Primary magmas at mid-ocean ridges, ‘hot-spots’ and other intraplate settings: constraints on mantle potential temperatures

David H. Green

Research School of Earth Sciences, Australian National University, Canberra.

David.H.Green@utas.edu.au

Most investigations of basalt petrogenesis define primary magmas as those which have formed by partial melting within the upper mantle and thus have chemical compositions (major element, trace element and isotopic signatures) determined by:

1. and 2. combine, in the time-scale of magmatic processes, to produce crystalline phases, a melt phase, and possibly a fluid phase, at equilibrium, i.e. the isotopic signature of the melt represents that of the source and the trace-element, minor-element and major-element compositions of the melt and all crystalline phases represent equilibrium partitioning. This rationale allows us (for example) to use the Fe/Mg partitioning between olivine and melt to screen natural glass compositions, or mantle xenolith-bearing magma compositions, to infer that for the Earth’s mantle with olivine Mg# ~90, primary magmas must have Mg# ~75. Similarly for the Earth’s mantle with modal Ol > Opx > Cpx > Ga, Sp or Plag, a primary magma must have some conditions of P,T, and activities of C-H-O fluid species, where it is multiply saturated with olivine, orthopyroxene and possibly spinel, clinopyroxene, and garnet or plagioclase.

This rationale has led to the two-pronged investigation of basalt petrogenesis by high-pressure experimental methods in which both the progressive melting behaviour of peridotite + (C-H-O) is investigated (defining compositions of residual phases and, under optimal conditions, of the melt phase also) and the liquidus phases of potential primary magmas are determined (e.g., for the suite of mantle xenolith-bearing magmas from alkali olivine basalts to kimberlites). The experimental method is obviously directly applicable to batch melting or melt extraction from upwelling diapirs but it is equally important for concepts of porous melt flow and melt pooling since melt compositions in porous flow are controlled by local compositions and P, T conditions. Melt increments reflect local phase compositions and P,T conditions and pooled melt likewise equilibrates (or attempts to) with wall-rock mineralogy, i.e. a new bulk composition at the P,T conditions of the postulated melt aggregation depth.

Just as the rationale above guides the identification and understanding of the genesis of primary magmas, so it may be used to identify and investigate evolved or derivative magmas formed by cooling and crystal fractionation at low pressure (sub-volcanic magma chamber processes) or at high pressures (mantle wall-rock reaction and precipitation, mantle magma chambers). The latter processes explain examples of evolved nepheline mugearites, and even phonolites in intraplate settings, which contain high-pressure mantle xenoliths. The same data on crystal/liquid equilibrium identify xenocrysts vs. phenocrysts, and evidence for magma mixing in channels or magma chambers.

The melting regime for mantle-derived basaltic magmas from intraplate settings (kimberlites, olivine melilitites, olivine nephelinites, basanites, alkali-olivine basalts) requires the presence of carbon and hydrogen (dissolved (CO₃ = ) and (OH-) in the melt phase. Magma genesis in intraplate settings can be understood in terms of the lherzolite + (C-H-O) system. Magmas are derived from an “incipient melting” regime which lies at temperatures below the volatile-free lherzolite solidus, marking entry to the “major melting” regime. Experimental petrology and observations of natural magmas concur in defining magma segregation at deep levels and magma transport within channels with only limited wall-rock reaction. In the intraplate settings, the role of an “incipient melting” field in lherzolite + (C-H-O) at depths of > 90 km to ?< 200 km creates conditions in which both depletion (from N-MORB to D-MORB sources) and enrichment (from N-MORB to E-MORB sources) occur by migration of a 1-2 % melt fraction.

These relationships are summarized in a specific model for a petrological lithosphere (below the silicate solidus for lherzolite +(C-H-O)) and asthenosphere (above the silicate solidus) (Figure 1) The proposed model is used to explain production of specific primary magmas in the P,T field between the regionally applicable conductive geotherm and adiabatic upwelling of mantle with potential temperature (Tp) ~1430°C (Figure 2).

Figure 1: Application of the experimentally determined melting relationships for pyrolite and pyrolite+(C+H+O) to model the intraplate lithosphere, asthenosphere and subasthenospheric mantle. The ‘oceanic intraplate’ geotherm represents a geothermal gradient appropriate to old oceanic crust and ‘young’ continental crust. This geotherm necessarily traverses the carbonatite stability field and incipient melting regime for the pyrolite compositions provided mantle fO2≥IW+2 log units. Attention is drawn to the petrological base of the lithosphere defined by the high pressure pargasite breakdown and silicate solidus at ~95 km. The lower boundary of the asthenosphere (incipient melting regime along the geotherm) is drawn arbitrarily at ~150 km as the intersection of a mantle adiabat of potential temperature Tp=1450°C with the pyrolite-(C+H+O) solidus for fO₂ = IW+1. Below this depth, a fluid with CH₄ ≥ H₂O is present but there is no silicate melt unless fO₂ > IW+1, i.e. a more oxidised region of the mantle. The carbonatite melt at > 95 km and T < 1000°C is represented as present or absent dependent on local variation in fO₂ i.e. if fO₂ ≥ IW+3 then carbonatite melt ± H₂O-fluid (rather than graphite (diamond)+H₂O-fluid) would be present.

