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Journal of Petrology Volume 41 Number 7 Pages 1041-1056 2000
© Oxford University Press 2000
LIP Reading: Recognizing Oceanic Plateaux in the Geological Record
DEPARTMENT OF GEOLOGY, UNIVERSITY OF LEICESTER, UNIVERSITY ROAD, LEICESTER LE1 7RH, UK
Received September 28, 1999; Revised typescript accepted March 2, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONTHE FORMATION OF OCEANIC...CASE STUDIES OF TWO...LIP READING THE GEOLOGICAL...INCORPORATION OF OCEANIC...CONCLUSIONSREFERENCES |
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Basaltic oceanic plateaux are important features in the geological record. Not only do they record ancient mantle plume activity, but they also are believed to be important building blocks in the formation of the continental crust. In this paper we review the salient features of two Cretaceous oceanic plateaux (the Ontong Java and the Caribbean–Colombian): thick sequences of predominantly homogeneous basalt; the occurrence of high-MgO basalt, including komatiites; and an apparent absence of sheeted dyke complexes. In addition, pyroclastic deposits may be scarce. We then explore ways of distinguishing plateaux from basaltic sequences erupted in different tectonomagmatic settings: continental flood basalt provinces; island arcs; back-arc basins; ocean islands and mid-ocean ridges. Using these criteria, potential Archaean and Proterozoic oceanic plateaux are reviewed and identified. Finally, we explore how these remnant oceanic plateaux became incorporated into the continents, by reviewing the proposed accretion mechanisms for the Cretaceous Caribbean–Colombian oceanic plateau, on the basis of evidence from South America and the tonalites of the southern Caribbean island of Aruba.
KEY WORDS: oceanic plateau; basalt geochemistry; large igneous provinces; plumes
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONTHE FORMATION OF OCEANIC...CASE STUDIES OF TWO...LIP READING THE GEOLOGICAL...INCORPORATION OF OCEANIC...CONCLUSIONSREFERENCES |
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One of the most successful paradigms in attempting to decipher the origin of mafic igneous rocks in the geological record has been the ophiolite model. This was developed in the early 1960s with the formulation of the theory of plate tectonics and the realization of the importance of sea-floor spreading. It was quickly realized that the spatial association of pillow basalts, sheeted dykes, and gabbroic and other ultramafic intrusions represent preserved cross-sections through oceanic crust and upper mantle (e.g. Cann, 1970
), which have been obducted and accreted onto the continental margins. Soon, however, it became evident that oceanic crust of this type formed not only at mid-ocean ridges but also at small spreading centres in back-arc basins associated with volcanic arc systems (e.g. Karig, 1971; Tarney et al., 1977; Watts et al., 1977). The main ways of distinguishing between ophiolites formed in these different tectonic environments have been based mostly on the geochemical signatures or affinities of lava sequences (e.g. Pearce & Cann, 1971, 1973) as well as on the nature of associated rocks. Although tectonic discrimination diagrams based on geochemical composition have been much abused in the geological literature, they are nevertheless a useful aid to deciphering the tectonic setting of volcanic suites, provided that they are used only as part of a framework that is also based solidly on other evidence.
The ophiolite model has thus served the geological community well over the last 30 years. However, like all scientific ideas, it must be modified to take account of new discoveries. One such discovery has been the existence of oceanic plateaux. In the mid-1970s it became apparent that, although most of the ocean crust is 6–7 km thick, there are several regions of the sea floor, e.g. the Ontong Java Plateau (Kroenke, 1974) and the Caribbean plate (Donnelly, 1973), where the ocean crust has a thickness well in excess of 10 km. As first proposed by Kroenke (1974) and later discussed by Burke et al. (1978), Ben-Avraham et al. (1981) and Nur & Ben-Avraham (1982), these oceanic plateaux, because of their thickness (and, if <20 my have elapsed between formation and attempted subduction, their residual heat) are inherently more buoyant than oceanic crust of normal thickness.
This buoyancy means that some of these plateaux have resisted complete subduction and have been accreted onto the margins of continents. They have thus been implicated by many researchers in the growth of the continental crust (Kroenke, 1974; Ben-Avraham et al., 1981; Nur & Ben-Avraham, 1982; Abbott & Mooney, 1995; Saunders et al., 1996; White et al., 1999). As shown by Tejada et al. (1996) and Kerr et al. (1998), it is not only the uppermost basaltic layers of these plateaux that obduct but also, if conditions are favourable, the deeper intrusive sections. The inherent difficulty in subducting thick, buoyant plateau crust means that it is more likely than other oceanic igneous rocks to be accreted onto the continents and preserved in the geological record. Consequently, in the interpretation of accreted igneous terranes within the continents, oceanic plateaux must now also be regarded as a potential tectonic setting alongside back-arc basins, oceanic arcs, continental volcanism, mid-ocean ridges and oceanic islands (‘hotspots’).
The possibility of an oceanic plateau origin for accreted mafic rock sequences is often one that is not considered. Therefore, our aim in this paper is not only to heighten the awareness of oceanic plateaux as potential contributors to the growth of the continental crust, but also to review much of our current understanding of the chemical and geological features of these sequences, to assist in the identification of oceanic plateaux in the geological record.
Our first objective will be to highlight the unique features of oceanic plateaux, that is, those that distinguish them from ophiolites formed in other plate tectonic settings. Second, we will discuss the mechanisms by which oceanic plateaux become incorporated into the continents and how they ‘ripen’ with time into mature continental crust.
