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Boundary layer fabric, dynamics and the origin of midplate volcanoes


Don L. Anderson

Seismological Laboratory, California Institute of Technology, Pasadena, CA 91125,

This webpage is an overview of the LLAMA model (Anderson, D.L., Hawaii, boundary layers and ambient mantle–geophysical constraints, J. Petrology, published online December 2, 2010, doi: 10.1093/petrology/egq068, available on-line at Journal of Petrology, Advanced Access).

The reference list in that paper is updated at the end of this contribution to include 2010 and 2011 publications.


The anisotropic seismic low-velocity layer (LVL) in the mantle under ocean basins is delineated by the Gutenberg (G) and Lehmann (L) discontinuities at average depths of about 70 and 220 km depth, respectively. A thinner, deeper, and less anisotropic LVL terminates near 220 km under cratons. The 220 ± 20 km depth level (L) is a universal and fundamental tectonic boundary in the mantle. The mantle above this level, Gutenberg’s Region B (Figure 1), is anisotropic because it is advected and sheared by plate motions and contains aligned melt arrays (Figure 2). Although it is weak and deforming, it supports a conduction gradient and is therefore part of the thermal boundary layer (TBL) and not part of the so-called convecting mantle. The mantle below this level is decoupled from plate motions and provides a relatively fixed shallow reference system. Oceanic volcanism taps magmas in the LVL, which ascend by hydraulic magma fracturing rather than by thermal erosion. Region B contains the key to a number of geodynamic and geochemical phenomena that have been attributed to much deeper boundary layers.


Figure 1: Nomenclature of the mantle. The classical subdivisions of the mantle are Regions B, C, D’ and D”. Region B constitutes the upper boundary layer of the mantle. It is much larger than other proposed reservoirs, such as D” and the continental lithosphere. Gutenberg (1959) placed the boundary between Regions B and C near 220 km. The classical lower mantle (Region D) starts at ~1000 km, not at 650 km. The low-velocity anisotropic layer (LLAMA) extends from the Gutenberg discontinuity (G) to the Lehmann discontinuity (L). A schematic potential temperature (Tp) geotherm is shown. Because of internal heating and secular cooling, D” may have a lower Tp than the base of B. Region B, the most heterogeneous and anisotropic region of the mantle, includes the seismic lid and the LVL, which extends from the Gutenberg or G-discontinuity (70 ± 30 km depth) to the Lehman or L-discontinuity (220 ± 20 km depth). Any point on a geotherm can be assigned a potential temperature, which is the temperature the point would have if adiabatically expanded to zero pressure. The concept of a Tp, and the ability to calculate it for any point on a geotherm, does not require that the mantle be adiabatic or convecting.


Figure 2: The LVL is anisotropic. Horizontally propagating shear waves that are polarized in the vertical plane, or along the symmetry axis, travel at a velocity VSV. VSH is the shear velocity of waves with particle motions in the symmetry plane. Left. Seismic shear velocities, VSH and VSV, in the central Pacific (Shapiro & Ritzwoller, 2002) and in the global reference model PREM. The steep conduction part of the geotherm extends to > 200 km depth. Right. The laminated boundary layer model (Kawakatsu et al., 2009) along with the interpretation that it is a sheared layer composed of refractory and melt-rich lamellae. The temperature increases by about 6-10°C/km with depth in a layer with a negative Vs gradient. The minimum in VSV corresponds to the maximum melt content. VSV is mainly sensitive to the temperature gradient. The G discontinuity is marked by a large decrease in shear wave velocity, typically between 6 and 9% (e.g., Kawakatsu et al., 2009). The L discontinuity is usually somewhat smaller. The shear wave anisotropy between G and L is of similar magnitude. This implies the presence of thin, low-rigidity sub-layers or sills in the LVL.


