The NetLander very broad band seismometer

Planetary and Space Science 48 (2000) 1289–1302

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The NetLander very broad band seismometer P. Lognonnea; b;∗ , D. Giardinic , B. Banerdtd , J. Gagnepain-Beyneixa; b , A. Mocquete , T. Spohnf , J.F. Karczewskib; g , P. Schiblera , S. Cachoa , W.T. Piked , C. Cavoith , A. Desauteza , M. Favedea , T. Gabsia , L. Simoulina , N. Striebiga , M. Campilloi , A. Deschampj , J. Hindererk , J.J. Levequek , J.P. Montagnerb , L. Riverak , W. Benzl , D. Breuerf , P. Defraignem , V. Dehantm , A. Fujimuran , H. Mizutanin , J. Obersto a Dept.

des Etudes Spatiales, Institut de Physique du Globe de Paris, 4 Avenue de Neptune, 94100 Saint Maur des Fosses Cedex, France b Dept. de Sismologie, Institut de Physique du Globe de Paris, 4 Place Jussieu, 75252 Paris Cedex 05, France c Institute of Geophysics, ETH-H onggerberg, CH-9093 Zurich, Switzerland d Jet Propulsion Laboratory, California Institut of Technology, 4800 Oak Grove Dr. Pasadena, CA 91109, USA e Laboratoire de Plan etologie et Geodynamique, Faculte des Sciences et des Techniques, 2, Rue de la Houssiniere, BP92 208, 44322 Nantes Cedex, France f Institute of Planetology, Wilhelm-Klemm Str. 10, D-48149 M unster, Germany g Dept. des Observatoires, Institut de Physique du Globe de Paris, 4 Avenue de Neptune, 94100 Saint Maur des Fosses Cedex, France h Centre de Recherche G eophysique, Institut National des Sciences de l’Univers, Garchy, France i Laboratoire de G eophysique Interne et de Tectonophysique Observatoire de Grenoble, IRIGM, BP53X 38041 Grenoble Cedex, France j UMR G eosciences Azur, 250 rue A. Einstein, 06560 Valbonne, France k Institut de Physique du Globe de Strasbourg, 5 Rue Ren e Descartes, 67084 Strasbourg, France l Physikalisches Instit ut, Universitaet Bern, Sidlerstrasse 5, 3012 Bern - Switzerland m Royal Observatory of Belgium, 3, avenue Circulaire, B-1180 Bruxelles, Belgique, Belgium n Institute of Space and Astronautical Science, Yoshinodai 3-1-1, Sagamihara, 229-8510, Japan o DLR, Institute of Planetary Exploration, Rudower Chaussee 5, D-12489 Berlin, Germany Received 2 June 1999; received in revised form 5 January 2000; accepted 12 April 2000

Abstract The interior of Mars is today poorly known, in contrast to the Earth interior and, to a lesser extent, to the Moon interior, for which seismic data have been used for the determination of the interior structure. This is one of the strongest facts motivating the deployment on Mars of a network of very broad band seismometers, in the framework of the 2007 CNES-NASA joint mission. These seismometers will be carried by the Netlanders, a set of 4 landers developed by a European consortium, and are expected to land in mid-2008. Despite a low mass, the seismometers will have a sensitivity comparable to the present Very Broad Band Earth sensors, i.e. better than the past Apollo Lunar seismometers. They will record the full range of seismic and gravity signals, from the expected quakes induced by the thermoelastic cooling of the lithosphere, to the possible permanent excitation of the normal modes and tidal gravity perturbations. All these seismic signals will be able to constrain the structure of Mars’ mantle and its discontinuities, as well as the state and size of the c Martian core, shortly after for the centennial of the discovery of the Earth core by Oldham (Quart. J. Geol. Soc. 62(1906) 456 – 475). 2000 Elsevier Science Ltd. All rights reserved.

1. Introduction Our present understanding of the interior structure of Mars is mostly based on the interpretation of gravity and rotation data, or on the chemistry of the SNC meteoroids, and its comparison with the much better-known interior ∗ Corresponding author. Tel.: +33-145114251; fax: +33-145114257. E-mail address: [email protected] (P. Lognonne).

structure of the Earth (e.g. Schubert and Spohn, 1990; Longhi et al., 1992). Most of the Earth’s internal structure was determined by seismic studies: this started at the beginning of the century by Oldham with the early discovery of the velocity increase with depth and, in 1906, by the detection of a low P velocity zone in the centre of the Earth (Oldham, 1906), identi ed later as the Earth’s core. Almost a century later, the state-of-the-art consists of 3D models of the Earth’s mantle for both the

c 2000 Elsevier Science Ltd. All rights reserved. 0032-0633/00/$ - see front matter PII: S 0 0 3 2 - 0 6 3 3 ( 0 0 ) 0 0 1 1 0 - 0

