Velocity-depth profile

Spherically symmetrical (1D) Earth models such as the Preliminary Reference Earth Model (PREM, Dziewonski and Anderson, 1981), iasp91 (Kennett and Engdahl, 1991) and AK135 (Kennett et al., 1995) are based on a large catalogue of body-wave travel times and normal mode observations (standing waves due to the free-oscillation of the Earth) and provide profiles of wave velocity, density, and attenuation throughout the Earth. In these models, velocities are refined to a set of polynomial functions that operate over certain depth ranges, with assumed seismic discontinuities at certain depths (Fig. 1.1). Several discontinuities in the upper mantle have been well established, the properties of which provide a link between mineral physics experiments and seismic observations as well as important constraints on models of mantle composition and dynamics. These discontinuity properties include the discontinuity depths and topography, the size of the velocity and

9 density increases and the sharpness of the boundaries. Discontinuities in the transition zone, for example, are considered to arise from nominally isochemical first order mineral phase transitions. The depths of these discontinuities can, therefore, be interpreted, through experimental studies, to provide information on the temperature of the mantle.

Figure 1.1 The density ρ, S wave velocity Vs and P wave velocity Vp as a function of depth according to the PREM model (Dziewonski and Anderson, 1981).

1.1.1.1 Discontinuity depth

The depths of seismic discontinuities can be obtained using a variety of approaches including analysis of seismic waves that bottom near the discontinuities (refractory seismology) or seismic waves that are either reflected or converted at the discontinuity (Shearer, 2000). The Mohorovicic discontinuity marks the base of the crust and the top of the Earth’s upper mantle. Two major global seismic discontinuities of the mantle transition zone at mean depths of about 410 and 660 km (hereafter termed 410 and 660) have been well established by observations in SS (S phase with one reflection at the surface) (e.g. Heit et al., 2010 and references therein), PP (P phase with one reflection at the surface) (e.g.

Deuss, 2009 and references therein), and P’P’(abbreviation of PKIKP, a P phase bottoming in the inner core with one reflection at the surface) precursors (e.g. Day and Deuss, 2013

10 and references therein), Ps (P-to-S converted wave) and Sp (S-to-P converted wave) conversions at the discontinuities (receiver functions) (e.g. Andrews and Deuss, 2008 and references therein), and ScS (S wave reflected from the core-mantle boundary) reverberations (e.g. Katzman et al., 1998 and references therein), whereas the presence of a discontinuity at about 520 km (hereafter termed 520) has been controversial (e.g. Kind and Li, 2015 and references therein). This discontinuity is clearly observed in some regions by SS precursor, PP precursor, ScS reverberation and receiver function studies but it is not visible in other regions. Therefore, the 520 does not appear to be a laterally homogeneous global discontinuity, but also it cannot be considered a local one because it has a relatively common distribution (e.g. Kind and Li, 2015 and references therein). All modern global reference Earth models (Brown and Shankland, 1981; Kennett and Engdahl, 1991; Kennett et al., 1995) include a sudden increase of the velocities of the elastic waves and densities at 410 and 660 km, while the 520 is absent in all but the mineral physics based model of Cammarano et al. (2005). The depth ranges reported for the 410, 520 and 660 are from 390 km to 430 km, from 500 km to 520 km and from 650 km to 680 km, respectively. The SS precursor observations may be the most suitable for globally averaged depth estimates due to the wide distribution of their bounce points which provides comprehensive global coverage (Shearer, 2000), resulting in mean discontinuity depths close to 410, 520 and 660 km. These average discontinuity depths to a first approximation match well with those expected for the pressure and temperature-induced phase transitions from olivine to wadsleyite, from wadsleyite to ringwoodite and from ringwoodite to bridgmanite (Brg) plus ferropericlase (Fp), respectively. If this is true, the discontinuity depth could provide relatively direct information on mantle temperatures using knowledge of the transformation boundaries of certain mineral phase changes.

1.1.1.2 Discontinuity topography

Differences among discontinuity depth estimates obtained in different studies imply that global variations in discontinuity depths, i.e. topography, exists (Shearer, 2000) that can be most likely attributed to mantle temperature variations. Discontinuity topography may be

11 detected using SS precursor techniques, ScS reverberations or Ps and Sp conversion techniques (Kind and Li, 2015). Precursor techniques can reach locations without local stations or earthquake sources, allowing a better global distribution of sampling points. The resolution of long period SS precursors does not allow small-scale topography of the discontinuities to be resolved because they are associated with maximum travel-time phases and can be contaminated by small scale, off-great-circle-path structure (Kind and Li, 2015; Shearer, 2000). Small-scale discontinuity topography less than a few tens of kilometers can be detected instead using Ps and Sp conversion techniques, i.e. waves reflected or converted at the discontinuities close to either the source or receiver (receiver function) (Kind and Li, 2015). The converted phases are minimum travel-time phases and allow a better resolution due to their smaller Fresnel zones but are restricted to near-station or earthquake locations (Kind and Li, 2015).

