No abstract

Mthe scahed. stibhotton s elocit v/at tentation st ruectuLre is essential input bor predictive propation nmodels. To estiratec this structure. bottoin-mounted soutces and receivers %%ere used to make mecasuirments of shear ;and compressional wave ptopitgatioti in shallow water sedimnents of the continental shell'. usually where boreholes and high-resolution retlecuion profleks give substantial suipporting geologic infoaniation ;about the suhsuu1'ace. This colocation provides an opportunity ito comipate seismically determined estimates oft physical properties of' the seabed with the "gound truth-properties. Measuirenments were miade iti 1986 %kith source/detector otffsets up to 2(M) mn producing shear wave velocity versus depth proliles of, Ithe uippet ( So it) fiiii h seabed l;and P' wase proliles to lesser depths). Measuremients tn 1988 were made with smailer source devices designed to emphasiz.e higher f'requencies and recoirded by an array of 3(0 sensors spaced atf I-m intervals to imiprove spatial samnpling and resolution of s hallow structure. These investigations with shear waves have shown that significa~nt laiteraml and vertical variations in the physical properties of the shallow seabed are cotumion and tare principally created b.% erosional and depositional processes assoctated %%ith glacial cycles and sea level oscillations during the Quaternarv. When the seabed structure is relatively uniform over the length of the profiles. the shear wave fields aire well ordered, and the iuatching of the data with full waveform synthetics has been suceessl'ul. producing velocity/attenuation models consistent with the subsurface lithology indicated by coring results. Both body waves and LILow-frequency sound propagation in shallow water environments is not restricted to the water __ column but also involves the subbottom. Thus. as well as being important for geophysical description of the seabed. subbottom velocity/attenuation structure is essential input for predictive propagation models. To estimate this structure, bottom-mounted sources and receivers were used to make measurements of shear and compressional wave propagation in shallow water sediments of the continental shelf, usually where boreholes and high-resolution reflection profiles give substantial supporting geologic information about the subsurface. This colocation provides an opportunity to compare seismically determined estimates of physical properties of the seabed with the -'ground truth" properties. Measurements were made in 1986 with source/detector offsets up to 200 m producing shear wave velocity versus depth profiles of the upper 30-50 m of the seabed (and P wave profiles to lesser depths). Measurements in 1988 were made with smaller source devices designed to emphasize higher frequencies and recorded by an array of 30 sensors spaced at I-in intervals to improve spatial sampling and resolution of shallow structure. These investigations with shear waves have shown that significant lateral and vertical variations in the physical properties of the...

The interleaving of parallel isotropic lamellae of contrasting mineralogical composition makes almost all marine sediments anisotropic, the form of anisotropy being transverse isotropy with a vertical axis of symmetry. Conventional marine seismic experiments, however, cannot quantify the anistropy because they do not record unconverted shear waves. In 1986, Rondout Associates, Inc. (RAI) and Woods Hole Oceanographic Institution (WHOI) recorded direct shear waves in shallow marine sediments by using a newly developed ocean-bottom shear source and a multicomponent on-bottom receiver. No single isotropic model could be adequately fit to the data, implying anisotropy. The seismic experiment was conducted in 21 m deep water about 10 km east of the New Jersey coast. In this paper, we describe the anisotropy in the top 50m of marine sediments beneath two of the RAI/WHOI refraction profiles. We use an anisotropic reflectivity program to produce synthetic seismograms to estimate the five independent elastic stiffnesses necessary for describing the transverse isotropy. Our synthetics fit the vertical and two horizontal components of the data for both profiles. The two intersecting refraction profiles are 150 and 200m long. These profiles are not long enough to constrain compressional wave velocities and anisotropy, but are quite adequate to find the shear wave anisotropy. A nearby drill hole showed that the sediments are interbedded silty clays, clays, and sands. The data require low shear velocities (<400m s-') and low Q, (400) in about the top 30 m of the sediments. In the top 10 m of the sediments, silty clay exhibits -12-15 per cent anisotropy for shear waves.

When velocity varies laterally as well as with depth, an exact Kirchhoff depth migration requires that rays be traced from each depth point in the section to each source/receiver location. Because such a procedure is prohibitively expensive, Kirchhoff migration is usually carried out by using a velocity function that depends only on depth. This paper introduces a new method, based on Fermat’s principle, which is a compromise between these two extremes. The slowness (reciprocal velocity) function is written as the sum of two functions, the first of which is large and depends only on depth, while the other is small and varies both with depth and position along the line. Raypaths are traced for the first slowness function and are used to calculate migration curves. For each depth point these same raypaths are used to calculate traveltime perturbations due to the laterally varying part of the slowness. The traveltime perturbations are added to the migration curve to obtain an approximation to the exact migration curve. More precisely, suppose that the slowness function can be written in the form [Formula: see text] where [Formula: see text] Using a(z), we generate a table giving two‐way traveltime T as a function of scatterer depth z and surface offset Δ, and a raypath table which gives the ray offset ξ as a function of scatterer depth, surface offset, and ray depth [Formula: see text] For fixed z, T(Δ, z) is a migration curve and conventional Kirchhoff migration of zero‐offset reflection data ψ(x, t) is performed by summation along such curves: [Formula: see text] by [Formula: see text] where W is a weighting factor. The raypath table is used to calculate the traveltime perturbation [Formula: see text] where the integral is taken over the unperturbed raypath. For fixed x and z the new migration curve is [Formula: see text] and migration is per formed by [Formula: see text] This new migration scheme is much less expensive than the exact Kirchhoff scheme because only one set of rays need be traced. Numerical tests have shown that this scheme works surprisingly well even when the lateral variation of velocity is large.

Samples of ocean bottom noise in the frequency band 0.003 to 5 Hz are analyzed for coherency and amplitude and phase relationships among pressure and the three components of particle motion. Data available from the Columbia‐Point Arena ocean bottom seismic station (38° 09.2′N, 124° 54.4′W) provide examples of different noise conditions. Coherent energy peaks near 0.14, and 0.06 Hz suggest fundamental mode Rayleigh wave motion propagating shoreward. Coherent energy near .30 Hz appears to be variable. Pressure variations near 0.01 Hz and lower frequency correlate with wave heights along the California coast and appear to produce forced deformation of the bottom.