The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol

Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 1

The Hawaii-2 Observatory: Observation of Nanoearthquakes

Rhett Butler The IRIS Consortium, Washington DC

Introduction

The Hawaii-2 Observatory (H2O, http://www.iris.edu/GSN/) was established on the seafloor between Hawaii and California (Figure 1) in 1998 using the retired AT&T Hawaii-2 undersea telephone cable for power and telemetry (Butler et al., 2000; Duennebier et al., 2002, Pettit et al., 2002; Harris and Duennebier, 2002). This is the first undersea seismic station in the Global Seismographic Network. The on-site instrumentation includes a broadband Guralp CMG-3 and 4.5 Hz Geospace HS1 geophones, buried 0.5 m into the seafloor mud, and an OAS E2PD hydrophone 0.5 m above the seafloor. All of the data are natively sampled at 160 sps continuously and sent via the Hawaii-2 cable to Oahu where the data are distributed by the University of Hawaii to IRIS. Data are available in real-time via a Live Internet Seismic Server (LISS) server, and are being used by the Pacific Tsunami Warning Center and the National Seismic Network. In early 2002, the Ocean Drilling Program drilled a borehole through 29±1 m of sediment into basalt about 1.5 km northeast of the current H2O instrumentation as a site for a future borehole sensor (Stephen et al., 2003).

H2O was primarily sited to fill a large gap in global coverage between Hawaii and California, and not to monitor any particular local or regional seismicity. However, in addition to recording seismic body and surface waves from global and circum-Pacific earthquakes, a diverse range of marine seismoacoustic phenomena are also seen. These include whales and high-frequency lithospheric P and S waves (Butler and Duennebier, 2000), oceanic noise (Bromirski, 2000), seismoacoustic T waves (Butler and Lomnitz, 2002; Butler, 2003), shipping noise, indigenous biological noise, and other signals of unknown source.

The H2O data are routinely monitored for quality control. Because of the broadband nature of the instrumentation and the dominance of noise in the microseismic band (0.1-2 Hz), the data are also viewed on a rectified high-pass filtered channel to see high- frequency events. At high frequencies (>10 Hz) the H2O site is one of the quietest sites in the Global Seismographic Network, with noise levels within 10 dB of the USGS low noise model (Peterson, 1993). During October 2001 an interesting pattern of arrivals was seen that merited further investigation: a small arrival followed within the hour by a curious triplet of arrivals. These data sets were requested from the IRIS DMC for further analysis. Looking at the full waveforms, it became apparent that the arrivals were small natural seismic events with sufficient energy to set up a pattern of water reverberations at the H2O site. In this report, I summarize the observation of these unexpected events. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 2

Magnitude

To determine the size of the small events the data were transformed to standard Wood- Anderson response to compute ML by deconvolving the seismic instrument response and then convolving the response of the torsion seismometer. Local magnitude is simply defined for this instrument as ML = log10(A) - A0, where A is the largest peak-to-peak amplitude in millimeters and A0is a normalizing factor that is a function of distance (Richter, 1958). Although A0also depends on the local wave propagation and attenuation conditions, the re-examination by Kanamori and Jennings (1983) of this distance term shows no systematic trend with respect to Richter’s values, and is within 0.1 of Richter’s value (-1.4) of A0 at short ranges.

The data for the small seismic events are shown in Figure 2. The measured ML is indicated, assuming the Southern California correction (Richter, 1958) for distances less than 10 km. The events are very small, less than magnitude zero. The magnitudes were calculated from both the Guralp and geophone channels, yielding the same results within 0.1 magnitude units.

The standard definition of microearthquakes in Japan, where the concept probably originated, is 1 ≤ M < 3 (H. Kanamori, personal communication, 2002), and events with M<1 are called ultra-microearthquakes. The Southern California Earthquake Center uses microearthquake as a term to describe earthquakes smaller than ML 2.0, and occasionally, slightly larger quakes, especially those not felt by people nearby (http://www.scecdc.scec.org/glossary.html). The National Earthquake Information Center also references a microearthquake as having a magnitude of 2 or less on the Richter scale (http://neic.usgs.gov/neis/general/handouts/glossary.html). The earthquakes observed at the H2O site in Figure 2 are about three orders of magnitude smaller than a microearthquake, and hence the term “nanoearthquake” is adopted for these tiny events with ML < 0.