The amplitude of the global 660 topography (38-50 km) appears to be larger than the 410 topography (22-40 km) (Flanagan and Shearer, 1998; Gu et al., 2012; Shearer, 1991, 1993). Whether the 410 and 660 are globally anticorrelated is still under debate due to contradictory observations (Gu et al., 2003; Gu et al., 2012; Hu et al., 2013; Humphreys et al., 2000; Ramesh et al., 2005). Both global and high-resolution local observations show depression of the 660 of between 20 and 50 km at subduction zones on a large scale (Tibi and Wiens, 2005; Tonegawa et al., 2005; Tonegawa et al., 2006; Tono et al., 2005). Evidence of 410 uplift at subduction zones, instead, is less convincing due to locally contradicting observations (Tibi and Wiens, 2005; Tonegawa et al., 2005; Tonegawa et al., 2006; Tono et al., 2005). The transition zone structure beneath hot spots, many of which are in oceanic regions, is less clear due to the limited data set (Kind and Li, 2015). To a first approximation the discontinuity topography provides information on lateral mantle temperature variations and depends on the Clapeyron slopes of the mineral transformations giving rise to the discontinuity, although other factors such as chemistry and metastable olivine due to sluggish kinetics at low temperature may also play a role (Kirby et al., 1996). The Clapeyron slopes of mineral phase transitions can be determined by means of high pressure and high temperature experiments, but to be useful they need to be determined with high accuracy.

12 The local elevation of the 410 and depression of the 660 near the slabs (Fig. 1.2) are consistent with the opposite sign of the Clapeyron slopes of corresponding phase transitions (Helffrich and Bina, 1994; Shearer, 2000). If the Clapeyron slope of the 410 phase transition is higher than that of the 660 as some mineral physics studies proposed (Hirose, 2002; Irifune, 1998; Ito and Takahashi, 1989; Katsura et al., 2004; Katsura et al., 2003), the above mentioned observation would indicate larger lateral temperature variations at 660 km depth.

Figure 1.2 Topography of 410 and 660 km discontinuities in the region of the Japan subduction zone (modified from Tonegawa et al., 2005). (a) Depth variations of 410 km discontinuity. (b) Depth variations of 660 km discontinuity. The black curves denote the depth contours corresponding to the top surface of the Pacific Plate (PAC). Colors indicate differences from 410 km and 660 km. Red to yellow shows the elevation and pale-blue to blue shows the depression. The black ellipse indicates the uplift portion of 410 km discontinuity.

1.1.1.3 Discontinuity sharpness

The sharpness of the transition zone discontinuities, i.e. the depth interval over which a discontinuity occurs, can be determined using observations of high-frequency data such as P’P’ precursors or locally reflected and converted seismic waves because such high-frequency seismic waves can only be influenced by a high impedance contrast (equal to the product of compressional velocity and density) across a narrow discontinuity. These data are consistent with a sharp 660 (≤2 km thick) and a more diffuse 410 discontinuity which

13 can be modelled as a 7-km-wide gradual transition with a sharp jump at the end (Xu et al., 2003). The variation of the sharpness of 660 discontinuity among different areas was quite small between 2-5 km (Benz and Vidale, 1993; Tibi and Wiens, 2005; Tonegawa et al., 2005;

Yamazaki and Hirahara, 1994) while the 410 may be more variable in sharpness than 660 which ranges from 2 to 35 km (Benz and Vidale, 1993; Priestley et al., 1994; Tonegawa et al., 2005; Yamazaki and Hirahara, 1994). Two exceptions were reported by Bostock (1996) and Petersen et al. (1993) who found a sharper 410 (5-7 km) than 660 (20-30 km) from Ps conversion studies. The 520 reflector was observed in long-period SS precursor studies but absent in high-frequency P’P’ precursor observations, suggesting that the thickness of 520 discontinuity is between 10 and 50 km (Shearer, 2000). The sharpness of the discontinuities provides important information of the deep mantle. If the phase transition of the olivine system is responsible for the discontinuities, the pressure interval of the phase transition should be consistent with the sharpness of the discontinuity.