzkoppe were high enough to maintain granulite-facies metamorphism.
Melt densities were also calculated during the magma modelling using pMELTS, described above. The closed-system fractionation and open-system (AFC) results for 0·6 GPa are plotted as continuous and dashed trend-lines in Fig. 11. The model AFC trend clearly fits the dyke data well. If the dyke magmas came from a fractionating system in which the picrites resided in the lower crust and the rhyolites in the upper crust, then the Sr isotope data in Fig. 10 can be understood. Ancient acid–intermediate granulite-facies metamorphic rocks—characteristic of the lower crust in many continental regions (e.g. Rudnick & Gao, 2003)—typically have relatively high Sm/Nd, and low Lu/Hf, Rb/Sr, U/Pb and Th/Pb (Moorbath et al., 1975; Dickin, 1981, 2005; Scherer et al., 1997), thus accounting for the distinctive AFC trends of the Spitzkoppe dyke magmas with >10 wt % MgO. Likewise, the acid magmas match isotopically the Damara belt exposed Proterozoic–Cambrian metamorphic rocks and granites. Between these extremes some of the Spitzkoppe dykes with 3–9 wt % MgO seem to have been reacting with crust that had similar Nd–Hf–Pb isotopic characteristics to the lower crust, but with relatively higher time-integrated Rb/Sr. Amphibolite-facies middle crust would fit this isotopic profile well (Dickin, 1981).
Clearly, it is possible to construct other speculative scenarios where, for instance, the metamorphic grade of the crust varies laterally. The chosen situation is the one that best fits published seismic data. Finally, it is important to stress that neither Fig. 10 nor the other isotopic plots specify the physical paths taken by individual dyke magmas. Thus a dyke with, say, 7 wt % MgO could: (1) be melt that completed AFC evolution in the middle crust, having previously followed a different AFC path as it rose through the lower crust; (2) have a parent melt that rose directly from the mantle to stagnate within the middle crust.
Figure 12 is a schematic illustration intended to summarize and clarify the concepts developed above, together with ideas discussed in the section below, which focuses on the sources of the Spitzkoppe picrites. The structure of the sub-Spitzkoppe crust is extrapolated from extensive seismic and gravity sections passing north and south of the area (Bauer et al., 2000). These are concordant with previous seismic studies (Green, 1983; Miller, 1983; Gladczenko et al., 1997). An unusual feature of this seismic structure is the lack of any significant amounts of mafic lower crust (Vp >7·0 km/s), as emphasized above. This means that it is reasonable to suppose that intermediate and possibly some acid granulite-facies metamorphic rocks (of either igneous or sedimentary parentage) dominate the lower crust below Spitzkoppe.
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The mafic dykes may originally have been emplaced laterally—possibly from the zone of thinned lithosphere marking the eventual South Atlantic rift (Gladczenko et al., 1997
; Bauer et al., 2000; Thompson & Gibson, 2000). Nevertheless, the combined volume of the Gross and Klein Spitzkoppe granite stocks suggests strongly that a substantial local magmatic system developed below Spitzkoppe. This system could either have fractionated by AFC directly to the magmas that produced the granite stocks or provided the heat that generated them by upper-crust melting. The oxygen isotope evidence of Harris (1995) and Trumbull et al. (2004a) favours the latter model for the granites but the differences in Sr, Nd and Pb isotopic ratios between the rhyolite dykes and the Gross Spitzkoppe granite favours the view that even the acid dykes around Spitzkoppe have a different origin from the granites, as extreme AFC products.
Elsewhere beneath the Etendeka igneous province, such local systems may have developed very large intracrustal magma chambers—sufficient to fuel the eruption of the enormous quartz latites (e.g. Milner et al., 1995; Ewart et al., 2004b). Around Spitzkoppe the remarkably complete picrite–rhyolite suite of dykes, each with their distinctive geochemistry, argues against the homogenizing effects of a large magma reservoir. The preferred model is a stack of sills and their interconnecting dykes and/or pipes, allowing an overall magmatic evolution column to be tapped from a range of depths throughout the crust, as proposed for magmatism elsewhere (Thompson et al., 1972; Morrison et al., 1985; Marsh, 2004; Thompson et al., 2005). It should be noted that the Spitzkoppe picrites form by far the thickest dykes. Presumably exceptionally intense extension events were required to tap such dense and deep magmas. Finally, the summaries of isotopic characteristics at each level in Fig. 12 are, of course, not directly observed but deduced from reasoning outlined above.
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FOCUS ON THE PICRITES |
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In all the diagrams above where geochemical data are plotted the picrites, notably sample VB32, form the mafic end-members to the Spitzkoppe dyke suite. Nevertheless, it is clear from Fig. 2 that the high MgO content of the picrites is not an indication that they were emplaced as exceptionally Mg-rich liquids. Instead, their high MgO contents reflect accumulation of olivine phenocrysts in liquids that, judging from the compositions of the aphyric members of the suite and the calculated equilibrium liquids, did not exceed
11 wt % MgO. This contrasts with the picrites in the penecontemporaneous dyke swarms around Brandberg and the Huab area to the NW (Fig. 1, inset), which frequently contain olivine phenocrysts more forsteritic than Fo91 (Fig. 2), rising exceptionally to Fo93·3 (Thompson & Gibson, 2000; Thompson et al., 2001). The sampling of Spitzkoppe picrites to date (Table 1) is insufficient to be certain that equally forsteritic olivine phenocrysts do not occur. Alternatively, the magmatic plumbing of the Spitzkoppe system may have prevented extremely Fo-rich olivine reaching the upper crust. It therefore seems prudent to suspect that the parental melts fuelling the Spitzkoppe system probably included liquid compositions with MgO in the 15–20 wt % MgO range, but that these were not emplaced unmodified.