In reviewing the events in the NOAA/PMEL database for microearthquakes detected or located by the Pacific hydrophone networks, no events were noted in the appropriate time windows nor were arrivals found on individual channels (R. Dziak, personal communication, 2001). Hence, H2O is recording local nanoseismicity that is not observed elsewhere.

Foreshock Sequence

Following the first nanoearthquake by about 48 minutes, an interesting foreshock sequence leads up to the second main event over about 5 minutes. This sequence of 4 events monotonically increasing in size has the appearance of a cycle of cracking and halted rupture. The time separation between the increasing foreshocks grows, with that between the last two foreshocks and the second main event nearly the same at about 113 seconds. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 3

Faulting Dimensions

The relationship between seismic moment, M0, and fault area, and the underlying linkage of moment magnitude and local magnitude, M = ML for earthquakes less than magnitude 6 (e.g., Thatcher and Hanks, 1973; Kanamori, 1978; Hanks and Kanamori, 1979), permit the estimation of the dimensions of these nanoearthquake ruptures. Heaton et al. (1986) have derived the equation M = 2 Log L + 2/3 Log Δσ + 2/3 Log CD – 10.7, which relates moment magnitude M, fault rupture length L (cm), stress drop Δσ 2 -6 (dyne/cm = 10 bars), and a constant CD ≈ 2 which is related to the shape of the rupture surface.

Both the stress drop and fault width of these nanoearthquakes are unknown, but a range of reasonable values indicates the possible fault dimensions. For simplicity, the fault length is assumed to approximate its width. The global average stress drop for intraplate earthquakes is about 60 bars (Kanamori and Anderson, 1975). The breaking strength of fresh rock is about 1 kilobar. For a stress drop Δσ in the range of 10 to 1,000 bars the magnitude of the largest nanoearthquake, M = -0.7, yields a characteristic fault dimension of 5.85 to 1.25 m. For the smallest nanoearthquakes with M ≈ -2, the range is then 1.3 to 0.3 m. However, the location for these events are found to be very shallow (< 1 km) and given the likelihood of fluid-saturated cracks/faults in this environment, it is plausible for lower stress drops, perhaps 0.1-1 bars (e.g., Hough and Kanamori, 2002). In this case, characteristic fault dimensions 5 to 2 times larger are possible. The source durations from the P and S arrivals for the larger events is less than 0.1 sec. Assuming a rupture velocity of about 1500 m/s in the shallow basalts, the corresponding dimension is less than 150 m, consistent with very low stress drop.

Location

The relative locations for the events can be estimated from polarization analysis of the P waves and differential times of successive acoustic reverberations in the ocean above the site. The polarization of P waves for the two main events and largest foreshock are all at the same azimuthal bearing of 220° ± 180°, as shown in Figure 3. The relative timing of first and second water reverberations, pW and pWW (e.g., Figure 4), depends solely on the distance between the H2O sensors on the seafloor to the point where the seismic energy enters into the water column. Since the velocity-depth profile of the ocean is very well known, this distance is well constrained. The pWW-pW differential times for all of the events that could be measured are each 6.4 seconds (Figure 5). This time corresponds to a distance of 3.7 km. Therefore, all of the events in the sequence appear to be tightly clustered at one of the two locations noted in Figure 1 relative to H2O.

The southwest nanoearthquake location falls on modest bathymetric feature trending NNW-SSE through the area. Single channel seismic (SCS) surveys conducted in August 1997 as part of the pre-drilling ODP survey fortuitously passed almost directly over the two possible locations for the nanoearthquakes (Stephen et al., 1997). The migrated SCS profile (Stephen et al., 2001) over the nanoearthquake location 3.7 km southwest of H2O shows a significant unconformity that may be a faulting feature associated with the Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 4 bathymetric trend observed in the Figure 1. Therefore, corroborative bathymetric and SCS data suggest the southwestern location for the nanoearthquakes.

Seafloor Velocity Structure

The nanoearthquakes have source dimensions estimated to be the order of 10 meters and less. Source durations appear to be less than 0.1 sec. Source spectra for the events show significant energy at 80 Hz Nyquist even though there is a 3-pole anti-aliasing filter at 20 Hz on the channels.