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seismic velocities (e.g. Romanowicz, 1991), the anisotropy (e.g. Montagner and Tanimoto, 1991) and even anelasticity (e.g. Romanowicz, 1995). Seismic 1D imaging was also applied to the Moon (e.g. Nakamura et al., 1976, Nakamura, 1983), although the very deep structure remains unknown, due to the non-detection of large epicentral distance quakes by the Apollo seismic network. For a general review on planetary seismology, see Lognonne and Mosser (1993). Mars is, unfortunately, a seismically unexplored planet. Most of the past seismic experiments have indeed failed, either due to a launch failure, as for the Optimism seismometer (Lognonne et al., 1998) onboard the small surface stations of Mars 96 (Linkin et al., 1998), or a failure on Mars, as for the seismometer onboard one of the two Viking landers. The only attempt to address seismology on Mars was, therefore, made with the Viking seismic experiment onboard the remaining Viking 2 lander (Anderson et al., 1977). However, it did not result in convincing marsquake detection, basically due to very strong wind sensitivity, as well as to a very low resolution during non-windy conditions. Seismology on Mars thus remains to be studied. After almost one decade of continuous activities and proposals, such as the ESA Marsnet, NASA Mesur, and ESA-NASA InterMarsnet (Chicarro et al., 1991; Solomon et al., 1991; Chicarro et al., 1993, 1994; Lognonne et al., 1994; Banerdt et al., 1996) the rst network mission on Mars is expected to be launched in 2007 under the name NetLander. Originally proposed, in parallel, as a response to the 1997 CNES Small Mission Announcement of Opportunity (AO) and of the ESA Mars Express 1998 AO, by a joint French, Finnish and German team, it will be carried to Mars by the French Mars Sample Return orbiter, after a launch performed by Ariane 5 (Cazeau et al. 1999). See Harri et al. (1999) for more information about the NetLander mission. One of the main scienti c objectives of this 4 landers network mission will be the determination of the internal structure of the planet using a geophysical package. Such package will have a seismometer, a magnetometer and a Geodetic experiment, allowing a complementary approach, and therefore more constrains on the mineralogy and temperature of the mantle and possibly core of the planet. Geodesy and magnetic objectives are detailed by Dehant et al. (2000) and Mocquet and Menvielle (2000). For a review about Mars and its internal structure, see Spohn et al. (1998). Of these three dierent instruments, the seismometer has the best potential for depth resolution. The scienti c objectives of a seismic experiment on Mars are described in this paper. We also give a rst, preliminary, technical description of the experiment as such, proposed successfully by our team to the Netlander steering committee following the announcement of Opportunity issued in the late 1999 for the Netlander payload con rmation.

2. Scientiÿc objectives 2.1. Deep internal structure The Very Broad Band seismometer will perform both the seismic and tidal measurements. It was proposed onboard the NetLander by a large team of scientists, mostly involved only in Earth seismology and Earth tides. The seismic data analysis will determine the mean values of the shear and bulk elastic moduli and seismic attenuation as a function of depth, mainly from the transmitted phases. The re ected phases will mainly constrain the position of the interfaces between the mantle and core, the state of the core, the position and characteristics of mantle discontinuities and crustal thickness. These seismic data will consist of the recording of the natural quakes of the planet rst, and possibly, the recording of seismic signals generated by meteoritic impacts. However, parallel to these quake data, the experiment will also search for continuous seismic and gravity signals. These signals are associated with two continuous sources: • The rst one, in the frequency band of 0.1–10 mHz, is the atmosphere, more exactly the atmospheric turbulences. As recently discovered by Nawa et al. (1998), Suda et al. (1998), such turbulences on the Earth excite indeed continuously the fundamental branch of spheroidal Earth normal modes. As shown by Kobayashi and Nishida (1998), such an excitation process on Mars might be almost as strong as on the Earth. The inversion of the detected free frequencies of the excited normal modes might then, without any quake, provide information on the shear structure of the upper Martian mantle, and on the state and size of the core for the gravest modes (e.g. Van Hoolst et al., 2000). • The second one is the tides of the Sun and of the two small Martian satellites, Phobos and Deimos. The observation will focus on the tides produced by Phobos and the Sun, the latter producing a displacement of a few cm associated with a tidal forcing of the order of 10−7 m=s2 . These observations do not only give information on the direct attraction of these bodies on Mars, but also on the surface deformations (Love number h) and on the induced mass redistribution (Love number k). The contribution of the core to the sun tidal response could be of the order of half a mm. Here, the signal used will be the amplitude of the tidal gravimetric factor, which is a linear contribution of the Love numbers h and k. It corresponds to the Mars transfer function to the external tidal forcing. Details on this transfer function are given in Dehant et al. (2000). Together with geodesy and magnetic sounding, a multiparameter characterisation of the Martian mantle will therefore be performed. This multi-parameter approach is the key issue for the understanding of the planet’s interior in terms of mineralogy and temperature pro le with depth.

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Fig. 1. Predicted variations of S wave velocities as a function of the iron content of the Martian mantle (Mocquet et al., 1996).