The maximum melt content occurs at the minimum in shear velocity, VSV, which is ~60 km depth under ridges and 140-150 km depth in the central Pacific (Figure 2). The strength and orientation of the anisotropy at those depths implies a laminated lithology with aligned melt accumulations (LLAMA). The upper mantle surrounding Hawaii has higher than average absolute seismic velocities for the age of the plate, relatively low inferred temperatures and a variable plate thickness. Ambient mantle in the boundary layer under the central Pacific and elsewhere, is not the same as the shallow mantle under ridges.

The shear wave delays in the upper 220 km of the mantle, both globally and regionally, account for most of the observed teleseismic delays in well-instrumented well-sampled areas. Similar delays across the Hawaiian swell, which is poorly sampled by seismic rays, are routinely attributed to deep cylindrical structures. The observed anisotropy and heterogeneity of the shallow mantle in the vicinity of Hawaii yields near-vertical, deep, plume-like structures when relative travel-times to Hawaii are modeled in the usual way, i.e., assuming a smooth isotropic shallow mantle. These are artifacts, however.

Refractory remnants of ancient high-temperature melting processes remain in the shallow mantle because they are relatively buoyant. The fabric of the subplate boundary layer arises from shear-bands and aligned, melt-rich lamellae embedded in the ancient buoyant refractory matrix. The seismic lid and the rheological lithosphere are in the cold, top-most refractory layer (Figure 3). The LVL is infiltrated by melts and fluids, which are aligned by shear. The anisotropy of the LVL is mainly due to this dynamically maintained macroscopic fabric, which also allows melts to be retained. Melts are released when the laminar fabric is disrupted by steps in the lid (Figure 1), by fracture zones, and by shear-driven upwellings. Since the fabric of the boundary layer involves melt-rich lamellae separated by thick depleted harzburgite layers (Figure 2), rather than a chaotic marble-cake of multiple components, it is easy to understand why geochemical arrays are simple two-component mixing arrays diverging from a common component. The major element, trace-element and isotopic compositions of oceanic and continental intra-plate magmas and their xenoliths do not require sources from deeper than the boundary layer of the upper mantle. In addition, many of the isotopic signatures are acquired from the atmosphere, hydrosphere and sediments.


Figure 3: The LLAMA model. The seismic lid grows at the expense of the LVL. The lithosphere-lid system overlies aligned melt arrays, which occur in shear-bands. The red regions are low-rigidity lamellae. If these are melt-rich, then when they freeze they will have high seismic velocities and the boundary layer will then have properties similar to cratonic mantle. Scaling relations suggest that shear-bands and melt lenses can be tens of kilometers apart. This macro-fabric is in addition to the grain- and vein-scale micro-fabric, e.g., marblecake and oriented crystal fabric.


The recycling of upper and lower continental crust, sediments and oceanic lithosphere into the shallow mantle, plus metasomatism by slab-derived fluids and residual ancient lithospheric components, can account for the entire range of chemical and isotopic signatures in intra-plate basalts. Even the largest of the large igneous provinces apparently originates in the shallow mantle. The compositional heterogeneity of the mantle is intrinsic to plate tectonics and does not require chaotic stirring of a limited number of physically separate mantle reservoirs into “the convecting mantle”. There is no requirement or evidence that any of this material circulated to the deep mantle before being sampled.

Intimate juxtaposition of diverse lithologies, is accomplished by boundary-layer shearing (Figure 4). Seismic results show that most of the lithologic heterogenity of the mantle occurs above 220 km depth. Since seismic velocities are strongly affected by melt and other fluids in the mafic parts of the mantle, this suggests that this is where midplate magmas originate. Fluids and magmas intercalated with high-velocity, refractory components are mainly responsible for both the heterogeneity and anisotropy of the shallow mantle. Aligned or anastomosing arrays of melt channels result in anisotropy of seismic velocity, viscosity, conductivity and permeability. This unusual region of the mantle is composed of laminated lithologies and aligned melt arrays (LLAMA) (Figure 4). The mechanism that leads to this fabric is lubricated lateral advection and magma accumulation. The temperatures, compositions and volumes of intra-plate and continental breakup magmas are consistent with these ambient mantle conditions and top-down tectonics. Upwellings are part of plate tectonics but they are generally not deep-seated or buoyant in their own right, i.e., they are not in the form of thermal plumes. Mantle plumes are not required by any petrological or geophysical data.


