After a detailed search for events that have suitable geometry for sampling the superdome edge and possible plume, we found four events (Figure 4.17) that could be analyzed as presented below. The record section in Figure A.1a covers the epicentral distances of 95° to 110°, where each record is plotted relative to PREM. Thus, each trace should start at zero time if the Earth is PREM-like and the event was properly located. However, since these four events are small, they are not well located, nor do they have accurate origin times. But because we are only interested in their relative travel times across the array, it does not cause a serious problem. Moreover, this array has been well studied, with only minor station corrections [James et al., 2001]. Consequently, the travel time delay for stations south of about 100° by 5s is caused by earth structure, assumed to be the LLSVP. A 200 km layer with reduced S velocity of 3% produces 1s of SKS delay relative to PREM. Unfortunately, determining these delays accurately in the presence of noisy oceanic crustal events containing depth phases is difficult. In the first step, we determine or define an empirical source function which is a wavetrain most simple and common to the entire array by a cross-correlation search. The top trace was used in this case. We then generate a synthetic seismogram for a reference model, PREM, and determine the best Δ LR for each record by a grid search along with the Δ T travel time delay
aperture arrays appears to be promising in reducing such bias, but, especially when inter-station distances are large, weak correlation of waveforms between adjacent stations, owing to the site effects (Kato et al., 2000), would be disad- vantageous in identifying branches of later phases. Station corrections for the short-period waveforms (Yamazaki et al., 1996) should be still a difficulty for waveforms recorded in regional distances owing to the nature of complex wave field. Lateral heterogeneity in the transition zone velocities has been reported for the Philippine Sea region (Brudzinski et al., 1997; Tajima and Grand, 1998; Nowack et al., 1999) so that it might be crude to discuss the structure in detail with our results from a one-dimensional approach. Nevertheless, our findings of high velocity anomaly in two paths qualita- tively agree with high velocity anomalies that are found in other regional studies (e.g., Suetsugu and Nakanishi, 1987; van der Hilst et al., 1991; Fukao et al., 1992). High velocity anomaly in the transition zone beneath the Philippine Sea has been linked to the accumulation of subducted slab which do not penetrate into the lowermantle effectively (Okino et al., 1989; van der Hilst and Seno, 1993; Ohtaki and Kaneshima, 1994; Tajima and Grand, 1998). The origin of the high velocity anomalies could be both thermal and chemical con- sidering the compositional difference between the oceanic plate and the ambient mantle (e.g., Gaherty et al., 1999).
tomography also shows that the slab mantle at 90–120 km depth range has a moderately low velocity (figure 2d of Zhao et al., 2000a). Many seismic events are located within that portion of the zone, similar to those beneath Kii Peninsula, suggesting that mantle dehydration events occur there. Al- though the remnant arc (Palau-Kyushu Ridge) (Fig. 9) sub- ducting beneath Kyushu makes the thermal structure of the PHS unclear there, it is older ( ∼ 50 Ma, see Seno, 1988) than the Shikoku Basin subducting along the Nankai Trough ( ∼ 20 Ma, e.g., Okino et al., 1994). The colder slab would develop the dehydration locus of serpentine more signifi- cantly at depth than shown in Fig. 5, if the mantle part of the PHS is hydrated. Thus the state beneath Kyushu might be regarded as an advanced stage of what is beneath Kii Penin- sula.