Several papers illustrated the sensitivity of the seismic and gravimetric measurements with respect to the internal structure. See Okal and Anderson (1978), Gudkova and Zharkov (1996), for the normal mode frequencies, or Lognonne and Mosser (1993), Dehant et al. (1999) for the tidal gravimetric factor. Let us illustrate here the sensitivity to mineralogy for the more classical seismic measurement of the mantle discontinuities, which can be determined from the arrival time analysis of the phases re ected at the discontinuities. If the Martian mantle composition is close to the Earth’s sharp density and seismic velocity discontinuities (∼0.5 km=s) are expected to be present between 1100 and 1500 km depth (Okal and Anderson, 1978). These discontinuities are mainly associated with the phase transitions of olivine in the 12–16 GPa pressure ranges. They should induce triplication of P and S wave travel times versus epicentral distance. In contrast, an enrichment in iron with respect to the Earth’s mantle (e.g. Ringwood, 1979; Dreibus and Wanke, 1985; Wanke and Dreibus, 1988) induces the coexistence of  and phases of olivine over a 2 GPa wide domain of pressures. This is corresponding to depth between 1000 and 1200 km (Bertka and Fei, 1996, 1997; Mocquet et al., 1996). Consequently an increase in the iron-content smoothes out the discontinuities over a thickness of one to two hundred kilometres (Fig. 1). This smoothing process should induce a focusing of seismic rays in this depth range, and reduce the importance of body wave triplication. In parallel, the attenuation will be measured both in the seismic band, from the amplitude of seismic waves, and possibly from the phase-lag of the tides. This measure will provide an improvement of the attenuation at the period of the Phobos tide, which is estimated from the Phobos secular acceleration to be Q ≈ 100 ± 50 at a period of about 320 min. (Smith and Born, 1976). Such a low value can be related to mineralogy, for example, to the existence of an asthenosphere associated with a hydrated mantle at a depth close to 300 km (Toksoz et al., 1978; Lambeck, 1979). It can also be related to temperature which is controlling

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either the tidal acceleration through temperature-dependent viscosity (Sohl and Spohn, 1997) or the seismic attenuation through the distribution of thermally activated relaxation mechanisms (e.g. Anderson and Given, 1982). To conclude, the main objective of the geophysical package will therefore be the determination by geophysical methods of these parameters, and subsequently of their simultaneous inversion. When done with mineralogical constrains obtained from state equations and laboratory measurements, such an inversion will allow the recovery of the thermal and mineralogical state of the interior of Mars. The seismic information that will be retrieved by such a network will however suer from the number of stations, especially if we have in mind the present state of the Earth networks. Only a 1D seismic model will then be determined by the experiment, from all the dierent seismic information available (body wave arrival times, surface waves velocities measurement, free oscillation frequencies). The a priori limitation of these methods are already described by Solomon et al. (1991), Lognonne and Mosser (1993) and Lognonne et al. (1996), including a priori synthetic seismograms. For body waves practically the main limitation will be related to error in the location of marsquake hypocentres (horizontal co-ordinates and depth), very likely strongly associated to lateral heterogeneities, including crustal thickness. Fig. 4a–d provide an illustration with an example taken for the Apollo seismic network, for which, however, the lack of broad band capacity of the sensors has resulted in large errors in the S phase arrival time picking. Such localisation is a prerequisite for all investigations of the planetary interiors and must involve the identi cation of as many seismic phases as possible (initially P, S, PKP and secondary phases) and the joint=iterative re ning of hypocentres and travel-time curves. A set of 50 quakes, all with records of P and S at 3 stations, will provide, however about 300 secondary seismic data. These will constrain the event’s position and origin time (4 × 50 = 200 unknowns), and the velocity models with transmission data. Probably, as shown by Fig. 3b, 25% of these quakes might be associated also with the PKP measurement, if all stations are operational. If augmented by the arrival time of re ected phase, such as PcP, P1100 P, etc, this set might provide about 150 degrees of freedom from the travel time data of direct and re ected phases, directly useful for the model determination. Such a data set might be sucient for the determination of a preliminary symmetrical model of the planet. 2.2. Tectonics and marsquakes Another key goal of the seismometers will be the detection, characterisation, and quanti cation of marsquakes, in terms of space, time, and size distribution of the events and of their source parameters (focal mechanism, source

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complexity and stress drop). It will allow the understanding of the presence and type of tectonic deformation and strain release on Mars’ surface, with respect to models of thermo-elastic cooling and lithospheric relaxation. The determination of marsquake source properties (hypocenter, focal mechanism and stress drop) will integrate other data from Mars’ missions. The images from orbiter’s high resolution camera and altimetry measurements will help in the identi cation of active structures and faults, the determination of the state of stress and strength envelope in the crust and lithosphere and the relation between vertical and horizontal tectonics in active areas of Mars. 2.3. Meteoroids Meteoroids constitute the “building blocks” of planets and are therefore of great importance for our understanding of Solar System origin and evolution. The traditional method of studying meteoroids has been the sporadic observation of their terrestrial atmospheric entries from the ground. More recent techniques include Spacewatch telescope or satellite infrared global observations that routinely detect meteoroids approaching near-Earth space and meteoroids encountering our planet. In addition, the 1740 meteoroid impacts detected on long-period records by the Apollo Seismic Station Network (Nakamura et al., 1982) provided a unique data set to characterise the present-day impact ux on the lunar surface. Studies of the meteoroid ux beyond Earth’s orbit are of great scienti c merit. They yield information on the overall meteoroid population in the Solar System and in particular the calibration information for estimating ages of the surfaces of Mars (Neukum and Hiller, 1981) and asteroids (Chapman et al., 1996) by crater statistics. Scienti c issues to be resolved include, whether the meteoroid population observed on Mars is the same as the one that bombards Earth and Moon (Neukum and Ivanov, 1994). Also, the absolute magnitude of this ux is uncertain. Shoemaker (1977) and Davis (1993) suggest that the entry ux of meteoroids entering the atmosphere of Mars is 2.6 times that on Earth due to proximity of the asteroids belt. Based on these data, Davis (1993) estimates that the number of large impacts per year detected by a seismic station on Mars (assuming a sensitivity of the Apollo seismometers) is similar to that on the Moon, i.e. approximately 100 events per year. However, the analysis of the rst high-resolution images returned from Mars Global Surveyor 96 indicate a large abundance of craters, 15 m in diameter and possibly smaller (Hartmann, 1998). Previously, minimum crater diameters were thought to be close to 50 m, with the atmospheric shield preventing the types of impacts that would form smaller craters. This new nding suggests that atmospheric ltering has little eect on the surface impact rates and that seismic detection rates should be larger than estimated by Davis (1993).