Figure 4: A variable thickness plate moves to the right and entrains the top of the underlying mantle, forming a shear boundary layer. Lithospheric steps and fracture zones cause disruption of this boundary layer. The hotter, deeper parts of the boundary layer are almost stationary relative to the surface. Observed teleseismic delays result almost entirely from anisotropy and lateral variation in the shearing horizontal boundary layer, rather than from continuous vertical features extending to >1000 depth (plumes). The lower portions of LLAMA are hotter than MORB. Most samples of the mantle, away from plate boundaries, are derived from this sublayer. Subduction can also displace material out of Region C but no deeper material is required by either mass balance or geochemistry of magmas and their xenoliths.


Laminated sublayer, laminar advection

Plate motions are decoupled from the deeper mantle by a ~200 km thick shear boundary layer with a high thermal gradient (Figures 3 & 4). I summarize the more important implications of this geodynamic model as follows:

  1. Potential temperatures in the deeper parts of LLAMA are higher than MORB temperatures and can be higher than in D” (Figure 1).
  2. The 220 ± 20 km depth level (L) is a fundamental tectonic boundary. There is a pronounced change in the fabric of the mantle and the shear velocity gradient at this depth. The mantle above this depth is by far the most anisotropic and heterogeneous part of the mantle (Figure 1).
  3. Most mantle samples originate above the L discontinuity (Figure 2) and most of the geochemical components originate in the surface layers of the Earth and are stored in LLAMA.
  4. LLAMA, along with the crust, accounts for most of the global variation in seismic travel times (Figure 4).
  5. The LVL under Hawaii starts deep. It can be converted to a craton-like boundary layer simply by lowering the temperature and freezing the melt-rich lamella (Figure 3).
  6. That part of the oceanic upper mantle that supports a conduction gradient is much thicker than the part that has been cooled by conduction since the plate formed. It is comparable to the steady-state thickness and the thickness of cratons. The implication is that new thermal boundary layers are emplaced atop pre-existing ones rather than atop isothermal or adiabatic interiors (Figure 5).
  7. Melts can be extracted from the upper boundary layer of the mantle. The “fixity” and fluxes of intraplate volcanoes, and the associated shallow bathymetry can all be explained by boundary layer dynamics, with sources in and near the base of a lubricating layer of ambient mantle, i.e., LLAMA (Figure 4).
  8. Ambient mantle surrounding Hawaii has relatively high seismic velocities and normal or low inferred temperatures. Nevertheless, it is a truism that half of the teleseismic delays to Hawaiian stations within the local region are slower than the other half.
  9. Waveform modeling, high-resolution seismology and the central limit theorem show that low-velocity anomalies in the mantle are more pronounced, sharper and less continuous than the tomographic images on which current geochemical and geodynamic speculations are based. Sharp boundaries are unlikely to be caused by long-lived thermal upwellings.
  10. The kind of shallow heterogeneity and anisotropy measured by seismology, when sampled by sparse near-vertical teleseismic rays and analyzed by the usual tomographic methods, will yield near-vertical or tilted plume-like structures centered under the seismic array. This will occur wherever the array is placed, providing misleading indications of plumes.

Figure 5: Schematic diagram of the growth of a thermal boundary layer that is emplaced atop a previously cooled part of the mantle. Mid-ocean ridges form by the fracturing of the top part of a mature boundary layer, which is then dragged away by diverging plates. The boundary layer is perturbed by the formation of a ridge (e.g., Presnall & Gudfinnsson, 2011). Mantle upwells from some shallow depth to fill the opening crack, imposing an adiabatic gradient on the part of the mantle above the level of upwelling, which subsequently cools, forming a new thermal boundary layer. The deeper parts of the boundary layer can have higher Tp than ridge basalts or even all or part of D”. The red arrows, projected to the surface, give the Tp of points on the geotherm.