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THE FORMATION OF OCEANIC PLATEAUX |
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TOPABSTRACTINTRODUCTIONTHE FORMATION OF OCEANIC...CASE STUDIES OF TWO...LIP READING THE GEOLOGICAL...INCORPORATION OF OCEANIC...CONCLUSIONSREFERENCES |
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One of the single most distinctive features of oceanic plateaux is their crustal thickness. In the case of Iceland (Staples et al., 1997
) and the Ontong Java Plateau (Gladzenko et al., 1997), this exceeds 30 km—approaching the thickness of continental crust. McKenzie & Bickle (1988) have shown that, to produce such crustal thicknesses, an elevated mantle potential temperature (Tp) above that of ambient mantle (1280°C) is required. One way of raising mantle Tp is to invoke a hot plume (up to 200–300°C hotter than ambient mantle) ascending from a thermal discontinuity in the mantle (e.g. Campbell & Griffiths, 1990). Initial decompression melting in the head of one of these plumes as it flattens along the base of the oceanic lithosphere results in the eruption of basalts over an area with a diameter of the order of 1000 km, producing a large igneous province (LIP). Oceanic plateaux, like continental flood basalts (CFB) (Cox, 1980), are a variety of LIP (Coffin & Eldholm, 1994) although, as we shall discuss below, there are some fundamental differences between them.
At this point it is important to mention that the crustal thicknesses of these plateaux vary, both in space and time. Higher mantle temperatures in the Archaean imply that mantle plumes would also have been hotter than at the present day (e.g. Nisbet et al., 1993). This would not only have resulted in hotter and thicker normal oceanic crust, but could have also resulted in even thicker (and so even less subductable) oceanic plateaux.
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CASE STUDIES OF TWO CRETACEOUS OCEANIC PLATEAUX |
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TOPABSTRACTINTRODUCTIONTHE FORMATION OF OCEANIC...CASE STUDIES OF TWO...LIP READING THE GEOLOGICAL...INCORPORATION OF OCEANIC...CONCLUSIONSREFERENCES |
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The Cretaceous period was a time of intense plume-related igneous activity, with the eruption of at least eight LIP, of both continental and oceanic affinity (Coffin & Eldholm, 1994
). Most of this igneous activity occurred between 130 and 70 Ma, with major peaks of activity around 122 Ma and 90–88 Ma (Mahoney et al., 1993a; Tejada et al., 1996; Kerr et al., 1997a; Sinton et al., 1998). The majority of these Cretaceous LIP are oceanic plateaux and include the Kerguelen Plateau in the Indian Ocean; the (now) Caribbean Plateau (formerly in the eastern Pacific); and the Ontong Java Plateau, the Manihiki Plateau, the Shatsky Rise, the Hess Rise and the Mid-Pacific Mountains, in the western Pacific. Although deep-sea drilling [via the Deep Sea Drilling Project (DSDP) and Ocean Drilling Program (ODP)] and remote geophysical surveys (e.g. gravity, seismic and magnetic) of Cretaceous plateaux have been carried out, these techniques really give us only a sketchy understanding of their structures. In particular, recovery of material is restricted to the upper 100 m or so of the basement, <0·3% of the total crustal thickness. However, over the past 5 years, our knowledge of the formation, structure and composition of oceanic plateaux has increased dramatically through detailed studies of the tectonically uplifted margins of two of these Cretaceous plateaux, namely the Caribbean Plateau [here termed the Caribbean–Colombian oceanic plateau (CCOP) because of its outcrop in western Colombia] and the Ontong Java Plateau (OJP). Obduction occurred during attempted plateau subduction and has resulted in the deeper levels of these two plateaux being exposed.
Caribbean–Colombian Oceanic Plateau (CCOP)
It is now widely accepted that the CCOP formed in the eastern Pacific on the Farallon plate during the mid- to late Cretaceous (Burke et al., 1978; Duncan & Hargraves, 1984; Pindell & Barrett, 1990). Less than 5 my after the major formational phase (
88 Ma), the eastward movement of the Farallon plate brought the young, still hot plateau into collision with the proto-Caribbean arc (White et al., 1999) and the NW margin of South America, although the timing of this latter event is not well constrained (Kerr et al., 1997b). This emplacement of the CCOP into the proto-Caribbean region, and its obduction onto the Colombian and Ecuadorian coast, resulted in the uplift and exposure of deep sections of the plateau around the Caribbean and along the north-western edge of South America.
Only a brief review of the CCOP will be given in this paper. All the pre-1996 work on the exposed sections of the CCOP has been summarized by Donnelly et al. (1990) and Kerr et al. (1996c, 1997b). More recent work on the plateau has focused on 40Ar/39Ar dating (Sinton et al., 1998), isotopic characteristics (Walker et al., 1999; Hauff et al., 2000), the structure of the plateau (Kerr et al., 1998) and the gabbros of the province (Révillon et al., 2000).