In conclusion, the goal of seismic meteoroid impact observations is to characterise the impact ux near Mars in terms of cumulative energy distribution of impactors, absolute meteoroid mass ux, variations of impact rate during the Martian day and clustering of meteoroids. The rst task will be to distinguish between quakes and impacts. Very likely, impacts will be dierent from quakes with respect to the P=S ratio, the P coda as well as their frequency dependence. With such identi cation, it will then be possible to interpret basic observations such as position and time of impact in terms of approximate orbital distribution, possible types, and possible meteoroid sources. Results will be put into the context of the various existing data sets of Earth and space-based observations of meteoroids. Implications of the cratering rate on Mars and asteroids will be evaluated. The data analysis will follow the methods developed for the analysis of meteoroid impacts detected by the Apollo Seismic Stations (Oberst and Nakamura, 1989, 1991). It will include discussions of atmospheric shielding eects and the question of seismic eciency (fraction of preimpact kinetic energy of the meteoroid that is converted to seismic energy). 3. Instrument description and performance 3.1. Expected surface and body waves signals Marsquakes may be generated through the release of thermal stresses. This activity is about 100 times greater than the shallow moonquake activity detected by the Apollo seismometers with good signal-to-noise ratio. The amplitude of the Mars seismic signal is still expected to be about 4 orders of magnitude lower than on the Earth. This was detailed by Golombek et al. (1992) from surface fault observation and by Phillips (1991) from a theoretical estimate of the thermo-elastic cooling of the lithosphere. It might provide about 14 quakes of seismic moment 1015 N m per year, with an increase=decrease of the frequency by 5 for a decrease=increase of the seismic moment by 10. The teleseismic signal will be observed as body and surface waves. For the body waves, the main limitation is related to attenuation and scattering. The high attenuation of Mars is con rmed by the high secular acceleration of Phobos, which needs a quality factor of about 100 at the tidal period of Phobos and was extrapolated in the seismic range to about 350 by Lognonne and Mosser (1993). Very likely, the Mars mantle, due to its low pressure, is in a thermodynamic state more comparable to the highly attenuating Earth’s upper mantle rather than to the lower mantle, which features lower attenuation. The eects on the amplitude of body waves are plotted on Figs. 2a–b, after Mocquet (1998). In contrast to the Earth, where the propagation of body waves is ecient at high epicentral distance, due to the low attenuation of the lower mantle, we observe here a severe reduction

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Fig. 2. (a) Fourier transformed amplitude of the P and S body wave packet. The amplitude is plotted with respect to frequency for dierent epicentral distances and for a source moment M0 =1015 Nm. (b) Maximum peak-to-peak amplitude in the frequency band as functions of epicentral distance for the same seismic moment as 1a.

in the amplitude of S waves at epicentral distances greater ◦ than 50 . The amplitudes are computed for two frequency bands (0.1–1 Hz, 0.5 –2.5 Hz). These amplitudes are for a quake of seismic moment of 1015 Nm, with a source at the surface, and an isotropic source. The crustal transmission as well as geometrical spreading and attenuation, is taken into account, for a Mars model compatible with the present a priori knowledge. The amplitudes at frequencies higher than 0.5 Hz (0.5 –2.5 Hz) show a strong decrease of amplitude with epicentral distance for body waves. P waves in the frequency band 0.1–1 Hz are in contrast not strongly sensitive. However, P amplitude in the band 0.5 –2.5 Hz will be especially sensitive to scattering. This was demonstrated on the Moon and is related to the multiple fractures of the crust, related to meteoroid impacts as well as tectonics structure (e.g. Dainty and Toksoz, 1977, Nakamura, 1977). On the Earth, scattering is also very strong in volcanoes,