A consequence of the short source durations is that the uppermost oceanic crust is illuminated internally by a “seismic strobe light” that creates distinct seismic arrivals that can be used to investigate the seafloor structure at the H2O site. This is shown in Figure 6 where a variety of individual arrivals can be identified over a period of three seconds—a time frame wherein only a single oscillation might be seen for a typical teleseismic arrival. Using these relative arrival times, a detailed picture of the H2O seafloor structure can be derived from simple ray tracing (Figure 7 and Table 1). The layer nomenclature for the seafloor is used. Layer 1 is implicitly the sediment layer. Layer 2 is a basalt layer that is often subdivided into the slower Layer 2A comprised by pillow basalts, and a higher-velocity Layer 2B comprised by sheeted basalt dykes and hydrothermally altered pillow basalts.

The differential time pW-P, coupled with the pWW-pW derived distance, provides the first constraint. If the nanoearthquakes are located within a couple of 100 m below the sediments, the compressional velocity of the basalt must be about 3.0 km/s. If the nanoearthquakes are deeper, then the velocity must be slower since P and pW will share a larger similar path. A velocity of 3.0 km/s is comparable to values for Layer 2A basalts found near the mid-ocean ridge (e.g., Houtz and Ewing, 1976). Layer-2A-like pillow basalts were observed in cores from ODP drilling 1.5 km north-east of the H2O site beneath a thin (~30 m) higher-velocity consolidated basalt flow just below the 29 m sediment layer (Stephen et al., 2003). The S-P time then constrains the shear wave velocity to be about 1.5 km/s. Note that if Layer 2B velocities (VP ~ 5.2 km/s) are used, the S-P time suggests a distance of about 9 km, which is inconsistent with the pWW-pW reverberations. The velocities estimated here are very similar to results reported by Chapman and Hannay (1994) from a reflection survey north of the H2O site at 34°N from 130-144°W for oceanic crustal age from 42-63 m.y.

Following the S wave one observes a series of shear reverberations in the sediments, Sss+… (Figure 6). Assuming the same sediment thickness, 29 m, as found drilling at the ODP site, and a ray parameter appropriate for a shallow source 3.7 km distant, the average shear wave velocity of the sediments can be constrained at 97 m/s. The compressional wave speed of the sediments is assumed to be 1% greater than the water velocity at the seafloor.

The Ps arrival is over twice as large on the radial than the transverse component, and though it appears larger than S on the seismometer, it is a minor arrival on the Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 5 hydrophone. Without benefit of the hydroacoustic channel, Ps might easily be confused with S. Sp proceeds S as a large arrival on the vertical and hydrophone, and provides some constraint on the source depth. Assuming the seismic velocities in Table 1, the nanoearthquakes are best fit with a source depth of about 200 m below the sediments.

Prior to the P wave there is a small arrival that corresponds to a refracted head wave from a layer 500 m below the source with a Layer 2B basalt compressional velocity. Since a similar precursor is also seen for the other main event, it is unlikely to be a foreshock. By combining the apparent source depth, 200 m, with the apparent 500 m layer thickness below the source that accommodates the time of a refracted-P precursor, the Layer 2A is modeled at 700 m thick. This thickness is only an estimate, and depends upon the assumed velocity for Layer 2B.

It is unlikely that the interpretation presented here, though internally consistent, is unique. The shallow depths, within a few wavelengths of the seafloor, might result in interface waves that may be misidentified as body waves. Also, the uppermost crust and sediments likely possess gradients that would affect interpretation.

There is fairly substantial coda following Sp on the hydroacoustic channel, which is not large on the vertical component. The timing of this energy corresponds to the water sound speed. Butler and Lomnitz (2002) and Butler (2003) have discussed seismoacoustic mode coupling of T waves to seafloor sediments at the H2O site, and these arrivals might be related. An alternative explanation is that the coda is a Stonely wave propagating near the shear velocity of the Layer 2A uppermost crust.