The radiogenic isotopic ratios (crucially excluding Os; see below) of the four picrites studied in detail, VB11, 18, 19 and 32, are appropriate to a convecting mantle source similar to those of OIB (Figs 5–7), with 87Sr/86Sr 0·707, Nd 0, Hf +2·5, 206Pb/204Pb 19·16, 207Pb/204Pb 15·70 and 208Pb/204Pb 38·67 (Table 3). In contrast, the elemental compositions of the samples are very different from either MORB or OIB. These differences are illustrated by a normalized plot of the abundances of a range of elements in the picrites (Fig. 13). In Fig. 13a the contrast between the ‘spiky’ Spitzkoppe patterns and average enriched MORB (E-MORB; Sun & McDonough 1989) is clear. It is also interesting to see how similar the E-MORB pattern is to that of a Horingbaai basalt dyke (97SB29) that appears to have undergone minimal crustal contamination (Thompson et al., 2001). Figure 13b shows that duplicate chemical analyses of samples VB11 and 32, from the same Black Range picrite dyke, vary negligibly. In the right-hand part of the diagram, between Sr and Lu, the inter-element ratios of the picrites are similar to those of normal MORB (N-MORB; Sun & McDonough, 1989), whereas all elements in the left-hand part of Fig. 13b show relative enrichment above MORB values.
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Osmium isotopes in the olivine phenocrysts of the Spitzkoppe picrites provide another way of discovering a feature of their original compositions. This is because the Os contents of olivine phenocrysts in the Spitzkoppe and Horingbaai dykes range from
800 to 8000t Os (D. G. Pearson and R. N. Thompson, unpublished data). These values are 15–150 times higher than average crustal values (Saal et al., 1998) and sufficient to make them very resistant to significant change by the addition of a small amounts of crust. Therefore the Os isotopes of the Spitzkoppe picrites should ‘see through’ any subsequent crustal contamination that may have affected them and retain values reflecting their equilibria with mantle rocks. Using an emplacement age of 132 Ma, the initial 187Os/188Os ratios of the picrite olivines (Table 4) are 0·126248 (VB32) and 0·124262 (VB18). Convincing 187Os/188Os ratios for MORB are notoriously hard to obtain and values for abyssal peridotites are usually quoted instead. Nevertheless, Gannoun et al. (2004) used mineral isochrons to obtain initial 187Os/188Os ratios of 0·1265 and 0·1290 for two Atlantic MORBs, at the lower end of the published range for MORB glasses (Alard et al., 2005). The 187Os/188Os ratios of the Spitzkoppe picrites and their olivine phenocrysts, and the Atlantic MORBs, are all well below the range for Atlantic OIB (Hauri, 2002). They are also slightly above the total range (at 132 Ma) for southern Namibian SCLM xenoliths and well above the average value of 0·11781 at 132 Ma (Pearson et al., 2004).
It has been clear throughout the foregoing discussion that the Black Range picrite (samples VB11 and 32; 19–20 wt % MgO) is isotopically the Spitzkoppe magma furthest removed from all observed and inferred crustal compositions in that region. How can this be if its elemental composition differs from MORB in ways that can be modelled satisfactorily by addition of a few per cent of crust to a MORB picrite melt (Fig. 13)? There is an alternative long-standing and popular view that the differences between these Etendeka-related picrites and MORB-source magmas were caused by the picrite sources being located entirely within veined and metasomatized SCLM (e.g. Marsh et al., 1997, 2001; Trumbull et al., 2004a). There is limited direct evidence of the nature of the SCLM beneath the Damara belt, in the form of suites of SCLM xenoliths sampled by Cretaceous magmatism at Okenyenya and near Swakopmund (Baumgartner et al., 2000; Whitehead et al., 2002). The closer Swakopmund xenoliths lack evidence of modal metasomatism and pyroxenite veining (Whitehead et al., 2002).
On the other hand, indirect evidence for veined and/or metasomatized SCLM beneath the Damara belt is widespread. Alkaline lamprophyric mafic and ultramafic dykes occur in the majority of the Damara belt Cretaceous intrusive complexes: Okenyenya, Erongo, Ondurakorume and Okorusu (Milner & le Roex, 1996; le Roex & Lanyon, 1998; Thompson et al., 2001); Paresis (Gibson et al., 2005); cutting the Spitzkoppe granite (Frindt et al., 2004a); between Spitzkoppe and the coast at Henties Bay (Marsh et al., 1991). Some are spectacular mica-rich mafic ultrapotassic rock-types, with >6 wt % K2O at 8 wt % MgO (Thompson et al., 2001). Others differ little from dolerites petrographically but all have broadly similar and distinctive chemical compositions that can clearly be seen from their normalized multi-element patterns (Fig. 13b), which emphasize their steep REE patterns and relative enrichments in Ba, Rb, U, Th, Nb, Ta, Sr and P. The analysed Spitzkoppe lamprophyre FH129 is distinctive in having a high Pb/Ba ratio—a characteristic shared by the Spitzkoppe picrites (Fig. 13)—which may therefore indicate a localized trace-element feature in their source regions. Although the Spitzkoppe lamprophyre FH129 itself is clearly unsuitable in composition to generate the Spitzkoppe picrites simply by mixing in small amounts with a MORB-like progenitor (Fig. 13b), its compositional similarities to them are enough to suggest that either lamprophyric melts or their SCLM sources were involved in the genesis of the picrites.