The CCOP is arguably one of the best exposed and documented oceanic plateaux in the world. Field observations of accreted oceanic plateaux suggest that when they are emplaced onto the continental margins, usually only the uppermost basaltic pillow lavas and shallow-level dolerite intrusions are obducted (Kimura & Ludden, 1995). However, the exposed sections of the CCOP comprise not only pillow basalts and dolerite sheets, but also high-MgO lavas (picrites and komatiites), gabbros and ultramafic rocks (Fig. 1) that have been interpreted as representing the deeper crustal levels of the plateau (Nivia, 1996; Kerr et al., 1998). These assemblages resemble parts of the standard ophiolite model. It has been suggested that the reason for the obduction of these deeper sections is that the CCOP commenced its collision with the Proto-Caribbean arc <5 my after its formation, while it was still relatively hot and buoyant (Kerr et al., 1998). This excess buoyancy may have meant that the CCOP was even less subductable, and so greater uplift occurred during emplacement. Alternatively, the difference may be related to the level at which the detachment from the rest of the plateau occurred (see Kimura & Ludden, 1995).
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The high-MgO picrites and komatiites are more heterogeneous in terms of their incompatible trace element contents and Sr–Nd–Pb isotope ratios than the basalts of the CCOP (Kerr et al., 1996a, 1996b), or the basalts from the OJP (e.g. Babbs, 1997). Examples of this heterogeneity can be seen in Fig. 2, with the high-MgO lavas having Nb contents that range from 2 to 25 ppm, and Lan/Ndn and Smn/Ybn ratios of 0·3–1·4 and 0·5–3·5, respectively (Fig. 3) (Smn/Ybn denotes chondrite-normalized Sm/Yb). In contrast, the basalts of the province possess Nb contents that range from 3 to 7 ppm, and Lan/Ndn and Smn/Ybn ratios of 0·6–1·1 and 0·8–1·4, respectively (Fig. 3). Radiogenic isotope ratios for the high-MgO lavas are also much more variable than for the basalts (Fig. 4), with (for example) the high-MgO lavas having 
Ndi values ranging between +6 and +12, whereas the basalts possess a much smaller range of Ndi values.
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The increased heterogeneity and higher MgO content of the lavas at deeper levels within the plateau are believed to be related. This is because, as noted by Kerr et al. (1995b), the mantle plume source region of oceanic plateaux is heterogeneous. As the magmatic system above a plume evolves, it is likely that large magma chambers will develop, which can ‘trap’ the more primitive, dense and heterogeneous magmas. This entrapment results in mixing and crystal fractionation of the heterogeneous picritic magmas and the eruption of the relatively homogeneous basalts of oceanic plateaux.
One notable feature of the obducted sequences of the CCOP (and indeed the OJP) is the lack of a sheeted dyke complex, which often characterizes ophiolites formed in back-arc basins and at mid-ocean ridges. This may be due to several factors (see, e.g. Saunders et al., 1996). It may simply reflect the tectonic setting of plateaux: if they formed away from a spreading axis, on top of pre-existing crust, then there is no a priori reason for a sheeted dyke complex. Alternatively, a plateau may form at a spreading centre, like Iceland, where any near-surface dyke systems are likely to be buried several kilometres beneath successive accumulations of magma (Pálmason, 1986). The observation that oceanic plateaux do not seem to possess sheeted dyke complexes is a useful aid to identifying these plateaux in the geological record. (However, the identification of the eruption conduits of oceanic plateaux remains a problem.) Finally, the geochemistry of the gabbros of the province clearly shows that they are genetically related to the basalts and picrites (Kerr et al., 1998).
To conclude, we can ask the question: how would we recognize the CCOP if we observed it, preserved, in 3 by time? First, the occurrence of picrites and komatiites should alert us to the possibility of an oceanic plateau. Second, submarine eruptions are indicated by pillow basalts. Third, the chemistry of the basalts [especially the flat, almost chondritic rare earth element (REE) patterns] is diagnostic. Fourth, the lack of substantial tephra layers and sheeted dyke complexes may suggest an oceanic plateau. Finally, the thickness of the extruded portion of the CCOP is at least 7 km (Klaver, 1987), which is much thicker than normal oceanic crust or in most Phanerozoic ophiolites. It may be more difficult, however, to distinguish early Archaean oceanic crust from a Proterozoic or Phanerozoic oceanic plateau. The higher potential temperatures in the Archaean (e.g. Richter, 1988) probably resulted in substantially thicker oceanic crust, perhaps approaching the thickness (tens of kilometres) of recent plateaux. The lithologies may also have been similar. Archaean plateaux, by analogy with modern examples, and if produced above Archean mantle plumes, may well have been substantially thicker than those we see today, and one of several settings for the formation of high-MgO komatiites (e.g. Storey et al., 1991).
Ontong Java Plateau (OJP)
The OJP is the largest known oceanic plateau, covering an area of >1·5 x 106 km2. It consists of two main parts: the western and northern lobe, believed to have formed at 122 Ma, and the smaller eastern lobe, which appears to have formed at 90 Ma (Mahoney et al., 1993b; Tejada et al., 1996; Babbs, 1997; Neal et al., 1997). The total crustal thickness of the OJP is thought to be in excess of 30 km, although estimates from seismic and gravity surveys vary between 25 and 43 km thick (Neal et al., 1997). From 90 Ma, the OJP moved westwards with the Pacific plate. However, at 25 Ma, the westward-moving plateau collided with the Solomon Islands arc, thus clogging the westward-dipping subduction zone (Petterson et al., 1999). This resulted in tectonic emplacement of the OJP on to the Solomon Islands (Petterson et al., 1999) and the uplift and exposure of the deeper parts of the plateau (Mahoney et al., 1993b; Tejada et al., 1996; Babbs, 1997). Continued plate movements eventually resulted in the initiation of an eastward-dipping subduction zone, i.e. the docking of the OJP with the island arc resulted in a reversal in the polarity of subduction from west to east.