Surprisingly a clear shear wave birefringence is evident for the radial/transverse orientation determined solely from the P wave polarization. The radial SV component is 0.085 seconds faster than the transverse SH, indicating 3% anisotropy in the basalt. This is consistent with vertically oriented cracks in the basalt (e.g., Crampin, 1993). Observing this polarization anisotropy is quite difficult using typical hydroacoustic air-gun or explosive sources, as only SV is transmitted across the seafloor interface from the water. Here, since the illumination is from within the media, the effect is evident. Barclay and Toomey (2003) have seen shear wave anisotropy of 8-30% in Layer 2A at the Mid- Atlantic Ridge using microearthquake sources.

Conclusion

The observation of nanoearthquakes at the Hawaii-2 Observatory is pure serendipity. Installed as part of a global network, H2O now reveals extremely-fine-scale details in the seismicity and structure of the interior of the Pacific plate. As a real-time seafloor observatory, small and seemingly insignificant arrivals in the H2O data may be examined on the spot, when interest and curiosity are greatest. The fidelity of the signals, particularly on the horizontal components, is so good that one can use the particle motion to identify phases at high frequencies, which is only possible with well coupled, buried sensors. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 6

The few nanoearthquakes examined here present a surprisingly detailed picture of the H2O site. The preliminary presented here could be further expanded with synthetic seismogram analysis and augmented with the seismic refraction and reflection data sets collected in conjunction with the H2O deployment and OPD site survey and drilling results. A cursory review of the H2O data for other days shows that these nanoearthquakes are not rare at the H2O site, and indeed dozens have now been seen.

Nanoearthquakes are only one example of the diversity of signals to be found at H2O. Seismologists are encouraged to explore the data from this seafloor GSN station.

Acknowledgements.

Supported by NSF Cooperative Agreement EAR-0004370. Ralph Stephen, Tom Heaton, Fred Duennebier, Andrew Barclay, Sue Hough, and Stuart Crampin provided a number of helpful insights.

References

Barclay, A. H. and D. R. Toomey (2003) Shear Wave Splitting and Crustal Anisotropy at the Mid-Atlantic Ridge, 35°N, submitted to J. Geophys. Res.

Bromirski, P.D. (2000) Vibrations from the “Perfect Storm”: Microseism source area implications, EOS Trans. AGU suppl. 81.

Butler, R. (2003) Seismoacoustic T waves and Oceanographic Spice, submitted to Geophys. Res. Lett.

Butler, R. and F. K. Duennebier (2000) Observations of high-frequency Po and So phases propagating in the Northeastern Pacific lithosphere at the Hawaii-2 Observatory, EOS Trans. AGU suppl. 81, F837.

Butler, R, and C. Lomnitz (2002) Coupled seismoacoustic modes on the seafloor, Geophys. Res. Lett. 29, no. 10, 57 1-4, doi:10.1029/2002GL014722.

Butler, R., A. D. Chave, F. K. Duennebier, D. R. Yoerger, R. Petitt, D. Harris, F. B. Wooding, A. D. Bowen, J. Bailey, J. Jolly, E. Hobart, J. A. Hildebrand, and A. H. Dodeman (2000) Hawaii-2 Observatory pioneers opportunities for remote instrumentation in ocean studies, Eos 81, 157-163.

Chapman, N. R. and D. E. Hannay (1994) Seismic velocities of the upper oceanic crust, Geophys. Res. Lett. 21, 2315-2318.

Crampin, S. (1993) A review of the effects of crack geometry on wave propagation through aligned cracks, Can. J. Explor. Geoph. 29, 3-17. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 7

Duennebier, F. K., D. W. Harris, J. Jolly, J. Babinec, D. Copson, and K. Stiffel (2002) The Hawaii-2 Observatory seismic system, IEEE J. Oceanic Engineering 27, 212-217.

Hanks, T. C. and H. Kanamori (1979) A moment-magnitude scale, J. Geophys. Res. 84, 2348-2350.

Harris, D. W. and F. K. Duennebier (2001) Powering Cabled Ocean Bottom Observatories, IEEE J. Oceanic Engineering 27, 202-211.

Hough, S.E. and Hiroo Kanamori (2002). Source properties of earthquakes near the Salton Sea triggered by the 10/16/1999 M7.1 Hector Mine earthquake, Bull. Seismol. Soc. Am. 92, 1281-1289.

Houtz, R., and J. Ewing (1976) Upper crustal structure as a function of plate age, J. Geophys. Res. 81, 2490-2498.