The clearest evidence that these Cretaceous lamprophyric rock-types in Namibia originated in veined and metasomatized SCLM sources comes from their close similarity to the Cretaceous lamprophyres and lamproites that occur associated with the Paraná flood basalts throughout southern Brazil and Paraguay (e.g. Gibson et al., 1995a, 1995b, 2005). The exotic Sr, Nd, Pb and Os isotope ratios of the South American examples point clearly to their sources as geologically ancient SCLM veins (Gibson et al., 1995a, 1995b, 2005; Carlson et al., 1996). In contrast, the Sr, Nd and Pb isotopic ratios of the Namibian lamprophyres fall close to those of the most Mg-rich Spitzkoppe picrites VB11 and 32 (Figs 5 and 7). As the Namibian lamprophyres (especially FH129) are relatively rich in Sr, Nd and Pb, only a small input from their SCLM sources would be sufficient to dominate the Sr–Nd-Pb isotope ratios of the Spitzkoppe picrites, even if the latter were originally MORB-like Mg-rich melts from the underlying convecting mantle (see also Figs 14 and 16). Milner & le Roex (1996) and le Roex & Lanyon (1998) emphasized that the isotopic characteristics of the Cretaceous Namibian lamprophyres could best be explained if their veined and/or metasomatized SCLM mantle sources developed only shortly before the alkalic magmatism.
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The Spitzkoppe picrites plot close to and slightly below the
Nd–Hf mantle array in Fig. 6. In contrast, the dashed vectors in this diagram show how the peridotites forming ancient SCLM mostly plot far above the array (Pearson et al., 2003). Such peridotites presumably form a substantial component of the SCLM beneath the granulite-facies sub-Damara lower crust. If so, ancient SCLM peridotites clearly played no significant role in the genesis of the Spitzkoppe picrites. Nevertheless, a minor component in the picrites derived from olivine-poor pyroxenitic SCLM veins is entirely consistent with the data in Fig. 6 because such veins plot scattered along the Mantle Array (Pearson & Nowell, 2004). Additionally, the lamproites and similar mafic potassic magmas that are generally supposed to be largely melts of SCLM mica-pyroxenite veins also plot along the Mantle Array in Fig. 6, extending to very low Nd and Hf ratios (Nowell & Pearson, 2003). Thus, the addition of a small amount of such melt to a MORB-like initial Spitzkoppe magma would affect its composition in the way seen in Fig. 6. This process would also not affect Os-isotope ratios of the Spitzkoppe magma significantly because Os is concentrated in mantle peridotites, rather than their olivine-poor veins (Hauri, 2002). Another scenario would be for the convecting mantle source of the picrites to incorporate wholesale some reactivated Thermal Boundary Layer mantle from the base of the sub-Spitzkoppe SCLM, including its population of veins (Thompson et al., 2005).
Finally, the behaviour of Eu in the Spitzkoppe picrite dykes is consistent with this view of their origin. Figure 14 shows how the Eu anomaly (Eu/Eu*, where Eu* is the average of Sm and Gd abundances; all normalized to chondritic meteorites) can be determined very accurately and remains extremely close to 1· 0 in average MORB and OIB (Sun & McDonough, 1989) and in typical continental small-scale suites of basalts, such as the Holocene Potrillo lavas, USA (Thompson et al., 2005). In contrast, virtually all the Spitzkoppe dykes—picrites to rhyolites—have small to large negative Eu anomalies. With model picrite VB19 as the starting composition, we used pMELTS to calculate an Eu anomaly curve for AFC involving local granite VB12, terminating when MgO fell to 5 wt %. This shows that a combination of fractional crystallization involving plagioclase separation with AFC can account for much of the Eu anomaly variation in this suite because likely crustal contaminants, such as granite, all have large negative Eu anomalies inherited from the processes that formed them. Unless the Spitzkoppe picrites fractionated from even more Mg-rich parents, AFC involving crust cannot be invoked to explain why they also have negative Eu anomalies.
The alternative explanation of this geochemical feature in the Spitzkoppe suite stems from the observation that in Fig. 14 the abundant Cretaceous REE-rich mafic potassic and associated lamprophyric dykes of Namibia, southern Brazil and Paraguay all have similar slightly negative Eu anomalies to VB32, even though they definitely have not precipitated plagioclase. This geochemical feature of the Spitzkoppe picrites could therefore also be derived from a small SCLM component in their source magmas, rather than crustal contamination, although the latter would subsequently progressively increase the negative Eu anomaly in the fractionated dyke melts. A mixing curve in Fig. 14 between picrite VB19 and 96SB30, a Cretaceous Mg-rich potassic dyke from Okenyenya (Thompson et al., 2001), shows that the Spitzkoppe picrite compositions would have been very sensitive to addition of small amounts of such potassic melt. Variable degrees of mantle fusion during picrite genesis would also have generated melts with variable Nb contents and similar low Eu anomalies. This process may have contributed to the range of Nb abundances in Spitzkoppe dykes with only small negative Eu anomalies.
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IMPLICATIONS FOR ETENDEKA IGNEOUS PROVINCE MAGMATISM |
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How are Spitzkoppe and Etendeka magmatism related?
The close spatial proximity of the HB–O dyke swarm to penecontemporaneous Etendeka lava outcrops, the Huab sill complex and Damara belt Cretaceous intrusive complexes (Fig. 1) has inevitably caused most of those who have written about this large igneous province (LIP) to suppose that all these comprise a single magmatic event (e.g. Erlank et al., 1984
; Marsh et al., 1991, 1997, 2001; Duncan et al., 1998; Ewart et al., 1998a; Harris & le Roex, 1998). Trumbull et al. (2004a, p. 23) stated explicitly that: ‘the (HB–O) dykes represent exposed feeders to a now-eroded southern extension of the Etendeka basalt field’.