The OJP has been tectonically uplifted and exposed only in the Solomon Islands (mostly in Malaita and Santa Isabel), and the rest of our geochemical knowledge of the plateau is based on one DSDP and two ODP drill sites, which penetrated to a maximum depth of 149 m into basaltic crust (Mahoney et al., 1993b). Thus we know significantly less about the structure and composition of the OJP than we do about the CCOP.
Possibly as a result of the poor exposure of the plateau, and because this exposure appears to be restricted, in Malaita at least, to the uppermost 1–2 km of the original plateau sequence (Babbs, 1997), most of the analysed plateau lavas (both pillowed and massive) possess a very uniform composition (Mahoney et al., 1993b; Tejada et al., 1996; Babbs, 1997). Interestingly, all of the recovered sections suggest that eruption of basalt occurred in the submarine environment, with massive flows and pillow basalts being the predominant lithologies. Volcaniclastic rocks are rare, at least in the 122-my-old sequences, and intercalated sediments within the basalt pile are virtually absent. In the previous section, we noted the trace element heterogeneity and the highly variable MgO contents of the lavas of the CCOP. In stark contrast, however, the exposed and drilled lavas of the OJP are entirely basaltic in nature, with >95% of samples possessing a relatively narrow range in MgO content of between 6 and 8·5 wt % (Fig. 1). Furthermore, the samples from the OJP are remarkably homogeneous in terms of their trace element contents (Fig. 2), and like the majority of the basalts from the CCOP, possess relatively flat patterns on chondrite-normalized REE plots, and primordial mantle normalized multi-element plots (Fig. 2). This homogeneity is similarly reflected in the isotopic signatures (Fig. 4), which display a relatively restricted range in values compared with the CCOP. For example, the OJP basalts range in Ndi from +6 to +3 (Fig. 4), and in 206Pb/204Pb from 18·2 to 18·7.
As a result of a detailed petrological study of the basalts of Malaita, Babbs (1997) proposed that the mantle plume source region of the Ontong Java Plateau was, like the source region of the CCOP, heterogeneous and that the relatively homogeneous nature of the basalts was due to mixing either in the melt column or in magma chambers. Although the deeper levels of the OJP are not exposed, several recent reinterpretations of seismic velocity data have resulted in various models of the crustal structure of the plateau being proposed (Farnetani et al., 1996; Gladczenko et al., 1997). It has been suggested that the middle crust consists of olivine gabbros that represent the magma chambers of the plateau. The lower crust of the plateau, having compressional P-wave velocities >7·1 km/s, has been interpreted as being composed of cumulates from the fractionation of the primary picritic melts of the mantle plume (Farnetani et al., 1996; Gladczenko et al., 1997). Furthermore, Gladczenko et al. (1997) proposed that deformation and hydrothermal fluid invasion may have recrystallized the lower-crustal cumulates into garnet granulite, which would have a similar compressional P-wave velocity to olivine cumulates.
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LIP READING THE GEOLOGICAL RECORD: RECOGNIZING OCEANIC PLATEAUX |
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TOPABSTRACTINTRODUCTIONTHE FORMATION OF OCEANIC...CASE STUDIES OF TWO...LIP READING THE GEOLOGICAL...INCORPORATION OF OCEANIC...CONCLUSIONSREFERENCES |
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Although chemical discriminants are useful in the assessment of original tectonic setting of a suite of igneous rocks, they cannot, and must not, be the only means. In this section we will, therefore, focus not only on chemical discriminants but also on geological means of identifying oceanic plateaux. We will look at how to distinguish oceanic plateau sequences from those of volcanic arcs, mid-ocean ridges, marginal (or back-arc) basins, ocean islands, continental flood basalt provinces and seaward-dipping reflector sequences. Table 1 summarizes the chemical and geological features that can be used to help in the identification of oceanic plateaux in the geological record, and these will be amplified in the following sections.
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Oceanic plateaux and volcanic arcs
Arc rocks are relatively easy to distinguish from oceanic plateau sequences on the basis of geochemistry. Condie (1999) has recently reviewed the chemical distinctions between these two types and has suggested several chemical discriminants. One of the most effective is Lapmn/Nbpmn (where Nbpmn is primordial mantle normalized Nb abundance). In Fig. 5a it can be seen that the Lapmn/Nbpmn ratio of Cretaceous oceanic plateaux (CCOP and OJP) is always <1·1, whereas modern and ancient continental and oceanic arcs all possess Lapmn/Nbpmn ratios that are >1·1. Some basalts recovered from the Kerguelen Plateau, especially those from the southernmost parts of the plateau, have slightly higher Lapmn/Nbpmn than do basalts from the OJP or CCOP, possibly because of involvement of subcontinental lithospheric mantle, or even continental crust (e.g. Storey et al., 1989). This diagnostic feature is perhaps one that is best observed on a primordial mantle normalized multi-element plot, where the high Lapmn/Nbpmn ratio of arc rocks is graphically represented as a dip in the relative abundance of Nb (Fig. 5b). In addition to this, arc rocks are generally more evolved and contain more volcaniclastic layers than oceanic plateaux, although subaerial emergence of oceanic plateau volcanoes can result in temporally and spatially localized occurrences of volcaniclastic-dominated sequences.