Jennings, P. C., and H. Kanamori. (1983). Effect of distance on local magnitudes found from strong-motion records, Bull. Seismol. Soc. Am. 73, 265-280.

Kanamori, H. (1978) Quantification of earthquakes, Nature 271, 411-414.

Kanamori, H. and D. L. Anderson (1975) Theoretical basis of some empirical relations in seismology, Bull. Seismol. Soc. Am. 65, 1073-1095.

Levitus, S., R. Burgett, and T. P. Boyer (1994) World Ocean Atlas 1994, Vol. 3: Salinity, NOAA Atlas NESDIS 3, U.S. Government Printing Office, 99pp, Washington D.C.

Levitus, S., T. P. and Boyer (1994) World Ocean Atlas 1994, Vol. 4: Temperature, NOAA Atlas NESDIS 4, U.S. Government Printing Office, 117pp, Washington D.C.

Peterson, J. R. (1993) Observations and modelling of background seismic noise. Open- File Report 93-322, US Geological Survey, Albuquerque, New Mexico, 94 pp.

Petit, R. A., D. W. Harris, B. Wooding, J. Bailey, J. Jolly, E. Hobart, A. D. Chave, F. Duennebier, R. Butler, A. Bowen, and D. Yoerger (2002) The Hawaii-2 Observatory, IEEE J. Oceanic Engineering 27, 245-253.

Richter, C. F. (1958) Elementary Seismology, W. H. Freeman and Company, San Francisco, pp 342-344.

Stephen, R. A., S. A. Swift, and R. J. Greaves (1997) Bathymetry and sediment thickness survey of the Hawaii-2 cable, WHOI-03-97, WHOI, Woods Hole, MA.

Stephen, R. A., J. H. Natland, R. Butler, K. Becker, A. D. Chave, F. K. Duennebier (2001) Drilling at the H2O Long Term Seafloor Observatory, Woods Hole Technical Memorandum WHOI-01-2001, WHOI, Woods Hole, MA. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 8

Stephen, R. A., Kasahara, J., Acton, G. D., et al., (2003) Proc. ODP, Init. Repts., 200, Ocean Drilling Program, Texas A&M University, College Station TX, in press.

Thatcher, W. and T. C. Hanks (1973) Source parameters of southern California earthquakes, J. Geophys. Res. 78, 8547-857.

Table

Table 1. H2O Seafloor Velocity Model

Layer Thickness, km VP, km/s VS, km/s Sediment 0.029 1.56 0.097 Layer 2A Basalt 0.7 3.0 1.51 (SV), 1.47 (SH) Layer 2B Basalt -- 5.2 --

Figure Captions

Figure 1. (left) The Hawaii2 Observatory at 27° 52.916 N, 141° 59.504 W is sited 1750 km east-northeast of Honolulu (Butler et al., 2000). The site lies between the Murray and Molokai Fracture Zones at 4979 m depth on predominantly terrigeneous, pelagic clay. The oceanic crust is of Eocene age (45-50 m.y.) and is normal but fast spreading with a spreading half rate of 7 cm/y. (above, center) The Hawaii-2 Cable terminates in western Oahu, and passes in between Kauai and Oahu before traversing northeast toward California. (right) The abyssal bathymetry is locally flat in a ~5,000 km2 region with abyssal hills rising less than 500 m high at about 30 km southwest. Two possible locations for the nanoearthquakes sequence analyzed in this study are indicated, as there is a 180° uncertainty in azimuth. Note trending bathymetric features that may correspond to earthquake locations. Hydrosweep bathymetry by R/V Thompson is from the 1998 H2O cruise.

Figure 2. Nanoearthquakes observed at H2O are plotted using a Wood-Anderson instrument response. The vertical components are shown (above) for the first main event, and (below) for the aftershock/foreshock sequence leading up to second main event about an hour later (window start times are noted). Multiple water reverberations follow the initial arrivals. ML measurements are indicated, computed using all three seismic components of motion. There is an interesting apparent periodicity of about 113 seconds between foreshocks leading up to the second main event.

Figure 3. The polarizations of the P waves are shown for the three largest events in the nanoearthquake sequence. All events are located at an azimuth of 220° ± 180° from H2O.

