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The Source of Oceanic Volcanism – Plumes vs the Low-Velocity Zone

Dean C. Presnall

Department of Geosciences, University of Texas at Dallas, Richardson, TX

presnall@utdallas.edu


This webpage is a summary of: Presnall, D.C. and Gudfinnsson, G.H. (2011). Oceanic volcanism from the low-velocity zone – without mantle plumes, Journal of Petrology, Peter J. Wyllie Volume, in press. (available on-line at Journal of Petrology, Advanced Access).


 

In a new model for oceanic volcanism, Presnall & Gudfinnsson (2011) argue that oceanic volcanism originates from magmas in the seismic low-velocity zone (LVZ) that ascend through fractures in the seismic lithosphere. This modeling is supported by:

  1. basalt glass compositions from Hawaii and mid-ocean ridges;
  2. data on the melting and crystallization behavior of natural lherzolite;
  3. phase relations of model lherzolite in the systems CaO-MgO-Al2O3-SiO2-CO2 (CMAS-CO2) and CaO-MgO-Al2O3-SiO2-Na2O-FeO (CMASNF) at pressures from 1 atm to 6 GPa; and
  4. shear-wave velocities and anisotropy for the oceanic upper mantle.

The LVZ, which extends from about 65 to 220 km depth in the ocean basins, was first suggested to be a region of partial melting by Anderson & Sammis (1970). Wyllie & Huang (1975) and Eggler (1976) then suggested that melting in the LVZ is caused by CO2. This argument has recently been reinforced (Presnall & Gudfinnsson, 2005, 2011; Hirschmann, 2010). A common assumption is that the melt fraction is very small throughout the depth range of the low-velocity zone. However, for a constant amount of CO2, phase relations along a normal oceanic geotherm for a lherzolite mineralogy in the CMAS-CO2 system (Gudfinnsson & Presnall, 2005) indicate a large tholeiitic melt fraction at ~ 4-5 GPa and smaller alkalic (low-SiO2) melt fractions at both lower (~ 3 GPa) and higher (> 6 GPa) pressures along the geotherm (Figure 2b in Presnall & Gudfinnsson, 2011). Temperatures for the model system are adjusted to those for natural lhezolite on the basis of the volatile-free solidus of garnet lherzolite at 5 GPa, 1500°C (Lesher et al., 2003) and the solidus for a natural carbonated lherzolite (Dasgupta & Hirschmann, 2006). The pressure range of Hawaiian tholeiitic melt-extraction at ~ 4-5 GPa (~ 130-150 km depth) is based on the olivine-controlled crystallization trend of Puna Ridge tholeiitic glasses at Kilauea Volcano, Hawaii (Clague et al., 1995) combined with experimental data in the CMAS-CO2 system (Gudfinnsson & Presnall, 2011, Figure 1).

In the CMAS-CO2 system, the melt-fraction at a given P, T, and bulk composition in equilibrium with a lherzolite mineralogy is a linear function of the amount of CO2 (Gudfinnsson & Presnall, 2005, Figure 4). Because virtually all of the CO2 is dissolved in the melt, this would also be true for a natural carbonated lherzolite. For the depth range where a geotherm grazes the volatile-free solidus (Figures 2a, b in Presnall & Gudfinnsson, 2011), tholeliites exist at significantly enhanced melt-fractions. At both higher and lower pressures, where the geotherm is at significantly lower temperatures than the volatile-free solidus, melt-fractions are much smaller. Thus, the equilibrium “magma stratigraphy” in the LVZ along a normal geotherm, as modeled by the CMAS-CO2 system, would consist of small melt-fractions of alkalic magma at shallow depths (~ 100 km), large melt-fractions of tholeiitic magma at intermediate depths (~ 120-170 km) and small melt-fractions of alkalic magma at extreme depths (~190 km).

