Time-domain observations of a slow precursor to the 1994 Romanche Transform earthquake

Science

Volumn 274, Issue 5284, 1996, Pages 82-85

  McGuire, Jeffrey J ^a   Ihmlé, Pierre F ^b   Jordan, Thomas H ^a  

^a MASSACHUSETTS INSTITUTE OF TECHNOLOGY   (United States)

^b INSTITUT DE PHYSIQUE DU GLOBE DE PARIS   (France)

Author keywords

[No Author keywords available]

Indexed keywords

ARTICLE; EARTHQUAKE; OCEANIC REGIONS; PRIORITY JOURNAL;

EARTHQUAKE PRECURSOR; EARTHQUAKES; ROMANCHE TRANSFORM EARTHQUAKE 1994;

ATLANTIC, ROMANCHE TRANSFORM;

EID: 0029664227     PISSN: 00368075     EISSN: None     Source Type: Journal    
DOI: 10.1126/science.274.5284.82     Document Type: Article

References (27)

1. H. Kanamori and J. Cipar, Phys. Earth Planet. Inter. 9, 128 (1974); see also I. L. Cifuentes and P. G. Silver, J. Geophys. Res. 94, 643 (1989).
2. H. Kanamori and J. Cipar, Phys. Earth Planet. Inter. 9, 128 (1974); see also I. L. Cifuentes and P. G. Silver, J. Geophys. Res. 94, 643 (1989).
3. T. H. Jordan, Geophys. Res. Lett. 18, 2019 (1991).
4. P. F. Ihmlé, P. Harabaglia, T. H. Jordan, Science 261, 177 (1993).
5. P. F. Ihmlé and T. H. Jordan, ibid. 266, 1547 (1994).
6. S. Kedar, S. Watada, and T. Tanimoto [J. Geophys. Res. 99, 17893 (1994)] found significant amplitude and phase anomalies for certain low-frequency spheroidal modes that could be accounted for by a slow precursor to the Macquarie Ridge earthquake [see also J. Park, Geophys. Res. Lett. 17, 1005 (1990)], but they rejected this possibility because no precursor was evident as time-domain signals preceding the P waves at high-gain seismic stations. Such signals would be below the noise level if the time function of the precursor were smooth and of sufficiently long duration (3).
7. S. Kedar, S. Watada, and T. Tanimoto [J. Geophys. Res. 99, 17893 (1994)] found significant amplitude and phase anomalies for certain low-frequency spheroidal modes that could be accounted for by a slow precursor to the Macquarie Ridge earthquake [see also J. Park, Geophys. Res. Lett. 17, 1005 (1990)], but they rejected this possibility because no precursor was evident as time-domain signals preceding the P waves at high-gain seismic stations. Such signals would be below the noise level if the time function of the precursor were smooth and of sufficiently long duration (3).
8. D. Agnew and J. Berger, J. Geophys. Res. 83, 5420 (1978); G. C. Beroza and T. H. Jordan, ibid. 95, 2485 (1990).
9. D. Agnew and J. Berger, J. Geophys. Res. 83, 5420 (1978); G. C. Beroza and T. H. Jordan, ibid. 95, 2485 (1990).
10. The hypocenter location given by the National Earthquake Information Center (NEIC) is 94:03:14:04: 30:07.7 UT (universal time), 1.08°S, 23.9°W, h = 10 km; and the Harvard centroid location [A. M. Dziewonski, G. Ekström, M. P. Salganik, Phys. Earth. Planet. Inter. 86, 253 (1994)] is 94:03:14:04:30:33.1 UT, 0.88°S. 23.0°W, focal depth = 15 km.
11. w-7.0 event is 94:09:01:15:16:00.6 UT, 40.59°N, 125.78°W, focal depth = 15km.
12. We picked 10 arrivals for subevent A and 11 for subevent B, obtaining at least two B - A differential travel times in each azimuthal quadrant (see Table 1).
13. We performed relocations by using the clustered-event algorithm of T. H. Jordan and K. A. Sverdrup [Bull. Seismol. Soc. Am. 71, 1105 (1981)], which yields relative locations that are independent of common path anomalies. We relocated subevents A and B together with all events having 30 or more P-wave arrival times in the International Seismological Centre catalog from 1964 to 1987 and in the Preliminary Determination of Epicenters catalog from 1990 to 1995. All event depths were fixed at 10 km. The hypocentroid of this seismicity cluster has been shifted 12 km in the direction N30°E to align the seismicity with the plate boundaries observed in the gravity field. Although the arrival time data for the 2-year period 1988 through 1989 were unavailable, the PDE catalog shows no events in the aseismic region between 22.3°W and 23.3°W.