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and therefore we may expect a very strong scattering in the Tharsis area too. For body waves, the scattering reduces the peak-to-peak amplitude by producing conversions of the wave (mainly P to SV, P to SH) and spreads this energy in time. The eect can be as strong as a reduction by 10 of the P wave integrated energy, and stronger for the maximum peak-to-peak amplitude due to the length of the coda. It will probably aect less the P waves at longer period (0.1– 1 Hz), and much less the S waves (Aki, 1992). We will take therefore the estimate in the band (0.1–1 Hz) as the most robust with respect to attenuation (mainly for S) and diraction (mainly for P). Let us now consider the noise level. Noise level on the Martian surface is also expected to be much lower than on Earth, as Mars has no ocean, responsible for the high noise level in the 30-1s range on the Earth. Typical Martian noise should be about 1000 times smaller at intermediate period (30-1s), a period range dominated on Earth by the noise associated with the oceanic waves. We can therefore almost expect noise levels less than 10−8 m=s2 =Hz1=2 , and very likely close to or less than 10−9 m=s2 =Hz1=2 . The lowest value in the band 0.1–1 Hz is corresponding to less than 3 × 10−9 m=s2 , respectively 5 × 10−9 m=s2 peak-to-peak noise level. A quake with a seismic moment of 1015 Nm ◦ will therefore generate at epicentral distances of 50 , signals with a S=N ratio of about 30 for the S wave and 200 for P. Such ratio will be excellent for the remote detection of P and S. Figs. 3a–b detail these perspectives for a possible Network con guration, assuming total detection of signals with accelerations greater than 10−8 m=s2 peak-to-peak. They show that a rate of 60% can be achieved for the detection of quakes with seismic moment greater than 1014 Nm, i.e., corresponding to Earth magnitude greater than 3.2. This might provide about 100 detected quakes out of the 140 quakes that are expected during the two years of the NetLander mission lifetime to occur with at least this seismic moment. The possibilities in the localisation of the quakes are illustrated by the Figs. 4a–d, based on the example of the Apollo Lunar seismic Network (Gagnepain-Beyneix and Lognonne, in preparation). Note that the 4 stations network con guration, in contrary to the Moon seismic network, will allow a much better characterisation of the core, with the antipodal stations near Hellas Planitia. Fig. 3 shows also the redundancy eciency, in the case of failure of one of the stations: three of the zones of highest sensitivity in Fig. 3a can indeed be superimposed to the three areas appearing in Fig. 3b, which shows the detection area if one of the stations failed. All these possibilities are however closely related to the performance of the instruments and their installation. Concerning the installation, the rst priority will be to have a direct contact with the ground together with a mechanical decoupling of the sensor from the lander, as opposed to the Viking design. In this con guration, it will still be necessary to shield the seismometer from the direct

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Fig. 3. (a) Detection zone of P and S waves for one of the studied con gurations (MEM, LYS, TTS and HEE are respectively Menomnia Fosae, Lycus Sulci, Tempe Terra South and Hellas East). The white zone on to the east of TTS corresponds to the joint shadow zone of MEM and HEE. Note that the detection eciency is very high in the Tharsis area, where small quakes of seismic moment of 1014 Nm might be detected. (b) Detection zone for a detection of a PKP wave in at least one station and of P& S waves in the three other stations. The success rate is 16%, leading to about 22 quakes detected during the two years of operation.

eect of the wind on its structure. Tests were performed with a mock-up of a seismometer platform covered by a simple windshield. They have shown that vertical noise with spectral amplitudes of a few 10−8 m=s2 =Hz1=2 near 1 s and few 10−9 m=s2 =Hz1=2 close to 20 s can be realistically achieved with seismometer deployed at the Earth surface (Lognonne et al., 1996). On horizontal components, the noise has a spectral density of a few 10−8 m=s2 =Hz1=2 . In both cases, a further improvement of the signal-to-noise ratio can probably be achieved by recording the variations of the meteorological parameters (temperature, pressure) in the seismic frequency band and by performing a real-time decorrelation, as demonstrated by Beauduin et al. (1996). A level of 10−9 m=s2 =Hz1=2 resulting from the conjunction of a windshield, a good thermal protection, and the a priori low ground noise is a reasonable assumption. It is therefore necessary to have an instrument with a smaller noise level, in the seismic band, as close as possible to 10−10 m=s2 =Hz1=2 .

The amplitudes of the possible permanent excitation of normal modes was estimated by Kobayashi and Nishida (1998), and is expected to be of the order of or below one ngal. With a noise spectral density of 10−9 m=s2 =Hz1=2 and a spacing of 0.2 mHz between two modes, used as frequency band for the search of a single peak, the detection level will be equivalent to rms level of 2 ngals and therefore comparable to the a priori expected signal. A very low residual instrument noise after the atmospheric noise decorrelation, as well as stacks, will therefore be necessary to detect this signal. It will be very likely the most challenging part of the experiment. Finally, at longer periods, the main signal will be the tides produced by either the Sun or Phobos. The amplitudes are given by Table 1, and will be less critical than the above normal modes amplitude of the possibly permanently excited normal modes. The long-period channel of the seismometers, as done on the Earth with equivalent STS-1 or STS-2 sensors

Fig. 4. a–d. Example of localisation performed on the Moon by the Apollo seismic array. d shows the data, for both the long period (LP) and short period (SP) outputs. P and S waves arrival times are shown. Note however that the amplitudes are small, and diraction very strong, leading to very large errors in the S arrival time. b shows the a posteriori probability for the localisation in depth normalised to its maximum value, and (a) – (c) the same with respect to longitude and latitude. The 1-sigma ellipse is shown on (c). Note that this is a pessimistic con guration, the quake being outside the area and the depth being not constrained by secondary pP or sS phases. The localisation of the Apollo stations is given on a.

