Moho depth of the European Plate from teleseismic receiver functions
- 1.5k Downloads
Crustal structure and the Moho depth are exceptionally well known beneath Europe. The first digital, high-resolution map of the Moho depth for the whole European Plate was compiled in 2007 and recently published in Geophysical Journal International. In the past few years, considerable developments have taken place in the receiver function techniques. Different receiver function techniques provide new, independent information, in particular on the S-wave velocity distribution in the crust and on the Moho depth. This gives an opportunity to compare the Moho depth from the Moho depth map of the European Plate (H MM) and the Moho depth from receiver function studies (H RF). Herein, we also compile and analyze the uncertainty of the crustal thickness determinations data obtained with receiver function analysis. The uncertainty is found to be ±2 km for 20-km-thick crust and about ±4 km for 60-km-thick crust. Comparison of the Moho depths shows an approximately linear trend between H RF and H MM. For the Moho depth of 30–40 km, the values are approximately equal, while for thin crust, H RF is about 5 km shallower than H MM, and for thick crust, it is about 5 km deeper than H MM. Possible reasons for this, the observed discrepancy between the Moho depths HMM and HRF, are discussed.
KeywordsTeleseismic receiver function Crustal thickness Moho discontinuity European Plate
The boundary between crust and upper mantle was discovered by the Croatian seismologist Andrija Mohorovičić, and it is defined as a distinct discontinuity in the seismic wave velocities (Mohorovičić 1910). Studies during the subsequent 100 years showed that this discontinuity was a worldwide phenomenon, and it was named the Mohorovičić discontinuity, Moho, or M-discontinuity. Today, the seismologically defined Earth’s crust is understood to be the outer shell of our planet in which the velocity of P waves is less than about 7.6 km s−1 and S-wave velocity is less than about 4.4 km s−1 (e.g., Meissner 1986). In general, P-wave velocity in the lower crust is about 7 km s−1 and in the uppermost mantle about 8 km s−1. Thus, the P-wave velocity contrast at the Moho discontinuity is quite large, up to 1–1.5 km s−1. S-wave velocity in the lower crust is about 4 km s−1 and in the uppermost mantle about 4.6 km s−1, and the corresponding contrast at the Moho is over 0.5 km s−1. This indicates a significant change in the elastic parameters, resulting from a significant change in the rock types between crust and uppermost mantle.
In the past decades, and in particular in the past few years, considerable developments in the receiver function (RF) techniques have taken place. A large number of new Moho depth (crustal thickness) determinations, beneath permanent broadband seismic stations and temporary arrays used in passive experiments, motivated us to compare the Moho depths H MM from the Moho depth map of the European Plate (Grad et al. 2009) and the Moho depths H RF from receiver function analysis.
2 Receiver function data
Sources of receiver function data for crustal thickness (Moho depth) used in this paper
References for receiver function data
East European platform
Mediterranean Sea and Alpine area
Bertrand and Deschamps (2000), Kummerow et al. (2004), Li et al. (2003), Lombardi et al. (2008), Marone et al. (2003), Mele and Sandvol (2003), Mele et al. (2006), van der Meijde et al. (2003), Piana Agostinetti and Amato (2009), Zor et al. (2006)
Atlantic and polar regions
European Plate surroundings
Al-Damegh et al. (2005), Angus et al. (2006), Dahl-Jensen et al. (2003), Doloei and Roberts (2003), Mohsen et al. (2005), Paul et al. (2006), Radjaee et al. (2010), Saunders et al. (1998), Sodoudi et al. (2009), Taghizadeh-Farahmand et al. (2010), Tezel et al. (2010), Weber et al. (2004), Zor et al. (2003)
3 Discussion of the deviations, uncertainties, and errors
In this section, we discuss the uncertainties and errors in the Moho depth determinations, as well as possible reasons for the difference between Moho depths H MM and H RF.
In construction of the Moho depth map for the European Plate (Grad et al. 2009), different types of data were used, and in most cases, the Moho depths were consistent. The uncertainties are different for different seismic techniques and can be different even for the same technique in different experiments and areas. The lowest uncertainty is in the order of 5% for new, modern, good-quality seismic refraction profiles, available in digital form (e.g., models obtained by ray tracing modeling). This gives about ±2 km uncertainty for 40-km-thick crust. Older, reinterpreted, compiled, and/or manually digitized profiles have lower quality, with uncertainty in the order 6–8%. The highest uncertainty (about 20%) was attributed to results obtained from surface waves and gravity modeling. For all the data points used to construct the Moho depth map (for the same latitude ϕ and longitude λ), corresponding values of uncertainties (in kilometers) were attributed. The map of the Moho depth uncertainty was constructed using the same projection, transformation, filtering, etc., as the Moho depth map (Grad et al. 2009; in digital form also at WWW pages). The uncertainty of the Moho depths H MM ranges from ±2 to ±10 km. The lowest uncertainty in the order of ±2–4 km is associated with the continental part of Western, Central, and Northern Europe. Similar values are associated with the oceanic crust. However, since the oceanic Moho depth is about 10–15 km, the relative uncertainty is larger. The largest uncertainty is observed for Greenland and Africa–Arabia transition, where the resolution of the present map is the lowest.