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Oceanic plateaux and Phanerozoic oceanic crust
On first appearance, the observation that normal thickness ocean crust appears to be readily subductable at the present day suggests that only in exceptional circumstances will it become accreted to the continents and so become preserved in the geological record (e.g. Cloos, 1993
). However, in the Archaean, the Earth’s mantle was substantially hotter than at the present day (Nisbet et al., 1993), and a corollary of this is that more melt was generated at mid-ocean ridges. ‘Normal’ Archaean oceanic crust would therefore have been thicker, and so potentially less subductable than Phanerozoic oceanic crust. Thus, although more recently formed oceanic crust is unlikely to be preserved in the geological record, the more buoyant Precambrian oceanic crust may have been obducted, and so when considering these older sequences this possibility should be borne in mind.
Present-day mid-ocean ridge and oceanic plateau basalt sequences have much in common, and there is a continuum between them. Both types of lava can be pillowed but because of uplift (both thermal and dynamic: e.g. Ito & Clift, 1998) oceanic plateau volcanoes can erupt subaerially, in a manner analogous to Tertiary to Recent plume volcanism on Iceland. Therefore a basaltic sequence in the geological record that preserves evidence of emergent to subaerial eruption (conglomerates, fossil soil horizons, vesicles, accretionary lapilli, tuffaceous horizons, reworking in shallow water) in combination with pillow basalts may well represent an oceanic plateau sequence. However, many of these lithological associations are also found in an arc settings and so it is vital to consider these features in conjunction with geochemistry (see above).
In contrast, eruption at mid-ocean ridges will occur in deep water, and never subaerially, and so any intercalated sediments in such sequences will be deep-water sediments. However, basalts forming the OJP were clearly erupted in a marine, and possibly deep marine environment (deeper than the Aptian calcium carbonate compensation depth), so the absence of shallow water criteria does not preclude an oceanic plateau origin. Furthermore, obducted sequences of Cretaceous oceanic plateaux do not possess sheeted dyke complexes, which are believed to be typical of oceanic crust formed at mid-ocean ridges (Saunders et al., 1996; Kerr et al., 1997b, 1998).
From the above discussion it is clear that the geological discriminants of oceanic plateaux may be ambiguous, especially in ancient terranes. It is for this reason that we must turn to geochemical criteria to distinguish mid-ocean ridge from plateau sequences.
Most Cretaceous oceanic plateau lava sequences possess relatively flat chondrite-normalized REE patterns (Lan/Ndn 1) and can contain picrites and komatiites, whereas most mid-ocean ridge basalt (MORB) suites have depleted light REE (LREE) and rarely contain high-MgO lavas. Although it is possible for plume-derived lavas to be depleted in the LREE (Kerr et al., 1995b; Fitton et al., 1997), LREE-depleted oceanic plateau lavas can generally be recognized by their higher MgO content and/or their higher Nb/Y ratios than MORB (Fig. 5c). As mentioned above, it may be difficult to identify and characterize Archaean oceanic crust, which will have enhanced thickness and an increased amount of high-MgO rocks (as a result of higher mantle source temperatures), and perhaps a less depleted composition.
Oceanic plateaux and back-arc basins
Basalts erupted in back-arc basin (BAB) settings are potentially the most difficult to distinguish from oceanic plateau basalts. Often they possess flat chondrite-normalized REE patterns, and frequently are found as pillow lavas (Saunders & Tarney, 1984). Additionally, like mid-ocean ridges, the small spreading centres related to volcanic arc activity may produce crust of a similar thickness to oceanic crust, and sheeted dykes may also be present. The preservation potential of back-arc basins is high, because they are on the upper plate and hence more likely than normal oceanic crust to be obducted during continent–arc or arc–arc collisions. For this reason, many ophiolite complexes are considered as having formed in a supra-subduction zone setting.
Depending on the nature of the mantle below the back-arc basin, two fundamental compositional types of back-arc basin lavas can form. First, the mantle can be composed of a similar material to the source region of MORB (e.g. Saunders & Tarney, 1991). Although these basalts possess relatively flat chondrite-normalized REE patterns they can be distinguished from oceanic plateaux by virtue of their lower Nb/Y ratios (Fig. 6b). Second, more trace element enriched, plume-derived material can also make up the mantle source region of back-arc basins (Leat et al., 2000) and, as a result, the erupted basalts appear similar in trace element contents to oceanic plateau basalts and plot with the plateaux and intra-plate ocean islands on a Zr/Y–Nb/Y plot (Fig. 6b). We can, however, employ other geological features to distinguish an oceanic plateau basalt sequence from a back-arc basin sequence. First, the proximity of back-arc basins to an explosively erupting island or continental arc volcano increases the likelihood that back-arc basin sequences will contain more abundant tephra layers than oceanic plateaux. Second, the lower temperature (Tp 1280°C) of the mantle below a back-arc basin when compared with a mantle plume (Tp > 1400°C) means that, as at mid-ocean ridges, the eruption of high-MgO lavas is relatively rare. This higher mantle temperature for oceanic plateaux results in the basalts generally possessing higher Ni and Cr contents (at a given MgO content) than basalts formed in a back-arc basin (Fig. 6a). In Fig. 6a, most of the back-arc basin lavas plot in a separate field parallel to the trend displayed by the oceanic plateau lavas. One of the major differences between MORB and BAB basalts is the presence of a supra-subduction zone signature in some BAB basalts, reflected in increased contents of volatiles and large-ion lithophile elements (e.g. Rb, Ba, Th; Saunders & Tarney, 1991). BAB basalts tend, therefore, to have higher Ba/Nb and Th/Nb ratios than unaltered plateau basalts, although this difference may be obliterated during secondary alteration.