A southeastward-propagating fracture system (Jackson et al., 1972, 1975; Jackson & Shaw, 1975) that extends into the low-velocity zone and taps progressively shallower magmas to the southeast would show exactly the sequence and relative volumes of magmas observed along the Hawaiian chain as well as for individual Hawaiian volcanoes. The depth of ~ 130-150 km for maximum melting, as indicated by the phase relations (Gudfinnsson & Presnall, 2005), is consistent with the depth of maximum shear-wave anisotropy (Ekström, 2000) and lowest vertical shear-wave velocity (Maggi et al., 2006) throughout the central Pacific Basin. The classical concept of decompression melting (for example in a rising mantle plume) is replaced by sampling of the equilibrium magma-stratigraphy that exists in the Pacific LVZ along a normal age-dependent geotherm.

At mid-ocean ridges, the range of basalt glass compositions is controlled dominantly by low-pressure crystallization of olivine, plagioclase, clinopyroxene and orthopyroxene, as at Hawaii. However, unlike Hawaii, no mid-ocean ridge basalt (MORB) glass compositions, including those at Iceland, exist that indicate more than an insignificant amount of initial olivine-controlled crystallization at low pressures. As the PetDB database used by Presnall & Gudfinnsson (2011) is very large (6,933 glass analyses), this is a robust constraint.

Figure 1. Comparison of horizontal shear velocities (Vsh) for a SW-NE section across the Pacific Basin (Tan & Helmberger, 2007) (c) with solidus curves and geotherms for Hawaii (b) and mid-ocean ridges (a). G/D is the graphite-diamond transition (Bundy et al., 1961). The geotherm for Hawaii assumes a 1500°C adiabat. The natural volatile-free lherzolite solidus (bold dashed line) is constrained at 5 GPa,1600°C by the determination of Lesher et al. (2003). The natural carbonated lherzolite solidus (bold continuous line) is from Falloon & Green (1989) and Dasgupta & Hirschmann (2006). Tp is potential temperature. The P-T space between these two curves is a region of olivine+orthopyroxene+ clinopyroxene+garnet+carbonate-bearing melt. The slight change in slope of the carbonated lherzolite solidus at 3.5 GPa is caused by intersection of the solidus with the dolomite-magnesite transition (Dalton & Presnall, 1998).

Phase relations for the systems CaO-MgO-Al2O3-SiO2 and CaO-MgO-Al2O3-SiO2-Na2O-FeO (Presnall & Gudfinnsson, 2011, and references therein) show that MORB magmas, including those at Iceland, are extracted at ~ 1240-1260°C, 1.2–1.5 GPa. Pressures of melt-extraction higher than this would show olivine-controlled crystallization paths like that found at Hawaii. Thus, melt-extraction at mid-ocean ridges occurs at the shallow depth of ~ 65 km, the maximum pressure for melt-extraction without olivine-controlled crystallization on subsequent cooling. This depth is in excellent agreement with seismic data (Nishimura & Forsyth, 1989) showing that the depth range of minimum shear velocity (maximum melt-fraction) at the crest of the East Pacific Rise is ~ 40-70 km. These very low P-T conditions of melt-extraction are within the thermal boundary layer, and indicate a perturbed geotherm that grazes the lherzolite solidus at these conditions (Figure 2, Presnall & Gudfinnsson, 2011). Such a geotherm requires a steepened conductive dT/dP gradient beneath ridges, which is consistent with heat-flow data (e.g., DeLaughter et al., 2005).

Modeling by Conder et al. (2002) indicates eastward and upward flow within the low-velocity zone beneath and just westward of the East Pacific Rise. In agreement with this modeling, a northwest to southeast cross-section of shear-wave anisotropy across the Pacific Basin (Ekström, 2000) shows that the depth of maximum anisotropy is fairly constant at ~130-150 km for the central Pacific but shallows significantly as the East Pacific Rise is approached from the northwest. Similarly, Nishimura & Forsyth (1989) found that the depth of minimum vertical shear velocity decreases from ~150 km to ~ 40-70 km as the East Pacific Rise is approached from the northwest. This essentially perfect agreement of the seismic and phase equilibrium data for the depths of maximum melting and tholeiitic magma production at Hawaii and mid-ocean ridges is powerful evidence that all oceanic volcanism comes from fractures that tap melts in the low-velocity zone, not from plumes.