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Table 1 Amplitude and frequencies of the part of tidal signals related to the solid planet Tidal modes (deg., order)

Amplitude gal (10−8 m=s2 )

Frequency (mHz)

Sun: 2,2 Sun: 2,1 Phobos: 2,2 Phobos: 3,3 Phobos: 4,4

4.6 4.3 0.45 0.20 0.09

0.02 0.01 0.05 0.07 0.11

(Pillet et al., 1994), will detect these signals. A careful analysis shows that even small tides of several hours period can be correctly determined (Freybourger et al., 1997). In particular, when recording in quiet sites, the diurnal and semi-diurnal principal luni-solar tides on the earth are easily detected with amplitudes close to those obtained from good gravimeters despite their low frequency with respect to the classical seismic band (below one hour) for which the seismometers are designed. Our capacity in detecting the tidal signals on Mars due to the Sun and Phobos (Deimos leads to a negligible signal) will of course depend on the noise level acting in the relevant band. It will be partly of instrumental origin (sensor and electronics) and also of physical origin (in particular temperature, atmospheric pressure, and winds). With broadband records, a frequency analysis will separate the lines at the frequency of Phobos from those related to any harmonics of the day. Additionally, the pressure and temperature data will be available for the seismic-tidal experiment and will help in reducing noise contamination (Crossley et al., 1995; Baudouin et al., 1996). The design of the seismometer package will therefore be optimised for the achievement of the highest sensitivities within the mass, power, and volume allocations, compatible with the above amplitudes. It will use two dierent types of sensors: • 2 VBB Long period oblique sensors: These sensors will have a very high mechanical ampli cation, providing a vertical Very Broad Band (VBB) output with very low noise, and a single horizontal output. The oblique design, for the mass and volume allocation, allows the best mechanical ampli cation for the recombined vertical as compared to a pure vertical axis. These channels will also provide the vertical gravity output. This very low instrument noise on the vertical seismic and tidal output will be perfectly suitable for the low noise levels, which are expected on the vertical components on Mars. • 2 Short period broad band sensors (BRB): A single horizontal axis sensor will be integrated inside the VBB sphere, and therefore levelled. The instrument noise level, relatively high with respect to the VBB, will probably be comparable to the horizontal noise on Mars. A full three-axis sensor will be xed on a spike in direct contact with the ground. It will not be levelled, but will proba-

bly have a better coupling for high-frequency measurements than the levelled axis, due to the direct coupling with ground. It will provide the vertical and both horizontal components, with an inclination provided by their DC output. The 2 LPs and the horizontal SP are in a levelled sphere under vacuum as shown on Fig. 6. 3.2. Long-period output and tidal output The Very Broad Band (VBB) seismometer was developed through a Research and Technology Program of CNES, in preparation of the InterMarsnet program, using some heritage of the OPTIMISM seismometer, which was on board the Mars 96 Small Surface Stations (Lognonne et al., 1998). This instrument is composed of two tilted axes (see Fig. 5a) providing, after recombination, data of one horizontal axis and one vertical axis. The limitation to two axes is related to the size of the lander, much smaller than the InterMarsnet lander. However the use of a micro-seismometer (see Fig. 5b) for the second “missing” horizontal component allows the recovery of all three axes (see Fig. 6a–b), even though the instrument noise on this second component is expected to be about two orders of magnitude greater (Fig. 7). The VBB instrument has been designed to study the very low accelerations associated with the small magnitude of Marsquakes, in a thermal environment featuring very strong temperature variations (Cacho et al., 1999). Therefore, this sensor must be very sensitive, must have a very low self noise and must have a design relatively insensitive to environmental parameters. The seismic mass therefore is in an evacuated sphere and less sensitive to pressure changes. Its has an internal temperature compensation and high thermal dewar-like protection, limiting the seismic mass displacement to less than 5 m over the daily cycle. And nally, the design is compatible with an operation both with Earth and Mars’s gravity. The sensor was therefore rst designed to reduce its Brownian noise, whose power spectrum is roughly inversely proportional to mQT , where m is the inertial mass, Q the quality factor of the sensor and T the period (e.g. Aki and Richards, 1980). For Earth seismometers, this is generally done by increasing the proof mass, up to typically 0.5 – 0.8 kg, and by having a free period from several seconds up to almost 10 s, as with the various versions of Streikeinsen or Guralp seismometers. The VBB inertial mass, due to both mass and shock constraints, is lighter than 0.1 kg. This leads to a limitation on the free period to less than 2 s, with some freedom in the adjustment of the free period in the range of 0.5 –2 s. The only alternative was a high Q. The sensor head therefore has a high-quality mechanical pendulum with a Q-factor close to 400, for a frequency close to 1.5 Hz (Fig. 8). As shown by the Fig. 8, the Brownian noise through the Q factor will be related to the residual pressure inside the

P. Lognonne et al. / Planetary and Space Science 48 (2000) 1289–1302

Fig. 5. (a) Left: Breadboard of one VBB axis. All the structure of the seismometer is realised in titanium, in order to reduce the thermo-elastic deformations. (b) Right: Silicium wafers of the 3-axis BRB seismometer proposed by the Jet Propulsion Laboratory. The assembly consists of the stack of the three wafers. (Note: the French coin, on the left, and the US coins, on the right have about the same size, 2:4 cm in diameter).

sensor enclosure. A primary vacuum will therefore be required and possibly maintained by getters. The second eort was to have a very ne thermal compensation, in order to reduce the displacement of the proof mass over the daily 60 K variations in external temperature at the surface of Mars. The elastic properties of the spring used have been adjusted to compensate for aggregate thermal drifts. This mechanism has been designed using nite element software and measured thermal sensitivities as low as 80 nm=K were reached. Even without feedback, the maximum displacement excursion of the proof mass will therefore be less than 5000 nm, and can be further reduced by almost a factor of two with ecient thermal protection. By

1297

Fig. 6. Top and side view of the evacuated sphere with the 2 VBBs and one SP/JPL axis. The diameter of the sphere is 130 mm and will be under vacuum.

using a 24-bits A=D converter, this will allow the use of a displacement transducer operating with a Least Signi cant Bit (LSB) below 1 pm. Two independent transducers will do the displacement measurement: • An Oscillating Cavity Sensor (OCS), producing a digital tidal output (f ¡ 10−3 Hz), with a dynamic range of 32 bits; • A Dierential Capacitive Sensor (DCS), providing an analogue long-period seismic output (10−3 –10 Hz), with a dynamic range of 140 dB (with a spectral noise close to 1 pm=Hz1=2 at 1 Hz).