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Oceanic plateaux and ocean islands
In some respects, ocean islands are the small-scale equivalents of oceanic plateaux. In general, the magmatic flux rates at ocean islands are substantially less than those estimated for plateaux. There is, however, effectively a continuum between plateaux and islands, controlled by mantle flux rates and the thickness of the lithospheric cap. Magmatism at ocean islands tends to be focused, which, on moving plates, forms chains of islands and seamounts, and aseismic ridges. Many oceanic islands exhibit considerable magmatic differentiation, resulting in formation of rhyolites or phonolitic sequences and their pyroclastic equivalents. Not all do, however; Hawaii, for example, erupts very few silicic differentiates and, conversely, some plateaux also contain some differentiated lavas (e.g. Iceland, and the Kerguelen Plateau: Frey et al., 2000
). Chemical data can be used to distinguish plateaux from ocean islands. Many of the lavas that erupt above hotspots are relatively enriched in incompatible trace elements and depletion of the heavy REE (HREE) suggests deep melting within the garnet stability field (Fig. 6c). Although a few oceanic plateau lavas from the CCOP are enriched in incompatible trace elements and show evidence of melting in the garnet stability field, the vast majority of lavas from oceanic plateaux possess flat REE patterns, indicative of shallow, or, alternatively, high-degree melting. A predominance of lavas, therefore, with an enriched trace element signature (high values of Smn/Ybn, Nb/Zr and La/Y) would suggest that the sequence is part of an oceanic island rather than a plateau.
Oceanic plateaux and continental flood basalt provinces (including volcanic rifted margins)
The lavas of CFB provinces are susceptible to erosion as a consequence of dynamic uplift from the plume and remnant uplift from magmatic underplating (Cox, 1989). Therefore, unless buried by sediment, CFB may not survive in the geological record. What do survive, however, are the dykes and vents through which these lavas erupted, and several Precambrian CFB provinces have been identified on the basis of their remnant dyke swarms (e.g. LeCheminant & Heaman, 1989; Cadman et al., 1994).
In contrast to oceanic plateaux, continental flood basalts are usually not pillowed, and may show extensive evidence of subaerial eruption, in the form of intercalated terrestrial sediments, weathered horizons and initial eruption onto an eroded land surface. Furthermore, because many continental flood basalts erupt through continental lithosphere, they may become contaminated en route to the surface (e.g. Fodor et al., 1985; Thompson et al., 1986; Kerr et al., 1995a; Baker et al., 1996). This contamination results in the enrichment of the large ion lithophile elements leading to elevated Ba/Nb (primordial mantle normalized Ba/Nb > 10).
Mantle plumes have been implicated both as a cause and as a consequence of continental break-up (e.g. Hill, 1991; Anderson et al., 1992; Saunders et al., 1992; Barton & White, 1995). In either case, the associated eruptives form thick lava sequences on the margins of the rifting continents, forming seaward-dipping reflector sequences (SDRS). The SDRS lavas may spill over onto the adjacent continent to form a CFB province. The opening of the North Atlantic at 58 Ma and the break-up of Gondwana to form the Indian Ocean from 120 Ma were both intimately associated with a vigorous mantle plume (White & McKenzie, 1989; Kent, 1991; Kent et al., 1992; Saunders et al., 1997). These lavas commonly display chemical signatures indicative of significant interaction with the continental lithosphere, and so can be distinguished easily from oceanic plateau basalts. However, it is also perfectly possible for an oceanic plateau to merge seamlessly with SDRS and CFB (e.g. Iceland–East Greenland Margin–East Greenland basalts: Larsen & Jacobsdóttir, 1988) as the magmatism straddles the continent–ocean divide.
In places, where fragments of continental lithosphere become isolated in the middle of oceans during break-up, complications can occur. The Nd isotope geochemistry of lavas from the Kerguelen Plateau in the Indian Ocean (negative Nd) implies that they have interacted with ancient lithosphere, and suggests that rafts of this old continental lithosphere occur within this region of the Indian Ocean basin (Frey et al., 2000).
Examples of oceanic plateaux in the geological record
Although Condie (1999) has argued that accreted oceanic plateaux are relatively uncommon in the geological record, we do not accept that this is necessarily the case. As demonstrated in the following section, we believe that oceanic plateaux may, in fact, be widespread in the geological record. Table 2 shows a summary of proposed oceanic plateaux and continental–oceanic margin plateaux from the Archaean to the late Jurassic. Furthermore, we hope that this discussion of the characteristics of oceanic plateaux will allow the identification of other examples throughout the world.