Figure 4. Schematic ray paths and associated names are shown for seismic arrivals identified for nanoearthquakes recorded H2O. Compressional waves are shown as dashed lines, shear waves are shown as solid lines. (a) Compression arrivals and (b) shear arrivals are illustrated. Not to scale. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003). 9

Figure 5. Hydrophone and seismic observations of the first main event show successive water reverberations (W) on the vertical, which are also seen on the noisier hydrophone channel and to a lesser extent on the radial seismic component. S-converted-to-P at the basalt-sediment interface generates spW reverberations. Horizontal components are oriented into radial and transverse polarizations, as determined from the polarization of the compressional waves. The principal constraint on event distance is the time between pWW and pW. The time between P and pW constrains the velocity of the basalt layer relative to the water velocity. Arrows indicate predicted arrival times from the velocity model in Figure 7. The other nanoearthquakes show essentially the same pattern of arrivals.

Figure 6. Hydrophone and seismic observations of the first main event are shown from the compressional arrivals through the shear wave reverberations in the sediments. The hydrophone channel is dominated by converted compressional wave energy, and effectively filters out shear waves. A small precursor is consistent with a refracted head wave from a Layer 2B basalt velocity. Strong conversion of P-to-S and S-to-P is evident from the sediment-basalt interface. Horizontal components are oriented into radial and transverse polarizations, as determined from the polarization of the compressional waves. Note that Ps is substantially larger on the radial component. Multiple reverberations in the 29 m sediment layer constrain the average sediment shear velocity to 97 m/s. Shear wave birefringence is evident with the radial SV faster than the transverse SH, indicating 3% anisotropy in the basalt. Arrows indicate predicted arrival times from velocity model in Figure 7. The other nanoearthquakes show essentially the same pattern of arrivals.

Figure 7. A simple velocity model with water, sediment, and two basalt layers predicts many aspects of the nanoearthquake records. Water depth is known (Butler et al., 2000) and sediment thickness determined by drilling (Stephen et al., 2003). The water velocity is derived from Levitus et al. (1994ab) and is 1.54 km/s at the seafloor decreasing to about 1.48 km/s at the axis of the SOFAR channel at 700 m depth. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003).

Hawaii–2 Cable H2O Bathymetry (150 m grid) 300 Km

28° 00′N

✳ ★ ✳ H2O SITE

27° 45′N H2O Murray F.Z. ★ Molokai F.Z.

27° 30′N

m 0 K 142° 15′W 142° 00′W 141° 45′W 150

5162.5 5075.0 4987.5 4900.0 4812.5 4725.0 4637.5 4550.0 4462.5 4375.0 4287.5

Figure 1. (left) The Hawaii2 Observatory at 27° 52.916 N, 141° 59.504 W is sited 1750 km east-northeast of Honolulu (Butler et al., 2000). The site lies between the Murray and Molokai Fracture Zones at 4979 m depth on predominantly terrigeneous, pelagic clay. The oceanic crust is of Eocene age (45-50 m.y.) and is normal but fast spreading with a spreading half rate of 7 cm/y. (above, center) The Hawaii-2 Cable terminates in western Oahu, and passes in between Kauai and Oahu before traversing northeast toward California. (right) The abyssal bathymetry is locally flat in a ~5,000 km2 region with abyssal hills rising less than 500 m high at about 30 km southwest. Two possible locations for the nanoearthquakes sequence analyzed in this study are indicated, as there is a 180° uncertainty in azimuth. Note trending bathymetric features that may correspond to earthquake locations. Hydrosweep bathymetry by R/V Thompson is from the 1998 H2O cruise. seconds between foreshocksleading uptothesecondmainevent. are indicated, computedusingallthree seismiccomponentsof motion.Thereisaninteresting apparentperiodicity ofabout1 about anhourlater(window starttimesarenoted).Multiplewaterreverberations followtheinitialarrivals. are shown Figure 2.Nanoearthquakes observedatH2Oareplottedusinga 3 3

-4 N-2 anometers x 10 -4 Na-2 nometers x 10 0 2 0 2 4 100 200 Butler, R.TheHawaii-2Observatory:ObservationofNanoearthquakes, Seismol.Res.Lett., W ood-Anderson Response (above) M forthefirstmainevent, and L ~-2 200 1 13 ±1sec M L =- (below) M L 300

0.7 =-1.1 fortheaftershock/foreshock sequenceleadinguptosecondmainevent Seconds 300 W ood-Anderson instrument response.Theverticalcomponents 1 13 ±1sec M 74(10),290-297,(2003). 400 L =-

0.9 500 M L measurements 600 400 13 Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003).