There is no longer any plausible evidence or argument for deep plumes under midplate volcanoes or for accessible lower mantle reservoirs. The widely quoted speculations to the contrary ignore the dominant roles of fabric and shallow boundary layers. The interpretation of seismic anisotropy in terms of aligned macroscopic lamellae and sills rather than as subsolidus olivine orientation, has immediate implications for mantle petrology and geochemistry and the locations and geometric relations of the sources of geochemical components.

A guide to the 2009-2011 & related literature

The message of Anderson (2010) and the references therein is that the surface boundary layer of the mantle, horizontal advection, and fabric are the keys to midplate magmatism and mantle dynamics. Recent studies have confirmed these views. It is now possible to model boundary layer scales of mantle dynamics (e.g., Adam et al., 2010; Schmandt & Humphreys, 2010a; Conrad et al., 2010; Stadler et al., 2010; Faccenna & Becker, 2010; Leahy et al., 2010; Van Wijk et al., 2010), to explain the low Tp in D” (Schuberth et al., 2009; Sinha & Butler, 2007), and to map boundary layer stratigraphy with xenoliths (Ishikawa et al., 2011). High-resolution seismic techniques confirm the importance of Region B (Schmandt & Humphreys, 2010b; Ritsema et al., 2010; Leahy et al., 2010; Sun & Helmberger, 2010). The components of oceanic basalts can all be found in the shallow mantle (Griffin et al., 2009; Salters & Sachi-Kocher, 2010). Laboratory experiments confirm the melt segregation mechanism for anisotropy (Kohlstedt et al., 2009; Holtzman & Kendall, 2010). The older references in the list below are mentioned in the figure captions or are essential for understanding the background of LLAMA, and some of the geochemical implications (e.g., Hart et al., 1992).