tectonic forces that resulted into normal faulting and associated horst and graben geometry. Structural traps include the faults, anticlines and duplex etc.  one of the most common ways of identifying seismic reflections is to compare a seismic section with another section to find the regularity in different horizon of the area to be investigated. To grasp strong command on seismic interpretation and structure delineation, synthetic seismograms are frequently used to identify the reflectors.  For new prospects in any area, seismic data interpretation for hydrocarbon traps is not sufficient, further detail study like Petro physical analysis, reservoir characterization, rock physics analysis and seismic modeling is required. According to  in order to get detail information of the subsurface, velocity modeling is indeed essential. However, if the geophysical data is sparse, then structural and stratigraphic interpretation is the most suitable method to extract more information regarding petroleum system. In general, velocity increases with depth as density and overburden pressure increases, velocity in the subsurface varies in both laterally and vertically. Vertical variations are due to lithological changes of layering and increasing pressure due to increasing depth. Lateral variations are due to slow changes in density and elastic properties due to changes in lithology or physical properties. Meanwhile acoustic impedance provides useful information about the lithological successions as well as variation in different rock properties. [8, 9] Acoustic impedance variations are directly related to the lithological variations and hydrocarbons content in a reservoir formation.”
The magnetic behavior of the atoms is specified differently in the two packages. In QE, one specifies the starting magnetization (values range between -1 and 1 for all spin down and all spin up, respectively) and may choose to fix the total magnetization or allow the self-consistent cycle to determine the magnetization (typically, the desired method). Thus, for the antiferromagnetic case, one site will have a positive starting magnetization while the other site will have a negative starting magnetization. Constrained magnetization calculations may be performed where the total or atomic magnetization or direction are constrained by adding a penalty functional to the total energy. In VASP, the user specifies the initial magnetic moment per atom in units of Bohr magneton. In either case, it may be difficult to obtain a magnetic spin state, particularly an antiferromagnetic state. Possible solutions are: (1) start with a different initial structure, relaxed with different parameters or at a different pressure, (2) start with higher magnetic moments in VASP or different magnetization in QE, (3) remove the symmetry constraint on k-points, (4) relax the structure with fixed total magnetization to get the correct structure for unconstrained calculations in QE, (5) relax the structure with correct magnetic charge densities to get the correct structure, and (6) specify starting eigenvalues of the occupation matrix in QE, perhaps holding it constant for several iterations (see ‘starting_ns_eigenvalue(m,ispin,l)’ and ‘mixing_fixed_ns’).
By comparing slab locations predicted from geodynamic models based on subduction history, both plate reconstruc- tions (e.g. Bunge and Grand, 2000; Hafkenscheid et al., 2006; Zhang et al., 2010; Shepard et al., 2012) and geo- dynamic model parameters, such as slab sinking rates and mantle viscosity (e.g. Ricard et al., 1993; Lithgow-Bertelloni and Richards, 1998; Liu et al., 2008; van der Meer et al., 2010; ˇ C´ıˇzkov´a et al., 2012), can be constrained. Towards that goal, we present here a simple geodynamic model of man- tle density based on subduction history, and compare it to seismic tomography. We both visually compare along cross sections and compute formal correlations (cf. Ray and An- derson, 1994; Becker and Boschi, 2002). Our work is essen- tially an update of Steinberger (2000), which we believe is appropriate now, as both models of seismic tomography and subduction history have changed since then. We also com- pare our results with those of a simple slab sinker approach (Ricard et al., 1993; Lithgow-Bertelloni and Richards, 1998), as well as an updated slab sinker approach based on our own subduction history model and the sinking rate 1.2 cm yr −1 inferred from van der Meer et al. (2010), in order to assess whether the geodynamic model in fact leads to an improve- ment.