Figure 2: The basis for the view that the modern mantle adiabat has a potential temperature of ~1450°C and all modern volcanism, whether hot-spot, back-arc or intraplate rifting, can be understood within the temperature envelope between the appropriate geotherm for the thermal boundary layer and the mantle adiabat (Tp^a1450°C)

Attention is also drawn to the inference that Archaean peridotitic komatiites imply mantle potential temperatures of at least 1650°C.

Also illustrated is the P,T field inferred for spinel lherzolite, pyroxenite and garnet pyroxenite xenoliths in the Newer Volcanics of southeast Australia. This is considered to represent a perturbed lithosphere geotherm during intraplate, rift-related volcanism and mantle metasomatism in southeast Australia. The field for South African kimberlite xenoliths has two components — the field enclosing the 40W/m² geotherm (Figure 2) and the high temperature ‘kink’ of enigmatic origin. In this figure, this region, and a similar kink at the lower end of the intraplate geotherm, are represented as transient phenomena associated with adiabatic upwelling and magmatism linked with lithospheric plate movement and thinning.

The composition of olivine phenocrysts in Hawaiian picrites and in MOR picrites varies up to Mg# 91.3 and Mg# 92.1 respectively. The compositions and liquidus temperatures of the magmas crystallizing the most magnesian phenocrysts can be estimated and anhydrous liquids temperatures (at 1 bar pressure) of Hawaiian tholeiitic picrites average 1365°C, and for MOR picrites average 1335°C. Water contents of the magmas decrease in the order Hawaiian picrites, E-MOR picrites to N-MOR picrites, and consideration of liquidus depression by these water contents leads to the conclusion that all primitive magmas had liquidus temperatures of approximately 1325°C at ~ 1 bar. The data from primitive magmas suggests that the temperature contrast between “hot-spot” and MOR primary magmas is ≤ 20°C. Application of information from partial melting studies of model (pyrolite) source compositions and of the liquidus phases of the hot-spot and MOR picrites leads to the conclusion that both ‘hot-spot‘ and MOR primary basalts are derived from mantle with potential temperature Tp ~ 1430°C. Insofar as primitive magmas may be used to infer the potential temperature of their sources, there is no evidence for a temperature contrast of Δ Tp = 100-250°C between ‘hot-spot’ or ‘deep mantle plume’ sources and ambient (MOR source) asthenospheric mantle.

Although magma temperatures are similar, the residual mantle compositions for Hawaiian picrites are refractory harzburgites, more refractory (including Cr/Cr+Al ratio) than the lherzolite to harzburgite residue from MOR picrite extraction. It is argued that the buoyancy plume and geophysically anomalous mantle beneath the Hawaiian swell is due to compositional and not temperature contrasts in the upper mantle. The four-component mixing identified in the Hawaiian source is attributed to interaction between old subducted lithospheric slabs, buoyant or suspended in the upper mantle, and surrounding ambient mantle at Tp =1430°C. A cartoon representing the model is shown in Figure 3.

Figure 3: Cartoon suggesting a model for intraplate ‘hot–spot’ volcanism attributing the primary cause of the melting anomaly (‘hot–spot’) to a compositional anomaly within the mantle beneath the asthenosphere, i.e. > 150–200 km depth. Increased (C+H+O) volatile flux is attributed to redox melting at the interface of ambient mantle with CH₄+H₂O fluid (fO₂=IW+1) and old subducted slab or delaminated cratonic lithosphere (fO₂=IW+3,4). The geochemical anomaly (‘plume source’) includes deep mantle (CH₄+H₂O fluid), sub-asthenospheric and asthenospheric mantle, and old, subducted sources in a mixing and upwelling process. The geophysical anomaly (e.g. ‘Hawaiian Swell’) is attributed to relative densities of subducted slab/delaminated lithosphere and ambient mantle, to the melt-enhanced column above the compositional anomaly, and the thermal anomaly and thermal erosion of the lithosphere by diapirism from >150 to 200 km depths leading to magma segregation within lithospheric depths (40-60 km) (from Green & Falloon, 1998).

References

updated 1st January, 2007