The Etendeka lavas show considerable geochemical diversity and as a result they have been subdivided in terms of both magma types, with supposed genetic individuality, and lineages of compositions linked by fractional crystallization (e.g. Erlank et al., 1984; Peate, 1997; Ewart et al., 1998a, 2004a; Marsh et al., 2001). The Etendeka tholeiitic basalt and andesite lavas appear to lack exposed comagmatic picrites; the picrite lavas at their base are distinctive geochemically and cannot be directly related to them (Ewart et al., 1998a; Gibson et al., 2000; Tuff et al., 2005). The dominant southern Etendeka basalt—andesite lavas form a chemically defined group variously known as Tafelberg (Erlank et al., 1984; Marsh et al., 2001) or ‘LTZ-L’ (Ewart et al., 1998a, 2004a)—equivalent to the Paraná Gramado lava group (Peate, 1997). A few geochemically similar basic dykes are found in the swarm west of Brandberg (Thompson et al., 2001). Figure 3 shows that TiO2 abundances in the non-picritic Spitzkoppe dykes are the same as in southern Etendeka low-Ti (Tafelberg/LTZ-L) lavas at equivalent MgO contents (Marsh et al., 2001). Likewise, Fig. 8a shows that the CIPW norms of the main southern Etendeka Tafelberg lava group plot in a field amongst the more fractionated Spitzkoppe dykes, whereas the associated more mafic Tafelberg/LTZ-L dykes (Ewart et al., 1998a; Thompson et al., 2001) show the same distinctive trend as the Spitzkoppe basalt–intermediate dykes in this diagram.
Sr–Nd–Pb initial isotope ratios show comparable similarities between the Spitzkoppe dykes and Etendeka low-Ti lavas and associated dykes. On an initial Nd vs 87Sr/86Sr plot (Fig. 5), the field of the latter is bisected by both the variation trend for less Mg-rich Spitzkoppe dykes and our model AFC curve, using pMELTS with Damara granite VB12 as the contaminant. The associated acid lavas of the Etendeka igneous province, represented by the Goboboseb quartz latites (Ewart et al., 1998b), have Sr–Nd initial isotopic ratios close to those of the Spitzkoppe rhyolite dykes (Fig. 5). Using only post-1985 Pb isotopic data, the Etendeka low-Ti basic lavas and dykes (Fig. 7) define a field remarkably like that of the Spitzkoppe picrite–basalt dykes but extending to lower 206Pb/204Pb ratios. A few Tafelberg and northern Etendeka Khumib–Kuidas lavas with the combination of relatively low 206Pb/204Pb (< 18·0), 87Sr/86Sr (< 0·708) and 143Nd/144Nd (< 0·5124) were claimed by Ewart et al. (2004a, p. 85) to show Pb isotope variations: ‘not consistent with crustal contamination’. Their Pb isotope trends are shown as fields labelled ‘Southern LTZ lavas and dykes’ in Fig. 7. It is clear that, although separated by 100–200 km, they are collinear with the trend defined by the Spitzkoppe picrites and Mg-rich basalts. The remarkable regular fall in MgO in the dykes along the latter trend leaves little doubt that it is AFC-related (see also Fig. 10, where R2 for 206Pb/204Pb vs MgO for such dykes is 0·97). Is it, therefore, really possible to be so certain that the parallel isotopic trends in some Etendeka lavas have nothing to do with crustal contamination of the magmas?
Erlank et al. (1984) and Ewart et al. (1998a) favoured alternative models in which the range of elemental and isotopic compositions of the Etendeka magmas was explained by: (1) the magmas being sourced, at least in part, from heterogeneous and partly enriched upper mantle areas, including the SCLM; (2) AFC-mixing processes in which upwelling mafic magmas had interacted with continental crust, respectively. However, each group of workers conceded that they could not rule out the other possibility based on the available data. If the intracrustal and sub-crustal processes affecting the Spitzkoppe parental picrite magmas (summarized in Fig. 12) are applicable to the evolution of the Etendeka magmas in general, the results of our present study accord approximately with—and add detail to—those of Ewart et al. (1998a). Probably it should be expected that the very large-scale Paraná–Etendeka magmatic systems would have developed large-volume magmatic systems in the middle–upper crust and that mixing processes would have erased most or all traces of any previous interactions involving lower crust. Likewise, it is more a matter of faith than logic to distinguish between crustal and SCLM inputs into such flood-basalt magmatism using only Sr, Nd, Hf and Pb isotopes.
Early clinopyroxene precipitation and Paraná–Etendeka magma group classification
The evidence above that ‘cryptic’ clinopyroxene crystallized at approximately mid-crustal depths from the Spitzkoppe basic magmas at 9 wt % MgO, and that AFC processes involving various crustal rock-types affected essentially all of these melts, has serious implications for the use of various elemental ratios in classifying these, and hence potentially all, Paraná–Etendeka magmas. Figure 15 shows three plots frequently used in this way (e.g. Erlank et al., 1984; Marsh et al., 1991, 2001; Peate, 1997; Ewart et al., 1998a, 2004a) with members of the Spitzkoppe dyke suite plotted on them and coded for MgO content. The large changes in all these inter-element ratios along the Spitzkoppe dyke trend are obvious. Furthermore, pMELTS–AFC modelling produces exactly the same inter-element ratio changes (Fig. 15b) and shows that this effect would be extremely variable, depending on the precise crustal contaminant. The point made in Fig. 15 would be irrelevant if all researchers kept to the stricture emphasized by Trumbull et al. (2004a, p. 27), that these ratios will only remain constant in closed-system fractional crystallization involving exclusively olivine. That would rule out all magmas in the Spitzkoppe suite with MgO <9 wt % and also all more Mg-rich melts that had interacted with lower crust, leaving virtually none to classify!