Low SKS birefringence may result from vertical varia- tions of the seismic anisotropy. Indeed, beneath the western Kaapvaal (Kimberley block), the Rayleigh wave azimuthal anisotropy shows a change of fast propagation directions from N–S in the crust to E–W in the mantle (Adam and Lebedev, 2012). A change with depth in the orientation of the fast direction from E–W to N–S was also detected at depths > 160 km beneath the Limpopo belt by a study associating P wave receiver functions and SKS waveforms inversion (Vinnik et al., 2012). The seismic discontinuity imaged by S wave receiver functions at ∼ 150 km depth (Wittlinger and Farra, 2007; Savage and Silver, 2008; Hansen et al., 2009) also points to vertical variations in the deformation structure within the cratonic root, since drastic changes in composition or S wave velocities with depth are not observed in neither our data set nor in that of James et al. (2004). Most cases presented in Fig. 7 show a decrease in anisotropy at 140 km depth, which would be consistent with the observations of Wittlinger and Farra (2007), but the associated gradient in seismic velocities is too weak to produce a strong impedance contrast. Peslier et al. (2010) and Baptiste et al. (2012) did
environments. In the British Isles, the main phase of tectonic activity was ∼400 Ma when the proto‐Atlantic Ocean closed. In the later stages of Atlantic rifting at ∼60 Ma, volcanic centers on the western coast of Britain created an extensive intrusive dike network in northern England and Scotland that extends across to the North Sea [Craig, 1991]. Arrowsmith et al.  found shear wave speed decreases of 1.5%, corresponding to ∼ 200 K hotter temperatures, that extended to 150 km depth under the British Isles. Tomographically determined shear wave speeds in Brazil ’ s Saõ Francisco craton [Schimmel et al., 2003] are 1% slower than the IASP91 reference velocity model [Kennett and Engdahl, 1991] in a columnar structure extending from about 100 km to 400 km depth beneath the Paraná Basin. The structure is attributed to activity of the Tristan de Cunha plume, active at ∼130 Ma, which modified the pre‐existing, 250 km deep, cratonic structure. While the anomalies discussed above all affect the lithosphere and asthenosphere under continents, oceanic
The second period of recording took place from the 19th January to the 21st March 1978. It was decided to reoccupy the same line, but the philosophy of the operation was different. Twenty nine seismic recorders were used, comprising ten N6's, four OBS's, and fifteen H 1 4 ' s . Of the N6 instruments, seven were deployed as mine monitors at the seven major open-cut coal mines, and three as control stations and tie points to the previous QHC line, occupying exactly the same sites as the stations QHC02, QHC09, and QHC16. The four OBS's were used in two separate experiments. The first was designed to aid the time term analysis, by placing the recorders on radial lines from the northern mines (Goonyella, Peak Downs, and Saraji), at approximately the same radius from these mines, on either side of Moura. The second experiment involved the deployment of the OBS instruments as a short refraction array to the east and west of the Moura and Kianga Mines, to determine the velocity structure of the Bowen Basin sediments directly beneath these mines.
1. Interseismic surface displacement normal to the Alpine Fault measured by GPS geodesy (Beavan et al., 1999; Beavan et al., 2004) shows a superposition of a short- wavelength signature of ca. 20 km whose amplitude is about half of the available normal convergent displacement across the plate boundary, and a longer wavelength signal (ca. 100 km) which has been interpreted as an elastic response to deformation of mantle lithosphere beneath the collision zone. The short-wavelength signal is similar to that shown in Model 2 (Fig. 3(b)). Beavan et al. (1999, 2004) have inverted the surface displacement and solved for slipping zones beneath the (currently locked) Alpine Fault. Their most recent best-ﬁt model predicts slip along the down-dip extension of the Alpine Fault, with the upper locked part of the fault ending at a depth of 7 km (Beavan et al., 2004). This is similar to the depth of transition from “locked” to “creeping” behaviour seen in Model 2 (Fig. 2(c) and (d)). Given the similarity, we interpret the slipping patch de- scribed in Beavan et al. (2004) as a zone of enhanced ductile creep, caused by a combination of thermal weakening and the interseismic response to seismic stress loading from the fault itself.
resolution tests (CRTs) to evaluate the resolution of our 3-D Vp and Vs models. Velocity perturbations of +/–2% were assigned to the grid nodes adjacent to each other, and then synthetic travel times were calculated for the checkerboard model. Figure 3 shows the plan views of the CRT results for Vp and Vs images at each depth, indicating that the resolution of the tomographic images is high at depths of between 25 and 90 km. The CRT results for the Vp and Vs models are consistent with each other (Figure 3). Although few earthquakes occurred in the lower crust and no local earthquakes occurred in the upper mantle under the study region, numerous rays of head waves (Pn, Sn) from the crustal earthquakes were refracted at the Moho discontinuity, allowing the near vertical rays from the teleseismic events to sample the lower crust and upper mantle. The checkerboard pattern in the lower crust and uppermost mantle was generally recovered, and therefore we believe that the seismic ve- locity variations in the study region were resolvable fea- tures.