The data sets used in the construction of the Moho depth map for the European Plate (Grad et al. 2009) were dominated by results from seismic profiles and their compilations (local and regional maps). Thus, the data density was relatively higher than for individual stations used for the Moho depth determination from receiver function analysis. The Moho map for the British Isles and surrounding areas constructed from many seismic profiles and gravity modeling (Kelly et al. 2007) had much larger weight than limited number of RF data (Tomlinson et al. 2006)—their influence on final map is rather small.
Although both seismic refraction and receiver function methods can reveal the structure of the Earth’s crust and uppermost mantle, their underlying principles are quite different. In seismic refraction method, we investigate the structures using mostly P waves (refracted and reflected) going down into the crust. In receiver function method, we investigate the structures using upgoing S waves from teleseismic events (P-to-S converted at the Moho and other discontinuities). This is the first source of deviations between the methods. On the other hand, this gives a chance to create a common, integrated P- and S-wave velocity model for the same area. For example, known P-wave velocity model from a refraction study can be used as a good starting point for S-wave modeling with receiver function method.
Different techniques are applied in the interpretation of RF. In the first step of RF processing, original components of seismogram (Z, N, E) are rotated into vertical, radial, and tangential (Z, R, T) components or into ray-parameter coordinate system (L, Q, T) which is useful to separate different wave types P, SV, and SH (Vinnik 1977; Geissler et al. 2008). In the estimation of the Moho depth, the delay time of P-to-S converted waves is compared to the direct P-wave (e.g., Langston 1979; Ammon 1991; Cassidy 1992). To investigate the 1-D S-wave velocity structure beneath the station the time-domain inversion methods are applied to the radial receiver function: linear, semi-linear, stochastic inversion—see, e.g., Ammon et al. (1990) and Hetényi and Bus (2007). However, the result of inversion depends on the starting model, in particular for noisy data. In order to receive independent results, we can run the inversion with many different starting models and then stack the results to calculate one mean model of the structure (e.g., Wilde-Piórko et al. 2002). Forward modeling using trial-and-error method may find a simple model which could well explain the observed receiver functions. In forward modeling, a generalization and simplification of the models could complement and correct the results of inversion.
In RF technique, the Moho depth is usually projected beneath the station. However, the observed converted and reverberated phases Ps, PsPmP + PpSmP and PsSmP (P, p—longitudal, S, s—shear waves, m—reflection from the Moho), come from the wide area around the station, which at the Moho is from a few tens up to a hundred kilometers in diameter. In seismic refraction method, the system of reciprocal travel times permits for determination of dipping boundaries along profiles. For 2-D refraction profiles, however, modeling does not take into account out-of-plane refractions and reflections, which could occur in structurally complex regions. In such case, 3-D approach should be used.
In the past decade, the RF method by Zhu and Kanamori (2000) for simultaneous determination of the Moho depth and Poisson’s ratio in the crust has become very popular. An average Poisson’s ratio σ in the crust and the Moho depth H are estimated in a grid search over the σ−Η space, and the (σ, H) pair which is in the closest agreement with the observed converted and reverberated phases is determined. This method was found to be very sensitive to the average Poisson’s ratio in the crust, but it works only for sharp Moho (with large contrast of elastic parameters), when clear Moho conversions and their associated multiples are observed (e.g., Kumar et al. 2007). Weak point of this method is the assumption of an average V P velocity in the crust. As result, the average Poisson’s ratio σ in the crust (or V P/V S ratio) may not be adequate and sufficient for the whole crust. For example, in the East European craton V P/V S ratio is about 1.67, 1.73, and 1.77 in the upper, middle, and lower crusts, respectively. This means that the S-wave velocity in the crust is not much differentiated and the relative contrast of V S velocity at the Moho is larger than that for P waves.