References

  • Anderson, D.L. and Sammis, C.G. (1970). Partial melting in the upper mantle, Physics of the Earth and Planetary Interiors, 3, 41-50.
  • Bundy, F.P., Bovenkirk, H.P., Strong, H.M. & Wentorf, R.H., Jr (1961). Diamond-graphite equilibrium line from growth and graphitization of diamond. Journal of Chemical Physics, 35, 383-391.
  • Clague, D.A., Moore, J.G., Dixon, J.E. and Friesen, W.B. (1995). Petrology of submarine lavas from Kilauea’s Puna Ridge, Hawaii. Journal of Petrology, 36, 299-349.
  • Conder, J.A., Forsyth, D.W. and Parmetier, E.M. (2002) Asthenospheric flow and asymmetry of the East Pacific Rise, MELT area. Journal of Geophysical Research, 107, 2344, d0i:10.1029/2001JB000607.
  • Dalton, J.A. & Presnall, D.C. (1998). The continuum of primary carbonatitic- kimberlitic melt compositions in equilibrium with lherzolite: Data from the system CaO-MgO-Al2O3-SiO2-CO2 at 6 GPa. Journal of Petrology, 39, 1953-1964.
  • Dasgupta, R. and Hirschmann, M.M. (2006). Melting in the Earth’s deep upper mantle caused by carbon dioxide. Nature, 440, 659-662.
  • Eggler, D.H. (1976). Does CO2 cause partial melting in the low-velocity layer of the mantle?, Geology, 4, 69-72.
  • Ekström, G. (2000). Mapping the lithosphere and asthenosphere with surface waves: Lateral structure and anisotropy. In: Richards, M.A., Gordon, R.G., and van der Hilst, R.D. (eds.) The History and Dynamics of Global Plate Motions. American Geopysical Union, Geophysical Monograph 121, 239-254.
  • Falloon,T.J. & Green, D.H. (1989). The solidus of carbonated, fertile peridotite. Earth and Planetary Science Letters, 94, 364-370.
  • Gudfinnsson, G.H. and Presnall, D.C. (2005). Continuous gradations among primary carbonatitic, kimberlitic, melilititic, picritic, and komatiitic melts in equilibrium with garnet lherzolite at 3-8 GPa. Journal of Petrology, 46, 1645-1659.
  • Hirschmann, M.M. (2010). Partial melt in the oceanic low velocity zone, Physics of the Earth and Planetary Interiors, 179, 60-71.
  • Jackson, E.D. and Shaw, H.R. (1975). Stress fields in central portions of the Pacific plate: Delineation in time by linear volcanic chains. Journal of Geophysical Research, 80, 1861-1874.
  • Jackson, E.D., Silver, E.A. and Dalrymple, G.B. (1972). Hawaiian-Emperor chain and its relation to Cenozoic circumpacific tectonics. Geological Society of America Bulletin, 83, 601-618.
  • Jackson, E.D., Shaw, H.R. and Bargar, K.E. (1975). Calculated geochronology and stress field orientations along the Hawaiian chain. Earth and Planetary Science Letters, 26, 145-155.
  • Lesher, C.E., Pickering-Witter, J., Baxter, G., and Walter, M. (2003). Melting of garnet peridotite: Effects of capsules and thermocouples, and implications for the high-pressure mantle solidus. American Mineralogist, 88, 1181-1189.
  • Maggi, A., Debayle, E., Priestley, K. and Barroul, G. (2006). Multimode surface waveform tomography of the Pacific Ocean: a closer look at the lithospheric cooling signature. Geophysical Journal International, 166, 1384-1397.
  • Nishimura, C.E. and Forsyth, D.W. (1989). The anisotropic structure of the upper mantle in the Pacific, Geophysical Journal, 96, 203-229.
  • Tan,Y. & Helmberger, D.V. (2007). Trans-Pacific upper mantle shear velocity. Journal of Geophysical Research, 112, BO8301, doi:10.1029/ 2006JB004853.
  • Wyllie, P.J. and Huang, W.-L. (1975). Influence of mantle CO2 in the generation of carbonatites and kimbelites, Nature, 257, 297-299.
last updated 5th May, 2011
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