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Fig. 7. Performances of VBB and JPL seismometers.

Fig. 9. Earthquake recorded by an STS2 and by a mock-up of the VBB seismometer, on 30 May 1998. The epicentre was in Afghanistan. The signal was recorded in a seismic vault, in Saint Maur, France. The quake had a magnitude of 6.9, with a focal depth of 33 km. The STS2 outputs were used to compute the oblique output compared to the mock-up signal. Note however that the breadboard was operating without feedback and at a slightly dierent place than STS2, which may explain the small dierences, between the two signals.

3.3. Integrated package and environmental sensors

Fig. 8. Brownian noise of 4 dierent mock-up units featuring dierent frequencies in the range 0.5 –2 s. Below 10 bar of residual pressure, and in all cases, the noise is expected to be less than 2 × 10−10 m=s2 =Hz1=2 .

Feedback, necessary for an instrument with such sensitivity and high Q will be performed with both sensors in order to reduce the mechanical recentering and to provide the ideally continuous record necessary for tidal analysis. The OCS contributes to the digital part, to prevent long-period drifts, and the DCS contributes to the analogue part, permitting an adjustment of the gain, of the stability of the instrument and of the reduction of the resonance peak. The performances expected by the VBB sensor are shown on the Fig. 7, for the mock-up presently being tested as well as for the future Flight Models. A comparison between the record of a STS2 and of a mock-up is given in Fig. 9. The performance will depend on the nal optimisation of the mechanical (free period, Q) and electrical parameters (feedback and displacement transducer gain) of the sensors, in order to minimise the sum of all the noise sources.

The VBB instrument hardware will consist of a primary subsystem, with a height of 15 cm and a diameter of 18 cm, to be mounted on the platform with a direct access to the ground plus electronics and a data-logger. The rst part, the sensor, levelling system, installation system, locking mechanism and environmental sensors, as well as the overall instrument integration will be under the responsibility of the Institut de Physique du Globe de Paris, France (PI: P. Lognonne), with support of CNES. The second part of the sensor will be the proximity electronics, a double-face electronic card for the CDMS I=F, A=D and feedback electronics, and the mass memory for the data logger. This part of the experiment will be under the responsibility of Institute of Geophysics, ETHZ, Switzerland (co-PI: D. Giardini), with support by ESA/PRODEX. The last part of the seismometer package will be the micro-seismometers, either in the VBB sphere or on the spike. This part will be under the responsibility of the Jet Propulsion Laboratory, USA (co-PI, B. Banerdt), with a possible support of NASA. The block diagram is shown in Fig. 10. The expected mass and power breakdowns are given in Tables 2 and 3. Pressure and temperature will be recorded with high resolution (b and K) in order to remove their in uence on the seismic signal. This will be done by the CDMS for the seismic signals with sampling rates higher than 1 sps, and on the Earth for the tidal outputs. The total mass of the complete instrument (with installing and levelling devices) will be about 1935 g without margins and is given in Table 2. The average power requirement will be about 600 mW (900 mW peak during the installation) and is given in Table 3.

P. Lognonne et al. / Planetary and Space Science 48 (2000) 1289–1302

1299

Fig. 10. Seismometer block diagram.

Table 2 (left): Mass breakdown of the NetLander seismometer

Table 3 (right): Power allocation of the seismometer

Sub-system

Volume

Mass (g)

Sensor heads, leveling system, vacuum enclosure and thermal protection Installing device Locking device Levelling device Micro-short period sensors House-keeping sensors (pressure and temperature) Electronics (1 card, double

all in a cylinder of h = 15 cm x ∅ = 18 cm

700

(Excluding the locking ngers) 12 × 15 × 2

mounted) Cables Total without margins

300 240 310 105 30 200 50 1935

3.4. Control, acquisition and data strategy Except during installation when a special acquisition sub-routine will be used, the standard process involved for

Sub-system

Mean power (mW)

Peak power (mW-duration)

CCD CCO P (12 MHz) Feed-back A=D Micro-sensors (JPL)

2 × 5 = 10 2 × 20 = 40 150 2 × 20 = 40 100 150

10 40 180 40 150 150

House-keeping Heater Installation

10 100

10 300 mW–8 h 1 Wh=d during 2 d 250 mW max

Total

600

900 max during installation 800 max peak

data acquisition will be based on a high-rate sampling of all parameters, and their transfer to the large-capacity on-board