The accuracy of the seismic velocity structure determined using teleseismic receiver functions is a complex problem since it strongly depends on many factors. Usually many tens or hundreds good-quality seismograms of teleseismic events are recorded by a single permanent broadband station (e.g., Piana Agostinetti and Amato 2009; Geissler et al. 2008). However, even a large number of seismograms do not guarantee a good azimuthal coverage (e.g., Radjaee et al. 2010). This is particularly important when the structure beneath the station could not be approximated by a 1-D model, i.e., in the case of dipping layers or anisotropy. These effects are usually visible in the azimuthal distribution of tangential components of RF. For temporary passive experiments, this is even bigger problem, due to the limited number of events. Usually during 1 year of a campaign, only about few tens of good quality records, with magnitude M ≥ 5.5, are collected (Kozlovskaya et al. 2008; Gregersen et al. 2006; Wilde-Piórko et al. 2008). Stacking of RF results is smoothed as that covers a wide range of azimuths and distances.
The frequencies of P waves recorded in seismic refraction studies are in the range of 5–10 Hz, while stacked receiver functions have pulses with smaller frequencies, about 1–2 Hz. Also, filtration of the receiver functions influences their quality and frequency (low-pass Gaussian filtration with parameter of 2–4, influence of water-level parameters). Because of that, the frequency resolution is better for the refraction method.
For both seismic refraction and receiver function methods, crustal anisotropy is a difficult problem to solve. To detect anisotropy in a seismic refraction studies, a dense system of 3-D recordings is needed. In receiver function, some opportunities give interpretation of transversal component of RF.
The velocity structure and crustal thickness determined for the same broadband station in different receiver function studies can also differ. This can stem from the use of different data sets (in particular a small number of seismograms) and interpretation techniques. For permanent stations in Bohemian massif, the crustal thicknesses presented in the papers by Wilde-Piórko et al. (2005) and Geissler et al. (2008) differ only by about ±1 km; however, in other regions, differences can reach more than ±5 km, as is the case for Greenland (Dahl-Jensen et al. 2003; Kumar et al. 2007).
In seismic refraction and wide angle reflection method observed PmP waves are weak for weak contrast at the Moho (gradient Moho zone). At the same time, good quality refracted Pn waves still give an opportunity for proper identification and determination of the crust–mantle transition. Ps wave converted at the same gradient Moho transition gives weak receiver function response, which weakens the precision of the Moho depth H RF determination.
4 Summary and conclusions
Herein, we have compared the Moho depths from the Moho depth map of the European Plate by Grad et al. (2009) and the Moho depth values obtained with RF methods by several authors. The best coverage of the crustal thickness from RF determinations comes from the continental part of the Western and Central Europe. There is practically no RF data available beneath the ocean and the seas. Based on 393 values of crustal thickness, we have determined linear relation (1) of the Moho depths H RF and their error estimates (Fig. 3a). Most of the error estimates seem rather optimistic, when compared to the total range of the values. Linear fit (2) between H RF and H MM (Fig. 3b) shows that the receiver function method gives shallower Moho when the Moho depth is smaller than 30 km and deeper Moho (than other methods) when the Moho depth is more than 40 km. It should be noted that in some cases for inversion of RF, instead of regional crustal velocity structure, global velocity models (e.g., iasp91, Kennett and Engdahl 1991), or other reference models with global crustal thickness of about 35 km are used (see for example Sodoudi et al. 2009; Hetényi and Bus 2007; Paul et al. 2006). This could be the main reason for the observed discrepancy between the Moho depths H MM and H RF. Distribution of the Moho depth values based on RF (in the range of 10–70 km) is clearly wider than the distribution of the Moho depths from other seismic methods (in the range of 15–55 km) for the study area.
The authors wish to thank Finnish Academy of Science and Letters, Väisälä Foundation for financial support. Geographic data handling and plotting was done with GMT software by P. Wessel and W.H.F Smith. The linear data fitting was done with program Glove from New Planet Software. The authors are grateful to four anonymous reviewers for helpful comments and raising numerous questions that were not answered in the first version of manuscript. The authors are grateful to Dr. Emilia Koivisto for improving the English language.
This article is distributed under the terms of the Creative Commons Attribution Noncommercial License which permits any noncommercial use, distribution, and reproduction in any medium, provided the original author(s) and source are credited.