14. D. Sandwell, Eos 76, 149 (1995).
15. We determined the moment tensor in 1-mHz bands from 1 to 11 mHz, using the free-oscillation inversion method described by M. A. Riedesel, T. H. Jordan, A. F. Sheehan, and P. G. Silver [Geophys. Res. Lett. 13, 609 (1986)]; no significant frequency dependence of the source mechanism was observed, which implies that the slow component of the 1994 Romanche Transform earthquake had a radiatton pattern similar to that of the main shock. The mechanism labeled LF in Fig. 5 is the average across the frequency band 3 to 6 mHz. The source mechanisms of subevents A and B, also shown in Fig. 5, were determined by waveform analysis. They are similar but not identical; for example, their long-period P-wave polarities are reversed at Naña, Peru (NNA).
16. Although rupture velocities of shallow-focus earthquakes have been known to exceed the local shearwave speed [R. Archuleta, J. Geophys. Res. 89, 4559 (1984)], they are rare. Typical rupture velocities of shallow-focus earthquakes are less than 3.5 km/s (14).
17. C. H. Scholz, The Mechanics of Earthquake Faulting (Cambridge Univ. Press, Cambridge, 1990).
18. E. Bonatti, et al., J. Geophys. Res. 99, 21779 (1994); R. C. Searle, M. V. Thomas, E. J. W. Jones, Mar. Geophys. Res. 16, 427 (1994), The morphology of the western portion of the Romanche Transform is extremely complex, exhibiting multiple paleotransform valleys that resulted from past changes in plate motions. The seismic gap on the main transform trace between 22.3°W and 23.3°W (Fig. 4) corresponds to a bathymetric high, which Searle et al. attributed to transpression caused by the northeastward bending of the fault trace at the western end of the gap. Locking of the main trace in this region could explain the offset of the B-subevent rupture to the north.
19. E. Bonatti, et al., J. Geophys. Res. 99, 21779 (1994); R. C. Searle, M. V. Thomas, E. J. W. Jones, Mar. Geophys. Res. 16, 427 (1994), The morphology of the western portion of the Romanche Transform is extremely complex, exhibiting multiple paleotransform valleys that resulted from past changes in plate motions. The seismic gap on the main transform trace between 22.3°W and 23.3°W (Fig. 4) corresponds to a bathymetric high, which Searle et al. attributed to transpression caused by the northeastward bending of the fault trace at the western end of the gap. Locking of the main trace in this region could explain the offset of the B-subevent rupture to the north.
20. We obtained the spectral estimates in Fig. 6 using the procedures described in (2-4) We measured the spheroidal free oscillations from vertical-component seismograms in the band from 1 to 19 mHz, using the methods of P. G. Silver and T. H. Jordan [Geophys. J. R. Astron. Soc. 70, 755 (1982)] and M. A. Riedesel and T. H. Jordan [Bull. Seismol. Soc. Am. 79, 85 (1989)]. Measurements of first-orbit Rayleigh waves (1 to 10 mHz) and long-period body wave trains (10 to 50 mHz) were obtained from vertical-component seismograms by the methods of Ihmlé et al. (3). In all cases, synthetic seismograms were used to account for radiation-pattern and propagation effects. The synthetics were computed by mode summation from the Harvard CMT (7) and the degree-12 aspherical earth structure of W.-J. Su, R. L. Woodward, and A. M. Dziewonski [J. Geophys. Res 99, 6945 (1994)]. We also corrected fundamental modes above 7 mHz for smaller scale heterogeneity using the degree-36 phase-velocity maps of G. Ekström, J. Tromp, and E. W. Larson [Eos 74, 438 (1993)].