1300

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memory. The short-period sensors outputs will be sampled at the highest rate, possibly 200 sps (16-bit word with maximum sampling rate of 500 sps) for all 3 axes. For the VBB seismometers (2 axes), sampling at a rate of 100 sps (16-bit word) is foreseen. An internal microprocessor will take care of all safety programs like the automatic digital feedback to ensure correct pendulum position and the digitisation of all auxiliary data. Such data that correspond to an additional rate of 50 sps (16-bit word) are temperature with K resolution, pressure with bar resolution, and tiltmeters with resolution of minutes of arc. The total digital stream transmitted to the Control and Data Management System (CDMS) for the seismometer will therefore be about 13; 600 bits=s. The data will then be digitally ltered to reduce the nal storage rate to 100 sps for the SP vertical sensor, 20 sps for all other seismic outputs (VBB and horizontal SP) and 5 sps for the housekeeping data. After digital ltering all 20 sps data will have 24-bit words and the data ow will correspond to 2400 bits=s with header, housekeeping data and time, and with a compression ratio of at least 2 by using a delta code. These data will be stored in a 4 Gbits turning buer, which will therefore complete a full turn in little less than 20 d. It will be impossible to transmit all these data. Instead, a low-frequency digitally ltered data set will be transmitted to the Earth almost continuously. It will consist of: • Very Long Period (VLP) data (e.g. with one sample every 10 s) from the vertical and horizontal components (for tide and a general quick look of instrument behaviour), for a total of about 100 kbit=d. • Long Period (LP) data at 1sps for the vertical output. This will be used to identify the fraction of time when full VBB data is to be transmitted (at a rate of 20 sps) • Spectral amplitudes of the Short Period (SP) vertical output at a rate of 1 multi-spectral estimate every second. We expect a daily transmission of these data in a volume of 2:5 Mbits=d. The seismometer team in at least two geographical locations will perform the quick-look on the Earth of these data, in order to maximise the turn-around time during regular shift hours (Paris and Pasadena, UT+1 and UT−8). From these data, a set of time will be identi ed, and a table of parameters will be up-linked to each lander in order to ag and to save the interesting data in the general memory of the CDMS (e.g. when quakes are tentatively identi ed). The corresponding VBB and SP data will then be progressively sent as EVENT data at a rate of about 5 Mbits=d. After a learning period, an automatic event detection algorithm will be implemented. The algorithm might be based on a STA=LTA or other classical event trigger or on the evolution with time of the compressed power spectral density computed on board. More sophisticated programs can be transmitted from the Earth if necessary.

3.5. Deployment and decoupling system The two axis VBB and a one axis SP micro-seismometer will be located in an evacuated sphere of 130 mm diameter. This sphere can be levelled and is supported by 3 small passive legs. This aggregate will be locked on the lander by 4 equatorial nger devices. All mechanisms will be designed to support a landing shock of 200 g during 20 ms. This shock level is similar to the Mars 96 speci cation, for which the Optimism hardware (Lognonne et al., 1998) was quali ed. The seismometer deployment will consist in a controlled drop of the instrument on the ground. In a critical case, the seismometer will stay coupled to the lander. When the three small legs contact the ground, the sphere containing the sensor will then be levelled automatically (with an accuracy of a fraction of one degree), and locked. Each VBB axis is equipped with a device allowing a complete and precise levelling of the pendulum. The system used for the seismometer deployment, even though not comparable with a full robotic or human installation, is therefore expected to provide a good installation. The in uence of the lander on the seismometer will be minimised due to the decoupling and coupling with ground at high frequencies (f ¿ 10 Hz) will be enhanced. Finally, the lander will act as a thermal and windshield for the instrument. This system will be designed to work even if the lander is tilted (up to ◦ about 30 ). 4. Conclusion The NetLander mission is expected to provide the rst network on Mars. The expected performances of the seismometer planned on the payload will allow the monitoring of the Martian seismic activity, especially in the Tharsis area. An attempt to monitor the continuous excitation of normal modes, as well as the gravity tides of the Sun and Phobos will also be performed. The NetLander mission, presently in study, is expected to start the realisation phase after the con rmation of the payload in April 2000. Delivery of the instrument is expected 1 to 2 yr before launch. The seismic network will operate from 2008 to early 2010, i.e., during the Martian year, after a launch in 2007. With these new seismic data, much of the unknown internal structure of Mars will be discovered. Possibly little more than one century after the discovery of the Earth’s core by Oldham (1906), the core of a second telluric planet might therefore be characterized. Acknowledgements The VBB seismometer was developed by IPGP-DT= INSU and SODERN inc. under a CNES R& T program,

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and with the academic technical sta supported by CNRSINSU. R. Tesseraud, D. Desmet, Y. Bouchet, C. Germon, J. Pinassaud also contributed to the VBB design. The VBB IPGP team is remembering the strong support of J. RunavotX . We thank G. Debouzy, Y. Dancet and G. Pont for active contribution during the dierent phases of the R&T program, and V. Courtillot, J.L. Le Mouel and C. Jaupart for the continuous eort in the development of planetary geophysics at IPGP. The french team thanks R. Bonneville, F. Rocard, O. Marsal, Y. Langevin, C. Sotin, J.P. Bibring, J.L. Birck for all their eorts during the preparation phase of the French Mars program and J.E. Blamont for his promotion of the program and for motivating discussions. Very constructive reviews were provided by E.A. Okal and K.F. Ma. This is IPGP contribution No. 1676 and Netlander contribution No. 3.

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