- Amante C, Eakins BW (2009) ETOPO1 1 Arc-Minute Global Relief Model: procedures, data sources and analysis. NOAA Technical Memorandum NESDIS NGDC-24, 19ppGoogle Scholar
- Ammon CJ (1991) The isolation of receiver effects from teleseismic P waveforms. Bull Seism Soc Am 81:2504–2510Google Scholar
- Ansorge J, Blundell D, Mueller St (1992) Europe’s lithosphere—seismic structure. In: Blundell DJ, Freeman R, Mueller St (ed) A continent revealed—the European geotraverse. Cambridge University Press, Cambridge, pp 33–69Google Scholar
- Cassidy JF (1992) Numerical experiments in broadband receiver function analysis. Bull Seism Soc Am 82:1453–1474Google Scholar
- Czuba W, Grad M, Guterch A (1999) Crustal structure of north-western Spitsbergen from DSS measurements. Polish Polar Res 20(2):131–148Google Scholar
- Czuba W, Grad M, Guterch A, Majdański M, Malinowski M, Mjelde R, Moskalik M, Środa P, Wilde-Piórko M, Nishimura Y (2008) Seismic crustal structure along the deep transect Horsted’05, Svalbard. Polish Polar Res 29(3):279–290Google Scholar
- Kortström J, Wilde-Piórko M, Tiira T, Komminaho K (2006) Receiver function analysis of the broad band data of Finnish Seismograph Network. Proceedings of the 37th Nordic Seminar on Detection Seismology, University of Iceland, AbstractsGoogle Scholar
- Kozlovskaya E, Kosarev G, Aleshin I, Riznichenko O, Sanina I (2008) Structure and composition of the crust and upper mantle of the Archean–Proterozoic boundary in the Fennoscandian shield obtained by joint inversion of receiver function and surface wave phase velocity of recording of the SVEKALAPKO array. Geophys J Int 175:135–152. doi: 10.1111/j.1365-246X.2008.03876.x CrossRefGoogle Scholar
- Meissner R (1986) The continental crust—a geophysical approach. Int Geophys Ser Acad Press Inc Orlando 34:1–426Google Scholar
- Mohorovičić A (1910) Potres od 8.X.1909. Godišnje izvješće zagrebačkog meteorološkog opservatorija 9(4/1), 1–56 (and English translation in 1992: Earthquake of 8 October 1909). Geofizika 9:3–55Google Scholar
- Olsson S, Roberts RG, Böðvarsson R (2008) Moho depth variation in the Baltic Shield from analysis of converted waves. GFF 130(3):113–122, Stockholm ISSN 1103–5897Google Scholar
- Silveira G, Vinnik L, Stutzmann E, Farra V, Kiselev S, Morais I (2010) Stratification of the Earth beneath the Azores from P and S receiver functions. Earth Planet Sci Lett. doi: 10.1016/j.epsl.2010.08.021
- Weber M, Abu-Ayyash K, Abueladas A, Agnon A, Al-Amoush H, Babeyko A, Bartov Y, Baumann M, Ben-Avraham Z, Bock G, Bribach J, El-Kelani R, Förster A, Förster H-J, Frieslander U, Garfunkel Z, Grunewald S, Götze HJ, Haak V, Haberland Ch, Hassouneh M, Helwig S, Hofstetter A, Jäckel K-H, Kesten D, Kind R, Maercklin N, Mechie J, Mohsen A, Neubauer FM, Oberhänsli R, Qabbani I, Ritter O, Rümpker G, Rybakov M, Ryberg T, Scherbaum F, Schmidt J, Schulze A, Sobolev S, Stiller M, Thoss H, Weckmann U, Wylegalla K (2004) The crustal structure of the Dead Sea transform. Geophys J Int 156(3):655–681Google Scholar
- Wessel P, Smith WHF (1991) Free software helps map and display data. EOS, Trans AGU 72(41):445–446Google Scholar
- Wessel P, Smith WHF (1998) New, improved version of Generic Mapping Tools released. EOS, Trans AGU 79(47):579Google Scholar
- Wilde-Piórko M, Geissler WH, Plomerová J, Grad M, Babuška V, Brückl E, Cyziene J, Czuba W, England R, Gaczyński E, Gazdova R, Gregersen S, Guterch A, Hanka W, Hegedűs E, Heuer B, Jedlička P, Lazauskiene J, Keller GR, Kind R, Klinge K, Kolinsky P, Komminaho K, Kozlovskaya E, Krüger F, Larsen T, Majdański M, Málek J, Motuza G, Novotný O, Pietrasiak R, Plenefisch Th, Růžek B, Sliaupa S, Środa P, Świeczak M, Tiira T, Voss P, Wiejacz P (2008) PASSEQ 2006–2008: passive seismic experiment in Trans-European suture zone. Stud Geophys Geod 52:439–448CrossRefGoogle Scholar
- Zhang J, Langston CA (1995) Dipping structure under Dourbes, Belgium, determined by receiver function modeling and inversion. Bull Seism Soc Am 85:254–268Google Scholar