Figure 2 Shear-wave splitting versus epicentral distance. Circles with error bars show observations and estimated uncertainty. Solid lines show predictions for a 500-km-deep source in models with a fast horizontally polarized shear wave: trace a, 4% transverse- isotropy in the uppermost 210 km of upper mantle; b, 5% anisotropy in a subducted slab that extends to a depth of 660 km; c, 5% anisotropy con®ned to a 100-km-thick layer immediately beneath the `660'; d, anisotropy that grades from 3% to 1% between the `660' and 900 km; e, 2% anisotropy in 100-km-thick layers above and below the `660', and 1% anisotropy in a layer between 760 km and 900 km; f, 4% anisotropy in the uppermost 210 km, 2% anisotropy in 100-km-thick layers above and below the `660', and 1% anisotropy in a layer between 760 km and 900 km; g, anisotropy that grades from 2.5% to 1.5% in a layer between 760 km and 900 km. Although there is some ambiguity as to the best model, only models with anisotropy in the lowermantle can explain the large splitting observations.
Figure 8. (a, b) Input P- (a) and S-wave (b) velocity anomalies (%) along a vertical cross-section through the model N10 oriented such that the two plumes are positioned approximately under Afar and MER. The lo- cation of the cross-sections (black line) is shown in Fig. 1b. The structure on the right represents the synthetic MS plume, the structure on the left the LS plume. (c, d, g, h) Vertical cross-sections through the recovered undamped P- (c) and S-wave (d) models and the damped P- (g) and S-wave (h) models. (e, f, i, j) Vertical cross-sections through the undamped (e) and damped (i) NEAR-P15 model and the undamped (f) and damped (j) NEAR-S16 model. The spacing between the contours is 0.25 % for P-wave models and 0.50 % for S-wave models. The undamped models (c-d) image the tail of the MS plume, but the LS plume recovery is almost completely lost. The damped recovered models (g-f) are able to resolve the MS plume structure and the head of the LS plume, but with relatively subdued amplitudes. The scale of the recovered structures is quite similar to that of the imaged features.
An upper mantleseismic profile for Fennoscandia was constructed from Russian explosion recordings made at NORSAR. The upper mantle P-wave velocity distribution beneath Scandinavia and western Russia was determined from this profile to be the same as that beneath the Canadian Shield (Masse, 1973), with discontinuities at depths of approximately 74, 107, 328, 430 and 710 km, and a gradual increase in velocity in the depth range 590 to 710 km. Although including a rather complex structure between 90 and 110 km, MA is a step-like model consisting largely of zones of constant velocity and first-order discontinuities. This simple structure entails simple travel times and synthetics with cross-overs near 2250 km (20.2°) and 2700 km (24.3°). At distances beyond the second cross-over, the direct, S710, S430 and S328 branches become well separated. The S710 branch has very large amplitudes (clipped on the plot) due to the strong positive velocity gradient immediately above the "710-km" discontinuity. Furthermore, the slowness of the S710 branch is too small for the B branch. The S320 branch matches the velocity of the D branch, but differs in amplitudes. This could be remedied by placing a more positive velocity gradient above the "320-km" discontinuity. The amplitudes and velocity on the S420 branch would fit the C branch, but the timing relative to the first arrival would need attention. MA generates a complex succession of multiples (2S107, 2S328, 2S430, 2S710) which appear to match quite well the style of the multiples of the observed data but the cusps of the presumed branches should be moved to smaller distances. Unfortunately, the poor misfit of the waveform data with the synthetic first arrivals exclude MA as a possibility for future refinement modelling.