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(ThN*LaN), normalized to primitive mantle]. E, Esmeralda lavas; G, Gramado lavas; UC, upper crust. Spitzkoppe dyke data are coded according to their MgO contents. MgO values are marked beside most basalt–picrite points. AFC trends are modelled using pMELTS (see text for details) shown as curved arrows. Curved-outline fields are from Ewart et al. (2004aUnfortunately, analyses of Paraná–Etendeka lavas and dykes with far less than 9 wt % MgO are routinely classified, and their ultimate sources discussed, using diagrams involving Ti, Zr and Y. For instance, Marsh et al. (2001
, Table 1) listed some of the analyses they used in their classification of Etendeka magmas, including Tafelberg (6·8–3·4 wt % MgO), Albin (6·8–4·5 wt % MgO) and Esmeralda (5·7–3·6 wt % MgO). Likewise, Ewart et al. (2004a, fig. 5) plotted Etendeka ‘basalts and dolerites’ with MgO slightly below 2 wt % MgO, and Peate (1997, fig. 7) used the term ‘basalt’ to include Paraná samples with just over 2 wt % MgO.
Can geochemical signals from below the lithosphere be recognized in Paraná–Etendeka magmas?
The most recent account of Etendeka magmatism (Ewart et al., 2004a, 2004b) included extensive discussion of Paraná–Etendeka magmatism in general. Ewart et al. (2004a) concluded that: ‘the emerging Tristan plume entrained (old) subducted material in the Transition Zone region’ of the convecting mantle, and that this is the main cause of the crust-like geochemical imprint in Paraná–Etendeka lavas. The thesis of Ewart et al. (2004a) is an application to continental magmatism of proposals that have been developed to explain geochemical characteristics of oceanic magmatism, especially in the South Atlantic region. Following Hart's (1984) identification of a large-scale isotopic anomaly in Southern Hemisphere mantle (inferred from its magmatic products), many researchers have claimed to identify the geochemical ‘footprint’ in these oceanic magmas of either ancient subducted sediments or delaminated ancient SCLM circulating within the convecting mantle. Gibson et al. (2005) recently summarized the magmatism of the Walvis Ridge, Rio Grande Rise and the islands of Tristan da Cunha, Inaccessible and Gough. They showed that shallow-mantle recycling of the SCLM metasomatized roots of several geochemically distinctive cratons in SW Africa and SE Brazil provided the most convincing explanation of the elemental and isotopic signatures of Cretaceous and subsequent pre-Recent oceanic basalts in the South Atlantic.
Figure 7 includes Pb-isotopic analyses of 79 Ma basalts from Deep Sea Drilling Project (DSDP) hole 525A, Walvis Ridge, erupted close to the Mid-Atlantic Ridge at that time (Gibson et al., 2005, fig. 1). Their 206Pb/204Pb ratios are so low that it is clearly feasible to see much of the isotopic characteristics of Etendeka magmatism as inherited from convecting mantle sources, thus partially agreeing with the Ewart et al. (2004a) interpretation of Etendeka magmatism and earlier work on Paraná (e.g. Peate et al., 1999). But do both these and other similar studies ignore the elephant in the room presented by the compositions of both continental crust and SCLM? The Paraná–Etendeka magmas clearly rose slowly through both these units, allowing extensive fractional crystallization to take place. To simplify the following argument, we shall concentrate only on the elements Nb and Pb. The 187Os/188Os ratios of Spitzkoppe picrite olivines (Table 4) are similar to those of depleted N-MORB. If, as argued above, the Spitzkoppe parental melts at the top of the convecting mantle contained 15–20 wt % MgO, what might their Nb and Pb contents have been? Su & Langmuir (2003) listed the average Nb content of depleted MORB (MgO 8 wt %) as 1·92 ppm. They gave no value for Pb but commented that their averages agreed well with those of Sun & McDonough (1989), who listed average Nb content of N-MORB (MgO 8 wt %) as 2·33 ppm and Pb as 0·30 ppm. The Pb content is probably a maximum because elemental Pb determination at such abundances requires careful acid leaching of sample chips before analysis (N. Wittig & D. G. Pearson, unpublished data) and any global average will probably include some unleached samples. Therefore, allowing for the MgO difference between average MORB and picrite melts, we consider realistic Nb and Pb contents of the Spitzkoppe parental picrites to be 1· 4 and 0·2 ppm, in agreement with the data of Kelley et al. (2005).
Attempting to derive average Nb and Pb contents for SCLM would be meaningless because it is demonstrably very heterogeneous (e.g. xenoliths and orogenic lherzolites) and hence variable in its fusibility. There is general agreement that the Cretaceous mafic potassic magmas of NW Namibia are the products of incipient SCLM fusion affecting its most fusible veins and patches (e.g. Milner & le Roex, 1996; le Roex & Lanyon, 1998; Thompson et al., 2001). Therefore these are the best exposed representatives of potential contaminants of rising Spitzkoppe picritic melts. Average Nb and Pb contents for such rock-types (Fig. 13) are 150 and 7 ppm, respectively.
We argued above that the crust below Spitzkoppe lacks a mafic lower crust and, judging from Cretaceous palaeogeothermal gradients, was metamorphosed to granulite facies during Etendeka magmatism and had probably been so since the Archaean. The fusible parts of the exposed Damara Belt crust are dominated by granites (sensu lato) with very variable Nb contents of 3–40 ppm (averaging 17) and Pb contents of 30–40 ppm (McDermott et al., 1996; Jung et al., 2003; this paper, Table 2). Although there are no obvious theoretical reasons why Nb contents should change in fusible rock-types within the middle or lower crust, Dickin (1981) showed that average Pb contents in Lewisian gneisses fall from 12 ppm in amphibolite-facies samples to 5 ppm in granulite-facies samples.
