Seismic arrays, first introduced in the late 1950s to detect underground nuclear explosions, have helped to improve our knowledge about the structure of the Earth for the last 40 years. During these years, numerous array processing methods have been developed that use the high signal coherence and accurate timing of array data to generate high-resolution images of Earth structure. Here, we present an overview of resolution issues related to seismic array studies of Earth structure by first introducing basic array processing techniques and then discussing more advanced techniques applied to array data recently. The increase of seismic stations deployed in experiments or permanently in many regions of the globe allows a much denser sampling of the seismic wavefield. This dense sampling enables the adaptation of controlled source analysis techniques for the study of Earth structure using earthquakes with higher resolution than previously possible. Here we will discuss different migration methods of teleseismic data that use the incidence angle information of scattered arrivals to obtain images of Earth structure. Finally, we show data examples how these methods can be used to increase our knowledge of the structure of the Earth’s deep interior.

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Improving Seismic Resolution Through Array Processing Techniques

We analyze two-ship multichannel seismic wide aperture profiles (WAP) acquired on the East Pacific Rise (EPR) between 9°30′C and 10°N in May and June 1985. For offsets between 2 and 5 km, most of the WAP common depth point (CDP) gathers exhibit a strong retrograde, reflection-like seismic arrival that corresponds to rays turned within a very strong velocity gradient zone underlying the uppermost crustal layer 2A. Stacking of this postcritical arrival for hundreds of CDP gathers along the WAP lines yields remarkably clear, and virtually continuous, images of the bottom of layer 2A for several tens of kilometers along and across the EPR. An innovative combination of this image with others of the seafloor (SF) and of the axial magma chamber (AMC) provides an unusually clear composite picture of young upper crust. We combine these images with detailed seismic velocity information, also obtained from the CDP gathers, to achieve an accurate determination of the layer 2A thickness in the area. Along the EPR axis between 9°30°N and 10°N, layer 2A shows a relatively constant thickness between 100 and 150 m (80–120 ms two-way travel time). Across the axis, however, a line near 9°30′N indicates that layer 2A thickness rapidly increases by approximately a factor of 2 within 2–4 km of the axis and remains nearly constant afterwards. Thus the final thickness of layer 2A is attained before the development of large-scale tectonic faulting, strongly suggesting that the thickening of layer 2A away from the axis is due to successive episodes of volcanic activity increasing the thickness of the uppermost crustal section from the top, rather than tectonic fracturing lowering the uppermost crustal seismic velocities and pushing the bottom of layer 2A downwards as the crust moves away from the axis.

Seismic imaging of oceanic layer 2A between 9°30′N and 10°N on the East Pacific Rise from two‐ship wide‐aperture profiles

We analyze a dense grid of seismic normal incidence and wide-angle data collected over an active back arc spreading center in the Lau Basin of the southwest Pacific. Our survey area covers the whole of a 35-km-long morphological segment and a small overlapping spreading center. A bright reflector, which is coincident with a velocity inversion at a depth of 3.2 ± 0.2 km below the seafloor, is observed on every one of the 40 across-axis profiles. We interpret this reflector as being the roof of a crustal magma chamber. The widest magma chamber reflector occurs beneath the overlapping spreading center, where it extends up to 4 km, being imaged beneath both ridges and, in places, beneath the overlap basin. Elsewhere the width of the reflector varies between 0.6 and 2.3 ±0.4 km. The narrowest reflectors are observed beneath deviations (devais) of the ridge axis. In addition to being observed on each of the kilometer-spaced across-axis profiles the reflector is also seen as a continuous event for at least 10 km on a profile along the ridge axis. The reflection coefficient of the event is between −0.34 and −0.65. This requires the melt at the top of the chamber to have a P wave velocity less than 2.7 km/s if the interface is planar or less than 3.8 km/s if the interface is layered.

Seismic mapping of a magma chamber beneath the Valu Fa Ridge, Lau Basin

SUMMARY Here we investigate the thermal structure of the mantle beneath the eastern Branch of the East African Rift system in Kenya and Tanzania. We focus on the structure of the mantle transition zone,asdelineatedbystackingofreceiverfunctions.Thetopofthetransitionzone(the410km discontinuity) displays distinctive topography, and is systematically depressed beneath the rift in Kenya and northern Tanzania and adjacent volcanic fields. This depression is indicative of a localized ! 350 " C thermal anomaly. In contrast, the bottom of the transition zone (the 660km discontinuity) is everywhere depressed. This region-wide depression is best explained as a Ps conversion from the majorite‐perovskite transition of anomalously warm mantle. We interpret this structure of the transition zone as resulting from the ponding of a mantle plume (possibly the deep-mantle African Superplume) at the base of the transition zone, which then drives localized thermal upwellings that disrupt the top of the transition zone and extend to shallow mantle depths beneath the rift in Kenya and northern Tanzania.

http://africaarray.psu.edu/publications/pdfs/Huerta_et_al_Mantle.pdf

Mantle transition zone structure beneath Kenya and Tanzania: more evidence for a deep-seated thermal upwelling in the mantle

[1] We investigate anisotropy surrounding a continental back arc spreading region, the Central Volcanic Region (CVR), in the North Island of New Zealand and present a simple method for spatial averaging to examine heterogeneous anisotropy. S phases from local earthquakes yield consistent trench-parallel fast directions (ϕ) in the southern, compressional region but varied directions in the northern, extensional region. In the forearc east of the CVR, ϕ is nearly trench-parallel, suggesting shear in the mantle or trench-parallel flow. In the western CVR, local phases from 50 to 80 km depth give ϕ parallel to extension (120°), and phases from 150 to 220 km depth yield similar ϕ (150°). These results suggest asthenospheric flow with olivine a axes oriented in the extension direction down to 100–150 km depth. In the eastern CVR, local phases show mixed ϕ, suggesting that it is a transition between the western CVR and the forearc. Trench-parallel ϕ for shallow local events within the CVR may be caused by fluid-filled cracks. Trench-parallel ϕ for deeper local events and for SKS phases can be explained by standard olivine fabrics developed by lithospheric shearing, by trench-parallel flow, and by fossil anisotropy in the subducting oceanic lithosphere. Alternatively, in the CVR they could be caused in part by hydrous minerals at the slab-mantle upper interface, which could cause olivine a axes to orient parallel to the trench due to a maximum shear stress oriented perpendicular to the trench along the slab. We suggest that trench-perpendicular flow in the extensional region helps to drive fluids away from the slab, while in the compressional region, trench-parallel flow fails to distribute the fluids, explaining the location of changes in geophysical properties.

https://agupubs.onlinelibrary.wiley.com/doi/pdf/10.1029/2003JB002932

Anisotropic structure under a back arc spreading region, the Taupo Volcanic Zone, New Zealand

Geophysicists commonly recognize that a sizeable portion of a reflector is involved in causing a reflection as seen on a seismic trace, but the areal extent is usually not calculated and hence not appreciated. Most commonly, concepts are simply transferred from classical physical optics and called Fresnel-zone effects. Confusion sometimes results due to the failure to distinguish whether incident waves are plane or spherical. An excellent explanation (but one which does not use the term “Fresnel zone”) is given by Hagedoorn (1954), from which Figure 1 is taken.

R. E. Sheriff

Papers

Nomogram for Fresnel-zone calculation