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Archaean greenstone belts in the Baltic Shield, such as the Kostomuksha belt (Puchtel et al., 1998b) and the Olondo belt in the Aldan Shield, Siberia (Puchtel & Zhuravlev, 1993), have been inferred to consist of oceanic plateau sequences. This interpretation is based on the occurrence of uncontaminated pillowed basalts and komatiites with no terrestrial sedimentary intercalations. Furthermore, the chemistry of the lavas (Ndi +2·8 to +3·4; Lapmn/Nbpmn 1) is consistent with an oceanic plateau origin (Puchtel & Zhuravlev, 1993; Puchtel et al., 1998b). Archaean greenstone belts of the Canadian Superior Province have also yielded lava sequences that contain remnants of oceanic plateaux (Desrochers et al., 1993; Kimura et al., 1993; Skulski & Percival, 1996; Fan & Kerrich, 1997; Polat et al., 1998; Hollings & Wyman, 1999; Kerrich et al., 1999). The argument for an oceanic plateau origin is based on the occurrence of komatiites and pillow basalts with no intercalations of terrestrial sediments or sheeted dyke sequences. Chemically, these lavas possess low positive Ndi values, Lapmn/Nbpmn µ 1 and flat to LREE-depleted chondrite-normalized REE patterns, similar to lavas from recent oceanic plateaux. De Wit et al. (1992) suggested that the core of the Kaapvaal Shield in southern Africa, particularly the 3·5 Ga komatiites and pillow basalts of the southern Barberton and Pietersberg Belts, may represent the remnants of one or more oceanic plateaux. This interpretation is supported by the geochemistry of the komatiites, which show no evidence for crustal contamination (Smith & Erlank, 1982).
Several sequences within the Canadian Superior Province contain basalts and komatiites with both the geochemical and geological characteristics of oceanic plateau lavas. However, these are in stratigraphic contact with similar basaltic and komatiitic sequences of lavas that possess a strong signature of continental lithospheric contamination (negative Ndi and Lapmn/Nbpmn > 1: Tomlinson et al., 1998, 1999; Hollings & Wyman, 1999). These sequences are interpreted as continental–oceanic margin plateau sequences, which have erupted in tectonic settings similar to the North Atlantic Tertiary igneous province and portions of the Cretaceous Kerguelen Plateau.
In addition to these Archaean occurrences, examples of accreted Proterozoic oceanic plateau crust have also been discovered. One of the first of these to be identified was in the Birimian province of West Africa (Abouchami et al., 1990; Boher et al., 1992), with pillowed basalts characterized by flat REE patterns and low positive Ndi. Subsequent to this discovery, oceanic plateaux have been identified in the Arabian–Nubian Shield (Stein & Goldstein, 1996) and the Flin Flon Belt in Canada (Stern et al., 1995; Lucas et al., 1996). As in the Archaean, examples of Proterozoic continental–oceanic plateaux have also been found in northeastern Finland (Peltonen et al., 1996) and in other parts of the Baltic Shield (Puchtel et al., 1997, 1998a). Lava sequences in the Proterozoic Cape Smith Fold Belt of northern Québec were proposed by Francis et al. (1983) and Dunphy et al. (1995) to have formed in a continental rifting event. The chemical composition of these basalts is consistent with their derivation from a mantle plume (Dunphy et al., 1995), and we agree that the stratigraphy of the Cape Smith Fold Belt represents the transition from a continental to an oceanic magmatic event.
Basalts from the Wrangellia terrane of western North America have been identified as an accreted oceanic plateau of Triassic age (Lassiter et al., 1995a, 1995b). This interpretation is based on both geological and geochemical evidence. First, the lava sequences contain both subaerial and subaqueous units, with no evidence of nearby arc volcanism or terrigenous sediment. Second, the lavas possess generally flat REE patterns, Ndi +4·8 to +7·3 and Lapmn/Nbpmn 1. Other postulated examples of accreted Cambrian to Jurassic oceanic plateaux are less well studied. However, one area that appears to display considerable promise includes the post-400 Ma accretionary complexes of Japan. It has been suggested that Permian oceanic plateau derived basalts are preserved in the Yakuno ophiolite in SW Japan (Herzig et al., 1994) and parts of the Mino Terrane in central Japan (Jones et al., 1993). Kimura et al. (1994) identified the accreted remnants of a Tithonian (150–145 Ma) plateau in northern Japan that they named the Sorachi Plateau. Possible Permian oceanic plateau-type basalts have also been proposed to exist within the Cache Creek terrane of the Canadian Cordillera (Mihalynuk et al., 1994), although more detailed work on these sequences is required to verify this proposal.
Thus, we cannot concur with the view expressed by Condie (1999) that oceanic plateau sequences are rare in the geological record. The evidence outlined above shows that oceanic plateaux are an important, and arguably a major, component of Archaean greenstone belts, and strongly supports the oceanic plateau model of crustal growth (see Ben-Avraham et al., 1981; Stein & Hofmann, 1994; Abbott & Mooney, 1995; Saunders et al., 1996; White et al., 1999). We believe that there are many more accreted oceanic plateaux still to be discovered, and the onus is therefore on Precambrian workers to bring these ‘out of the closet’. It should be noted that the observation by Bickle et al. (1994) that ‘Archean greenstone belts are not oceanic crust’ may be partially resolved if some of the belts represent plateaux. The absence of a complete ophiolite sequence in greenstone belts would, we argue, reflect the absence of a full sequence in oceanic plateaux.
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INCORPORATION OF OCEANIC PLATEAUX INTO THE CONTINENTS |
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In view of the fact that oceanic plateaux are common in the modern oceans, it may indeed seem surprising that they are not an even more obvious constituent of the accreted collages of the continental crust. A number of factors may explain this paradox.
First, it may be that only a small proportion of plateaux present in the oceans become accreted to the continents. Cloos (1993) has argued that the minimum crustal thickness of an inherently unsubductable oceanic plateau is 17 km. As discussed elsewhere in this paper, buoyancy is the driving force for the obduction, rather than subduction, of an oceanic plateau. This buoyancy derives from the crustal thickness of the plateau and any residual heat from its formation. It is logical that the less buoyant plateaux will be less prone to obduction than their thicker, or hotter, counterparts. The second possibility is that some plateaux are accreted only temporarily to the continents. As suggested by Saunders et al. (1996), the lower parts of an accreted plateau may be converted into eclogite, in which case, subduction could occur ‘piecemeal’ after accretion to a continental margin has taken place.