P W ave Polarizatio n 1st Main Earthquake 2nd Main Earthquake Largest Foreshock

ML = - 0.7 ML = - 0.9 ML = -1.1 h h h t t t r r r o o o N N N

220 ±180 220 ±180 220 ±180

East East East

Figure 3. The polarizations of the P waves are shown for the three largest events in the nanoearthquake sequence. All events are located at an azimuth of 220° ± 180° from H2O. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003).

(a) Compressional Waves

pW pWW

Water e l a c S

o t

▲ t

o Sediment N ✴ P Ps Basalt, 2A

Prefracted Basalt, 2B

(b) Shear Waves Water

e ▲ l

a Sp Sediment c S Sss o ✴ t S

t Basalt, 2A o N

Basalt, 2B

s pW - P = 5.8 sec l H2O Hydrophone a c s a P o n

a Bandpass 5-20 Hz

n pW SpW ) s r e

t pWW - pW = 6.4 sec H2O Vertical e m o n a n (

r e

t pW pWW pWWW e m o m

s H2O Radial 220º i e

n o s r e d n A - d o

o H2O Tangential 310º W

d r a d n a t S

Seconds Figure 5. Hydrophone and seismic observations of the first main event show successive water reverberations (W) on the vertical, which are also seen on the noisier hydrophone channel and to a lesser extent on the radial seismic component. S-converted-to-P at the basalt-sediment interface generates spW reverberations. Horizontal components are oriented into radial and transverse polarizations, as determined from the polarization of the compressional waves. The principal constraint on event distance is the time between pWW and pW. The time between P and pW constrains the velocity of the basalt layer relative to the water velocity. Arrows indicate predicted arrival times from the velocity model in Figure 7. The other nanoearthquakes show essentially the same pattern of arrivals. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003).

s P l Sp H2O Hydrophone a c s a P o n

a Bandpass 5-20 Hz

n Layer 2B Prefracted ) s r e

t P H2O Vertical e Sp m o n a n (

r e

t Ps S Sss e m o m

s SV H2O Radial 220º i e

n o s r e d

n SH A - d o

o SH H2O Tangential 310º W

d r a d n a t S

Seconds Figure 6. Hydrophone and seismic observations of the first main event are shown from the compressional arrivals through the shear wave reverberations in the sediments. The hydrophone channel is dominated by converted compressional wave energy, and effectively filters out shear waves. A small precursor is consistent with a refracted head wave from a Layer 2B basalt velocity. Strong conversion of P-to-S and S-to-P is evident from the sediment-basalt interface. Horizontal components are oriented into radial and transverse polarizations, as determined from the polarization of the compressional waves. Note that Ps is substantially larger on the radial component. Multiple reverberations in the 29 m sediment layer constrain the average sediment shear velocity to 97 m/s. Shear wave birefringence is evident with the radial SV faster than the transverse SH, indicating 3% anisotropy in the basalt. Arrows indicate predicted arrival times from velocity model in Figure 7. The other nanoearthquakes show essentially the same pattern of arrivals. Butler, R. The Hawaii-2 Observatory: Observation of Nanoearthquakes, Seismol. Res. Lett., 74(10), 290-297, (2003).

H2O Velocity-Depth Model 0

Vp 2

Water (4979 m)

3 h t p e D 4 *

Vs Sediment layer (29 m) 5 ✴ Earthquake Layer 2A ? (700 m)

6 Layer 2B Basalt

7 0 1 2 3 4 5 6 Velocity (km/s) Figure 7. A simple velocity model with water, sediment, and two basalt layers predicts many aspects of the nanoearthquake records. Water depth is known (Butler et al., 2000) and sediment thickness determined by drilling (Stephen et al., 2003). The water velocity is derived from Levitus et al. (1994ab) and is 1.54 km/s at the seafloor decreasing to about 1.48 km/s at the axis of the SOFAR channel at 700 m depth.