The import of these figures is that the rising Spitzkoppe parental melts from convecting mantle sources will travel through: (1) SCLM melting locally to a contaminant with 100 times the Nb and 35 times the Pb content of the Spitzkoppe magmas at that depth; (2) overlying crust with much lower Nb contents, close to those of the (now fractionated) magmas; (3) lower crust with fusible rock-types similar to SCLM veins in Pb content but followed by higher crustal layers with Pb rising to 200 times the original picrite values. Faced with rising through both SCLM and crust having such huge relative enrichments of both Nb (SCLM) and Pb (SCLM and crust), is it realistic for Peate et al. (1999), Ewart et al. (2004a) and others to suppose that any—let alone most—continental flood basalts rise from their convecting mantle sources, fractionate extensively and reach the surface without substantial modification of their Nb and Pb contents?
Figure 16 is a plot of Sr/Pb vs Nb/Nb* (defined in the legend) used by Ewart et al. (2004a) to argue their case for subducted ancient sediment in the convecting mantle sources of Paraná–Etendeka magmas. In this version the Walvis Ridge DSDP hole 525A basalts, Spitzkoppe dykes, pMELTS modelling and other data are added. If our analysis above of the SCLM Nb budget is correct, the overall SCLM average compositions of McDonough (1990) and Ionov (1996) marked in Fig. 16 are probably irrelevant. The averages for N-MORB and OIB (Sun & McDonough, 1989) and the fields for Tristan da Cunha and DSDP hole 525A lavas show that oceanic magmas in the South Atlantic diverge considerably in Nb/Nb* from typical MORB or OIB but less than most Spitzkoppe dykes. The fields for Cretaceous mafic potassic rock-types show very large variations in Nb/Nb* (0·4–2·6) within both the Paraná and Etendeka regions. If, as argued above, these Nb-rich SCLM melts dominated the Nb contents of picrite magmas rising through them (Gibson et al., 1995a), it is unsurprising that the Spitzkoppe picrites have relatively low Nb/Nb* ratios that show no systematic correlation with MgO content, in contrast to the behaviour of other geochemical parameters in the dykes (Figs 5–7 and 10). A mixing curve in Fig. 16 between average N-MORB and Etendeka mafic potassic sample OKU-94-1 (le Roex & Lanyon, 1998) shows that the addition of only 1% of the latter shifts Nb/Nb* of N-MORB into the Spitzkoppe picrite range.
The two AFC curves calculated using pMELTS show that this process affects the plotted parameters extremely variably, depending on the crustal contaminant. The VB12/AFC model curve is similar in length and orientation to the elongation of the Esmeralda, Gramado and Tafelberg lava group fields given by Ewart et al. (2004a). Although the Spitzkoppe data broadly conform to this AFC model curve (Fig. 9), the scatter shown by the picrite data is repeated by the less Mg-rich dykes. This is as would be expected, if individual magma batches pursued individual paths through the Pb-rich continental crust.
The lesson from Spitzkoppe appears to be that such problems may be so complex as to be insoluble with certainty at present, unless there is access to complete extensive comagmatic suites, such as the one around Spitzkoppe, and much more knowledge of the composition of the lower crust throughout the region. This principle has been known for a long time (e.g. Moorbath & Thompson, 1980; Thompson et al., 1982, 1986; Dickin et al., 1984, 1987; Morrison et al., 1985) and is regularly rediscovered and embellished (e.g. Baker et al., 1996; Kent et al., 2002; Yaxley et al., 2004; Harlou et al., 2006).
Acid magmas
The Spitzkoppe rhyolite dykes and granite (Harris, 1995; Frindt et al., 2004a, 2004b; Trumbull et al., 2004b), throw similar light on the genesis of acid and sub-acid magmas in the Paraná–Etendeka LIP. Erlank et al. (1984) considered that the large-scale Etendeka latites and quartz-latites were crustal melts. Subsequent studies have divided between emphasizing the concept that both these erupted rock-types and the substantial acid magmatism associated with the Damara belt Cretaceous intrusive centres were products of extensive AFC interaction between mafic magmas and continental crust (e.g. Ewart et al., 1998b) and the alternative that the observed geochemical range of the acid magmas can be attributed to extensive crustal melting by the Cretaceous mafic magmas and variable mixing between these two contrasting components (e.g. Harris, 1995; Trumbull et al., 2004b). As Harris (1995, p. 318) pointed out: ‘The distinction between assimilation and crustal melting when the former exceeds 60–70% is probably one of semantics’. Our study adds detail to this picture in one case and emphasizes that the complexity in intermediate melt compositions caused by AFC within the lower–middle crust may be effectively disguised by a subsequent upper-crust geochemical ‘overprint’ in the acid residua. This may be the reason why Trumbull et al. (2004b) concluded, on the basis of Nd and O isotope data, that the crust below the Damara belt was relatively uniform laterally on a large scale.
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SUMMARY |
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The
132 Ma dyke swarm in the vicinity of Spitzkoppe, in the Damara Pan-African orogenic belt of western Namibia (Fig. 1), comprises dykes ranging continuously in composition from picrite (20 wt % MgO) to rhyolite. Samples from the thickest (10–20 m) picrite dykes have olivine phenocrysts that appear, from microprobe analyses, to be accumulative (Fig. 2). They also contain both small amounts of orthopyroxene and traces of biotite in their groundmasses (Table 1). Plagioclase phenocrysts join the olivine in less Mg-rich basalt dykes and other phenocryst phases, such as clinopyroxene and Fe–Ti oxides, are confined to intermediate–acid dykes. Nevertheless, CaO and Sc variation in the dyke suite (Fig. 3) suggests that liquidus Ca-rich clinopyroxene precipitated from the melts when their MgO content reached 9 wt %. This phenomenon can occur if fractional crystallization takes place at high pressures, and a plot of CIPW-normative di–hy–ol–ne-Q (Fig. 8) in the Spitzkoppe dykes with MgO >3 wt % suggests that the magmas may have fractionated at mid-crustal depths.