Our studies of the Caribbean–Colombian oceanic plateau, however, suggest that the major obstacle to identifying an accreted oceanic plateau is the fact that secondary processes may obscure the original plateau characteristics. Whether plateau obduction occurs in an intra-oceanic or continental margin setting, a consequence of the obduction process is that subduction will either reverse in polarity or take a step back, leading to subduction occurring beneath the plateau. Thus, a recently accreted plateau is likely to become obscured by burial with the volcanic products of continuing subduction beneath the region. It may suffer deformation and metamorphism, making original volcanic and sedimentary features difficult to recognize and interpret. Finally, and importantly, intrusion of silicic magmas into the accreted plateau crust ‘matures’ it into something resembling continental crust.
The island of Aruba, now located in the southern Caribbean plate boundary zone, contains a sequence of basalts, dolerites and volcaniclastic rocks belonging to the 88 Ma Caribbean–Colombian oceanic plateau. This sequence is metamorphosed at greenschist (and locally, amphibolite) facies, as a result of the intrusion of a predominantly tonalitic batholith dated at 85–82 Ma (White et al., 1999). The tonalitic magmas have been shown to be derived from partial melting of the mafic plateau material, being triggered by the injection of subduction-related basaltic magmas into the newly obducted plateau crust (White, 1999).
The above example illustrates two important points about the maturing of accreted plateau material, and its transformation into continental crust. First, it demonstrates that the addition of silicic magmas can occur extremely quickly after obduction has taken place. Indeed, it is interesting to speculate that this process may be instrumental in stabilizing (i.e. adding to the buoyancy of) a tract of accreted mafic material, increasing its chances of survival in the geological record.
Second, the generation of tonalites via remelting of mafic rocks in this setting is considerably assisted if the accreted material retains residual heat from its formation. This is the case for the Caribbean–Colombian oceanic plateau, which was obducted only shortly after its formation. Thus, the association of a basaltic sequence with voluminous tonalites (having geochemical signatures consistent with remelting of basaltic rocks) may be a key factor in recognizing a ‘hot’ accretionary setting. This is more likely, logically, to be associated with accretion of an oceanic plateau (which may be still hot when accretion takes place), rather than with, say, an island arc or normal ocean floor.
Thus, even given that the processes of obduction, accretion and continuing subduction can obscure the primary characteristics of an accreted terrane, we argue that the recognition of oceanic plateau successions within the geological record should still be possible. The association of basalts and tonalites may draw attention to an accretionary setting having a high thermal energy; field and geochemical studies should then be able to confirm whether the basaltic sequence has oceanic plateau affinities. This is still possible, in theory, even if the rocks of interest are metamorphosed, as the geochemical discriminants discussed earlier in this paper are based on those elements that remain relatively immobile during metamorphism.
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CONCLUSIONS |
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- Oceanic plateaux are more common in the geological record than has hitherto been realized. On the basis of detailed studies of Cretaceous oceanic plateaux, we propose a series of geochemical and geological discriminants, which should, in most cases, permit the identification of Archaean and Proterozoic oceanic plateau sequences.
- These salient features of oceanic plateaux are as follows: the occurrence of high-MgO lavas (picrites and komatiites); chemically homogeneous basalts with relatively flat chondrite-normalized REE patterns and low Lapmn/Nbpmn ratios (<1·1); pillowed lavas; low abundance of volcaniclastic deposits; lack of sheeted dyke complexes and a relatively thick (5 km) extrusive section. Some plateaux show a predominance of subaerial eruptions, with associated flow morphologies.
- We would strongly caution that none of the above criteria can be used, on its own, to identify unequivocally an oceanic plateau. A suite of rocks of unknown tectonic origin with a significant number of the above features is most likely to be an ancient oceanic plateau. Although this is rather an obvious statement, we feel it is necessary given the misuse of geochemical discrimination diagrams over the past 25 years.
- A problem arises when attempting to extrapolate discrimination analysis to the early Archaean. High mantle temperatures at that time would have resulted in thicker oceanic crust, possibly similar to recent oceanic plateaux. The Archaean equivalents of modern plateaux may have been even thicker structures, with many tens of kilometres of basalts, high-MgO komatiites and plutonic equivalents.
- Finally, we would stress that accreted suites of rocks cannot be considered in isolation. There must also be an appreciation of the large-scale structure of the accreted terrane under investigation.
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ACKNOWLEDGEMENTS |
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There are many people whose assistance, advice and help has been invaluable during the formulation of the ideas contained in this paper; these include John Tarney, Ray Kent, Tanya Babbs, Mike Norry, Giz Marriner, Alvaro Nivia, Godfrey Fitton, Nick Arndt and Dallas Abbot. Constructive reviews by Bob Duncan, John Lassiter and John Ludden are greatly appreciated. A.C.K. acknowledges receipt of a Leverhulme Special Research Fellowship at Leicester, and R.V.W. was supported by the Natural Environment Research Council, UK (GT4/95/197/E).
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FOOTNOTES |
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*Corresponding author. Telephone: +44-116-2523638. Fax: +44-116-2523918. e-mail: ack2{at}leicester.ac.uk
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