21. We obtained the spectral estimates in Fig. 6 using the procedures described in (2-4) We measured the spheroidal free oscillations from vertical-component seismograms in the band from 1 to 19 mHz, using the methods of P. G. Silver and T. H. Jordan [Geophys. J. R. Astron. Soc. 70, 755 (1982)] and M. A. Riedesel and T. H. Jordan [Bull. Seismol. Soc. Am. 79, 85 (1989)]. Measurements of first-orbit Rayleigh waves (1 to 10 mHz) and long-period body wave trains (10 to 50 mHz) were obtained from vertical-component seismograms by the methods of Ihmlé et al. (3). In all cases, synthetic seismograms were used to account for radiation-pattern and propagation effects. The synthetics were computed by mode summation from the Harvard CMT (7) and the degree-12 aspherical earth structure of W.-J. Su, R. L. Woodward, and A. M. Dziewonski [J. Geophys. Res 99, 6945 (1994)]. We also corrected fundamental modes above 7 mHz for smaller scale heterogeneity using the degree-36 phase-velocity maps of G. Ekström, J. Tromp, and E. W. Larson [Eos 74, 438 (1993)].
22. We obtained the spectral estimates in Fig. 6 using the procedures described in (2-4) We measured the spheroidal free oscillations from vertical-component seismograms in the band from 1 to 19 mHz, using the methods of P. G. Silver and T. H. Jordan [Geophys. J. R. Astron. Soc. 70, 755 (1982)] and M. A. Riedesel and T. H. Jordan [Bull. Seismol. Soc. Am. 79, 85 (1989)]. Measurements of first-orbit Rayleigh waves (1 to 10 mHz) and long-period body wave trains (10 to 50 mHz) were obtained from vertical-component seismograms by the methods of Ihmlé et al. (3). In all cases, synthetic seismograms were used to account for radiation-pattern and propagation effects. The synthetics were computed by mode summation from the Harvard CMT (7) and the degree-12 aspherical earth structure of W.-J. Su, R. L. Woodward, and A. M. Dziewonski [J. Geophys. Res 99, 6945 (1994)]. We also corrected fundamental modes above 7 mHz for smaller scale heterogeneity using the degree-36 phase-velocity maps of G. Ekström, J. Tromp, and E. W. Larson [Eos 74, 438 (1993)].
23. We obtained the spectral estimates in Fig. 6 using the procedures described in (2-4) We measured the spheroidal free oscillations from vertical-component seismograms in the band from 1 to 19 mHz, using the methods of P. G. Silver and T. H. Jordan [Geophys. J. R. Astron. Soc. 70, 755 (1982)] and M. A. Riedesel and T. H. Jordan [Bull. Seismol. Soc. Am. 79, 85 (1989)]. Measurements of first-orbit Rayleigh waves (1 to 10 mHz) and long-period body wave trains (10 to 50 mHz) were obtained from vertical-component seismograms by the methods of Ihmlé et al. (3). In all cases, synthetic seismograms were used to account for radiation-pattern and propagation effects. The synthetics were computed by mode summation from the Harvard CMT (7) and the degree-12 aspherical earth structure of W.-J. Su, R. L. Woodward, and A. M. Dziewonski [J. Geophys. Res 99, 6945 (1994)]. We also corrected fundamental modes above 7 mHz for smaller scale heterogeneity using the degree-36 phase-velocity maps of G. Ekström, J. Tromp, and E. W. Larson [Eos 74, 438 (1993)].
24. 2 measure of data misfit and a quadratic form measuring the smoothness of the source time function, subject to the constraint that the source time function be nonnegative. The smoothing varied from high values before the subevent-A origin time (110 ≤ < t < 0 s), which ensured that the precursor did not generate significant high-frequency arrivals, to low values during the main-shock phase of the rupture (16 ≤ t < 45 s), when the high-frequency amplitudes were largest; intermediate values of smoothing were assumed between the initiation of subevent A and the initiation of subevent B (0 ≤ t < 16 s) and in the interval after the main shock (40 ≤ t < 200 s).
25. 17 Nm and a duration of about 2 s for an earthquake the size of the 1994 Romanche Transform main shock.
26. 2.
27. We thank H. Webb for assistance with the gravity data and G. Ekström, J. Tromp, and E. Larson for the use of their unpublished phase-velocity maps. We are grateful to D. Wiens and an anonymous reviewer for helpful comments that improved the manuscript. Sponsored by NSF under grant EAR-9305081 and by NASA under grant NAG5-1905. P.F.I, was supported in part by the Swiss National Science Foundation.