Variations in Sr, Nd, Hf and Pb isotope ratios within the dyke suite (Figs 5–7) show clearly that their fractional crystallization was under open-system conditions (AFC). For each radiogenic isotopic ratio, there is a good to excellent correlation with MgO content in parts of the dyke suite. These correlations indicate the likelihood that various crustal rock-types—some exposed at the present surface and others widely considered to form the sub-Damara-belt lower crust—were assimilated by the Spitzkoppe magmas. When all the measured isotopic ratios are plotted synoptically as functions of MgO (Fig. 10), it becomes clear that the isotopic nature of the inferred crustal contaminant changed systematically as a function of MgO. Thus, isotopic variation in the picrites and Mg-rich basalts is consistent with AFC processes involving a crustal end-member with Sr–Nd–Hf–Pb isotope ratios appropriate for intermediate to acid granulite-facies metamorphic rocks. Such rock-types are thought to be likely to form the lower sub-Spitzkoppe crust (Bauer et al., 2000) and are indirectly ‘sampled’ by the distinctive Sr–Nd isotopic ratios of acid lavas in the nearby early Cretaceous sub-volcanic complex at Paresis (Mingram et al., 2000). Likewise, the radiogenic isotope systematics of most of the basaltic–intermediate dykes ( 9–3 wt % MgO; Table 3) point to an amphibolite-facies metamorphic (middle-crust?) end-member. Finally, the rhyolite dykes and two basalts appear to have interacted predominantly with crust Pb isotopically similar to the Damara belt rock-types exposed at the present surface and forming the local upper crust (e.g. Miller, 1983; McDermott et al., 1989, 1996; McDermott & Hawkesworth, 1990).
The calculated densities of the dyke magmas fall progressively from picrite to rhyolite and this relationship, combined with the geochemical variation, leads to the concept that the sub-Spitzkoppe magmatic system was probably a stacked sequence of interconnected sill-like bodies, extending vertically throughout the sub-Spitzkoppe continental crust (Fig. 11). Episodic extension across the region tapped the magma reservoirs at varying depths, leading to the emplacement of the dyke swarm and also the contemporaneous Spitzkoppe granitic stocks.
The geochemistry of even the most Mg-rich picrites suggests that, although they lack significant crustal contamination, they nevertheless incorporate a small component from a source that was not the convecting upper mantle. Os isotopic ratios in the picrite olivines (Table 4) are at the low end of the MORB range (e.g. Gannoun et al., 2004; Alard et al., 2005). The easiest way to explain many other trace-element and isotopic ratios in the Mg-picrites is to envisage that they originated from the same convecting mantle source as MORB and subsequently, during uprise, interacted with partial melts from the same sort of metasomatized SCLM as formed the source of a suite of Early Cretaceous mafic potassic magmas emplaced sparsely throughout the southern Etendeka igneous province.
Petrogenetic modelling adapted and expanded the Adiabat_1ph version of the MELTS and pMELTS thermodynamically based algorithms (Smith & Asimow, 2005). Thus, the model began with only two input parameters: (1) the composition of the Depleted Mantle source proposed by McKenzie & O’Nions (1991, 1995); (2) the thermodynamic properties of the relevant components and minerals. Extensive trial-and-error calculations using pMELTS found a set of P–T conditions where this mantle generated a picrite melt with 20 wt % MgO, which, after olivine fractionation, would evolve to compositions close to those of the less Mg-rich Spitzkoppe picrites. Further pMELTS fractionation modelling confirmed that mid-crustal pressures (most probably 0·6 GPa; equivalent to 24 km depth) were probably necessary to precipitate Ca-rich clinopyroxene at 9 wt % MgO (Fig. 9). An expansion of the Adiabat_1ph software to model heat-balanced AFC processes can reproduce the essential features of a wide range of elemental and isotopic ratios within the Spitzkoppe dyke suite (e.g. Figs 5 and 14–16).
Finally, the processes deduced and modelled for the Spitzkoppe dyke suite appear also to help explain some geochemical features of the southern Paraná–Etendeka magmatism in general (Figs 15 and 16). The crucial wider message of this research is that elemental ratios widely used for classifying basic igneous rocks, and thought to reflect their mantle sources, become variable if either clinopyroxene fractionation or AFC processes involving continental crust (let alone both) take place. Thus, this study may contribute to the overall understanding of how the magmas of both Paraná–Etendeka and other continental large igneous provinces evolved. We show that the hot upwelling picrite parental melts—generated within the convecting upper mantle beneath such regions—become contaminated during their uprise through both SCLM and continental crust (and their extensive fractional crystallization within the latter) because of the huge disparities between the concentrations of key elements, such as Nb and Pb, in the parental picrites and the more fusible parts of both SCLM and crust.
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SUPPLEMENTARY DATA |
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Supplementary data for this paper are available at Journal of Petrology online.