How far one can trust the conclusions made? It seems im- possible yet to give a quantitative estimation for each value received, for there is no quantitative estimations of the data on temperatures and pressures used. We carried out calcula- tions with several models of distributions of thermodynamic parameters, which were presented in different publications (Stacey, 1972; Zharkov, 1983; Brown and Mussett, 1984), for the same model of electrical conductivity. No principal differences were found. Chosen model of global distribution of electrical conductivity in the middle and lowermantle meets fairly well (see Fig. 5) the results of laboratory stud- ies performed by (Shankland et al., 1993). That is why we preferred the very model. Assurance in rightfulness of our results would strengthen if they don’t contradict the modern views about a mantlestructure. Let us turn to Fig. 4. The
plates have been made through a number of seismological studies at different locations beneath the Pacific Ocean (Ga- herty, 1999; Tan and Helmberger, 2007; Kawakatsu, 2009; Kumar and Kawakatsu, 2011; Kumar et al,. 2012; Rychert and Shearer, 2011; Schmerr, 2012, etc.). Logistic problems and the extreme costs compared to land surveys impose se- vere constraints on obtaining similar information about the nature of other MORs. Another factor which hampers body- wave observations from the oceanic data is the water rever- berations that mostly contaminate the vertical components of the seismograms of the ocean bottom seismometers (Ku- mar et al., 2012). In this study, we attempt to investigate the seismicstructure using the available data from five seismo- logical stations situated on the islands located close to the mid-oceanic ridges. Although the crustal structure may not truly represent the nature of the oceanic plate due to possi- ble influence of the islands, the deeper structure is devoid of such effects.
We conducted a seismic survey consisting of MCS and wide-angle seismic measurements using dense ocean bot- tom seismograph (OBS) receivers to update the velocity model of the SB, which is a typical backarc basin. Since Bouguer gravity anomaly data indicate that the SB crustal thickness decreases toward the south (e.g. Ishihara and Koda, 2007), we located seismic lines at the north, center, and south of the basin to characterize this variation (Fig. 1). The direction of the seismic lines is roughly east-west, par- allel to the direction of the backarc basin spreading, with a half-rate of 2.3–4.7 cm/year in 27–20 Ma except for the last NE-SW spreading stage, which occurred at a slower rate of 2–3 cm/year in 20–15 Ma (Okino et al., 1999). A tuned array of 36 airguns with a total volume of 8,040 inch 3
structure contains nominally no non-bridging oxygens (NBO) per tetrahedron (T) (NBO/T=0), meaning that each oxygen atom connects two silicon atoms forming a network of silica tetrahedra. The incorporation of network modifiers, such as magnesium oxide, increases the number of non-bridging oxygens per tetrahedron, e.g. nominal NBO/T=2 for MgSiO 3 . Note that the concept of NBO/T breaks down at high
Seismic refraction and multi-channel seismic reflection surveys were conducted on the northeastern Hawaiian Arch to examine the effect of hotspot volcanism on the seismicstructure of the crust and uppermost mantle. The crustal thickness deduced from the refraction data was typical for oceanic crust, which suggests that magmatic underplating does not occur, at least in our survey area, although the crustal seismic velocity may be influenced by flexure of the lithosphere on the arch. We identified high P-wave velocities (~ 8.65 km/s) in the uppermost mantle parallel to the paleo-seafloor spreading direction, which indicates that the shallower mantlestructure immediately below the Moho preserves the original structure formed at a mid-ocean ridge. Moreover, we observed wide-angle reflection waves at large offsets in ocean-bottom seismometer records. The travel time analysis results showed that these waves were reflected from mantle reflectors at depths of 30–85 km below the seafloor, which are considered to represent hetero- geneities consisting of frozen melts created during the cooling of the plate.