  • Adam, C., M. Yoshida, T. Isse, D. Suetsugu, Y. Fukao, G. Barruol, (2010), South Pacific hotspot swells dynamically supported by mantle flows, Geophys. Res. Lett., 37, L08302.
  • Bryan, S.E., I. Ukstins Peate, D.W. Peate, S. Self, D.A. Jerram, M.R. Mawby, J.S. Marsh, J.A. Miller, (2010), The largest volcanic eruptions on Earth, Earth-Science Reviews, 102, 207-229.
  • Cañón-Tapia, E., (2010), Origin of large igneous provinces: The importance of a definition, in Cañón-Tapia, E., and Szakács, A., eds., What Is a Volcano?, Geol. Soc. Am. Special Paper 470, 77–101, doi: 10.1130/2010.2470(06).
  • Conrad, C.P., Benjun Wu, E.I. Smith, T.A. Bianco, A. Tibbetts, (2010), Shear-driven upwelling induced by lateral viscosity variations and asthenospheric shear: A mechanism for intraplate volcanism, PEPI, 178, 162-175.
  • Dasgupta, R., M.G. Jackson, C.-T. A. Lee, (2010), Major element chemistry of ocean island basalts–conditions of mantle melting and heterogeneity of mantle source, Earth Planet. Sci. Lett., 289, 377-392.
  • Faccenna, C. and T.W. Becker, (2010), Shaping mobile belts by small-scale convection, Nature, 465, 602-605, doi:10.1038/nature09064.
  • Griffin, W.L., S.Y. O’Reilly, J.C. Afonso, and G.C. Begg, (2009), The composition and evolution of lithospheric mantle: a re-evaluation and its tectonic implications. J. Petrology, 50, 1185-1204.
  • Gutenberg, B. (1959). Wave velocities below the Mohorovicic discontinuity, Geophys. J. Royal Astr. Soc., 2, 348-352.
  • Hamilton, W.B., (2011), Plate tectonics began in Neoproterozoic time, and plumes from deep mantle have never operated, Lithos, doi:10.1016/j.lithos.2010.12.007
  • Hart, S.R., E.H. Hauri, L.A. Oschmann, and J.A. Whitehead, (1992), Mantle plumes and entrainment–isotopic evidence, Science, 256, 517–520.
  • Holtzman, B.K., Kohlstedt, D.L., Zimmerman, M.E., Heidelback, F., Hiraga, T. and Hustoft, J., (2010). Melt segregation and strain partitioning: Implications for seismic anisotropy and mantle flow, Science, 301, 1227-1230.
  • Holtzman, B.K., and J.-M. Kendall, (2010), Organized melt, seismic anisotropy, and plate boundary lubrication, Geochem. Geophys. Geosyst., 11, Q0AB06, doi:10.1029/2010GC003296.
  • Ishikawa, A., D.G. Pearson, C.W. Dale, (2011), Ancient Os isotope signatures from the Ontong Java Plateau lithosphere: Tracing lithospheric accretion history, Earth Planet. Sci. Lett., 301, 159-170.
  • Kawakatsu, H., Kumar, P., Takei, Y., Shinohara, M., Kanazawa, T, Araki, E. & Suyehiro, K. (2009), Seismic evidence for sharp lithosphere-asthenosphere boundaries of oceanic plates. Science, 324, 499-502; doi:10.1126/science.1169499.
  • Kohlstedt, D.L., Zimmerman, M.E. & Mackwell, S.J., (2009), Stress-driven melt segregation in partially molten rocks, J. Petrology, doi:10.1093/petrology/egp043.
  • Leahy, G.M., J.A. Collins, C.J. Wolfe, G. Laske and S.C. Solomon, (2010), Underplating of the Hawaiian Swell: evidence from teleseismic receiver functions, Geophys. J. Int., 183, 313–329, 2010, doi: 10.1111/j.1365-246X.2010.04720.x.
  • Ritsema, J., A. Deuss, H.J. van Heijst and J.H. Woodhouse, (2010), S40RTS: a degree-40 shear-velocity model for the mantle from new Rayleigh wave dispersion, teleseismic traveltime and normal-mode splitting function measurements, Article first published online: 14 Dec 2010, doi: 10.1111/j.1365-246X.2010.04884.x
  • Salters, V.J.M., A. Sachi-Kocher, (2010), An ancient metasomatic source for the Walvis Ridge basalts, Chem. Geol., 273, 151-167.
  • Schmandt, B., and E. Humphreys, (2010a), Complex subduction and small-scale convection revealed by body-wave tomography of the western United States upper mantle, Earth Planet. Sci. Lett., 297, 435-445.
  • Schmandt, B., and Humphreys, E., (2010b), Seismic heterogeneity and small-scale convection in the southern California upper mantle, Geochem. Geophys. Geosyst., 11, Q05004.
  • Schuberth, B.S.A., H.P. Bunge, G. Steinle-Neumann, C. Moder, and J. Oeser, (2009). Thermal versus elastic heterogeneity in high-resolution mantle circulation models with pyrolite composition: High plume excess temperatures in the lowermost mantle. Geochem. Geophys. Geosyst., 10, Q01W01; doi:10.1029/2008GC002235.
  • Sinha, G. and S.L. Butler, (2007), On the origin and significance of subadiabatic temperature gradients in the mantle, J. Geophys. Res., 112, B10406; doi:10.1029/2006JB004850.
  • Shapiro, N.M. and M.H. Ritzwoller, (2002), Monte-Carlo inversion for a global shear velocity model of the crust and upper mantle, Geophys. J. Int., 151, 88-105.
  • Stadler, G. et al. (2010), The dynamics of plate tectonics and mantle flow: From local to global scales, Science, 329, 1033-1038, doi: 10.1126/science.1191223.
  • Sun, Daoyuan and D. Helmberger, (2010), Upper-mantle structures beneath USArray derived from waveform complexity, Geophys. J. Int., published online: 30 Nov 2010, doi: 10.1111/j.1365-246X.2010.04847.x.
  • Willbold, M. and A. Stracke, (2010), Formation of enriched mantle components by recycling of upper and lower continental crust, Earth Planet. Sci. Lett., 297, 188-197.
last updated 16th March, 2011