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APPENDIX |
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Several experimentally calibrated algorithms have been proposed for predicting the crystallization of basaltic and associated melts. The MELTS (Ghiorso & Sack, 1995
), pMELTS (Ghiorso et al., 2002) and pHMELTS (Asimow et al., 2004) algorithms solve for thermodynamic equilibrium between silicate liquid and solid phases. Adiabat_1ph (Smith & Asimow, 2005) is a text-based front-end to the (pH)MELTS family of algorithms—meaning MELTS, pMELTS and pHMELTS—that can simultaneously calculate trace-element and isotopic ratios and was employed here. MELTS and pMELTS may be used with anhydrous or water-saturated systems whereas pHMELTS (Asimow et al., 2004), an extension of the models in which trace amounts of water are partitioned into nominally anhydrous minerals (e.g. olivine), can only be used for melting calculations.
The original release of Adiabat_1ph (Smith & Asimow, 2005) included a ‘source-mix’ function for one-off calculation of a new bulk composition from two sources. A later development was a ‘flux-melting’ mode in which the source region could be enriched at each calculation step. The new AFC functionality is an extension of this mode, in that the thermal as well as chemical effect of the added material is constrained and solid phases may be fractionated simultaneously. Similarly, although Smith & Asimow (2005) described many geologically plausible thermodynamic paths, such as isobaric, isentropic and isochoric, there had not been any user-demand for isenthalpic calculations. Recently we have enabled the isenthalpic mode, so that heat-balanced assimilation in an evolving magma chamber may be modelled, like the calculations described by Reiners et al. (1995). The scheme they used alternated between solution of a simple heat-balance equation for addition of crustal material, which used MELTS output as its input, and an fO2 buffered fractional crystallization step performed with the MELTS software so that the equilibrium phase compositions and relevant thermodynamic quantities were updated for the new bulk composition and P–T conditions. Shortly after the Reiners et al. study, Asimow and Ghiorso completed modifications to the original MELTS model of Ghiorso & Sack (1995) that allowed isentropic or isenthalpic constraints to be imposed directly (Asimow, 1997). Asimow et al. (1995) compared ‘isentropic’ melting calculations approximated using an iterative procedure analogous to that of Reiners et al. (1995) with the isentropic decompression path predicted by the newly adapted MELTS algorithm. Asimow et al. (1995) found that both methods produced identical results and there is no doubt that a repeat of the Reiners et al. (1995) calculations using the isenthalpic MELTS function would do likewise.
The implementation of isenthalpic calculations is essentially that of the GUI version of MELTS, as maintained by M. S. Ghiorso (see http://melts.ofm-research.org/index.html), which is in turn similar to the method used for isentropic calculations (Asimow et al., 1995). At a given pressure, an isothermal calculation involves minimizing the Gibb's free energy of the assemblage, whereas an isentropic calculation requires enthalpy minimization and an isenthalpic one employs entropy maximization. It should be noted that, as no compensation is made for gravity, isenthalpic calculations should really only be performed at constant pressure. Reiners et al. (1995) performed their calculations at equally spaced temperatures, at each stage solving for the amount of assimilant added. Conversely, in both the GUI version of MELTS and the Adiabat_1ph front-end, a user-specified mass of solid (or liquid) is added and the MELTS algorithm returns the new temperature. However, the ways in which the initial conditions for the assimilant are specified differ in the details. Calculation of the initial enthalpy requires knowledge of both the temperature and the phase assemblage. In the GUI version of MELTS the user must list phase and end-member proportions, which will usually be inferred from microprobe analyses and estimates of modal abundance. In Adiabat_1ph the initial state of the assimilant is one that has been saved from a previous simulation for the given bulk composition, run to sub-solidus conditions. The difference in the predicted initial enthalpy calculated by either method is insignificant compared with that related to the uncertainty in the ambient temperature of the surrounding crust.
As currently formulated, it is not possible for (p)MELTS calculations that fix the total enthalpy to be simultaneously fO2 buffered, as this would imply that oxygen entering or leaving the system carried no heat. This problem also exists for calculations along an isentrope or isochore. Instead, we have added an option to approximate fO2-buffering in such situations. The procedure used for each incremental AFC step can therefore be summarized as follows: 1 g of material (equilibrated at 500°C) was added to the system and the composition and total enthalpy were adjusted accordingly. Next, pMELTS was run isenthalpically, with the fO2 unbuffered, to determine the equilibrium outcome of this crustal addition. Then an extra equilibration step re-established the fO2 buffer for the newly constrained P, T conditions and updated the reference enthalpy for any net change in oxygen content. Any resultant solid phases were discarded from the system and the liquid became the starting point for the next increment of added crust. This alternation between closed- and open-system behaviour is similar to that used by Reiners et al. (1995), and is also comparable with the way melt extraction is handled in most geochemical models of adiabatic (‘isentropic’) decompression melting, including Adiabat_1ph.
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ACKNOWLEDGEMENTS |
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The hospitality of the Damara people around Spitzkoppe made our sampling there a pleasure. Nick Marsh did us the substantial favour of analysing the major elements in the samples at Leicester University, and Dave Plant kindly found time on the SX-100 microprobe at Manchester University. Jeremy Preston swiftly designed and produced a bespoke spreadsheet for using partial molar volumes to calculate melt densities. Mark Ghiorso and Paul Asimow generously allowed us to modify their MELTS code. Bruce Charlier facilitated the mass spectrometry at the Open University and provided substantial help and encouragement. The Director and staff of the Geological Survey of Namibia strongly encouraged this study and greatly facilitated its progress. The paper was immensely improved by the thoughtful and constructive comments of Andy Saunders, Chris Harris and an anonymous referee, plus the Herculean efforts of editor Gareth Davies.
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FOOTNOTES |
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Present address: Department of Earth Sciences, The Open University, Walton Hall, Milton Keynes MK7 6AA, UK. 
Present address: Directorate of Diamond Affairs, Ministry of Mines and Energy, 1 Aviation Road, Windhoek, Namibia.
*Corresponding author. e-mail: r.n.thompson{at}durham.ac.uk
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