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Seismologia ja maan rakenne 762321A Seismology and structure of the Earth Lecture 6: Inner structure of the Earth: Seismic methods to study lithosphere and the Earth’s crust References: Stein, S., Wyession, M., 2003. An introducHon to seismology, earthquakes, and Earth structure. Blackwell Publishing. Aki, K., and P.G. Richards, 2002, QuanHtaHve Seismology, 2nd EdiHon, 685 pp., University Science Books, Sausalito. New Manual of Seismological Observatory PracHce (NMSOP-‐2): hZp://bib.telegrafenberg.de/publizieren/vertrieb/nmsop/ hZp://www.soes.soton.ac.uk/teaching/courses/oa405/GY405/crustal/Lecture_3 SEISMIC METHODS ARE USING SEISMIC WAVES IN ORDER TO OBRAIN THE STRUCTURE OF SUBSURFACE Types of Waves • Seismic Wave – Body Waves • Primary or p-‐wave – Compression wave • Secondary or s-‐wave – Transverse wave – Surface • Love wave • Rayleigh wave SEISMOGRAMS: LOCAL, REGIONAL AND GLOBAL SEISMOLOGICAL STUDIES: 1. Local seismograms occur at distances up to about 200 km. The main focus is usually on the direct P waves (compressional) and S waves (shear) that are confined to Earth’s crust. Surface waves are not prominent although they can someHmes be seen at very short periods. 2. Regional seismology studies examine waveforms from beyond ∼200 up to 2000 km or so. At these distances, the first seismic arrivals travel through the upper mantle (below the Moho that separates the crust and mantle). Surface waves become more obvious in the records. Analysis of conHnent-‐ sized data sets is an example of regional seismology, such as current USArray project to deploy seismometers across the United States. 3. Global seismology is at distances beyond about 2000 km (∼20◦), where seismic wave arrivals are termed teleseisms. This involves a mulHtude of body-‐wave phase arrivals, arising from reflected and phase-‐converted phases from the surface and the core-‐mantle boundary. For shallow sources, surface waves are the largest amplitude arrivals. Data come from the global seismic network (GSN). SEISMIC RAYS: REFLECTION AND REFRACTION OF BODY WAVES (RAY APPROXIMATION) times along the ray, they will be separated by di↵erent distances in the di↵erent RAY PARAMETER AND SLOWNESS layers, and we see that the ray angle at the interface must change to preserve the v1 v AY PATHS FOR LATERALLY HOMOGENEOUS MODELS 2 f the wavefronts the interface. Figure 3.2: across A plane wave crossing a horizontal interface between two homogeneous halfspaces. The higher velocity in the bottom layer causes the wavefronts to be spaced further apart. he case illustrated the top layer has a slower velocity (v1 < v2 ) and ngly In larger u2v).elocity The parameter maylarger be express Fig. 3.2 tslowness he top layer h(u as 1a s> lower (v1 <ray v2) and a correspondingly the slowness (u1 > u2). slowness and m ray from the overtical within each The ray parameter ay bangle e expressed in terms f the slowness and ray angle layer: from the verHcal within each layer: p = u1 sin ✓1 = u2 sin ✓2 . hat this is simply the seismic version of Snell’s law in geometrical opt han the layer above. The ray parameter p 16 CHAPTER 3. nstant and we have RAY THEORY v u1 sin ✓1 = u2 sin ✓2 = u3 sin ✓3 . (3.5) city continues to increase, ✓ will eventually z and the ray will be traveling horizontally. = 90 Figure 3.3: Ray paths for a model with a continuous velocity increase with depth will curve back toward the surface. The ray turning point is defined as the lowermost point on the ray also truepath, forwhere continuous gradients 3.3). If we let the slown the ray directionvelocity is horizontal and the incidence (Fig. angle is 90 . the m odel, in which velocity increases with depth. T takeo↵ angle be ✓ , we have ace be uConsider and the 0 we let the slowness at the surface be u0 and the takeoff angle be θ , we If 0 have 0 dT/dX= p = ray parameter u0 sin ✓0 = p = u=sin ✓. slowness horizontal (3 = constant for given ray When θ = 90◦ we say that the ray is at its turning point and p = u , where u is the tp tp 90 slowness we sayathat the ray is at its turning point and p = u , where utp is tp t the turning point. X Since velocity generally increases with depth in Earth, the slowness decreases with depth. Figure 3.4: A travel time curve for a model with a continuous velocity increase with depth. t theSmaller turning Since velocity generally with depth in Ear Each point on the curve results from a di↵erent ray path; of the w travel time ray ppoint. arameters are more steeply dipping at the the slope sincreases urface, ill turn dcurve, eeper in Earth, dT /dX, gives the ray parameter for the ray. and generally travel farther. ss decreases with depth. Smaller ray parameters are more steeply dipp into its horizontal and vertical components. The length of p = u sin s Figure 3.3: Ray paths for a model with a continuous velocity increase with depth will curve = u cos backcomponents. toward the The ray length turning point is defined as the /dX, givesand the vertical ray parameter forsurface. the ray. p =lowermost u sin point on the ray its dT horizontal The of where the ray direction is horizontal and the incidence angle is 90 . CHAPTER path, 3. RAY THEORY sx given by u, the local slowness. The horizontal comT p = u sin horizontal andisvertical TheInlength ent,into sx , its of the slowness the raycomponents. parameter p. an of u sx s isway, givenwebymay u, the localthe slowness. horizontal ogous define vertical The slowness ⌘ by com- for a model with a continuous velocity increase with depth will curve e. The ray turning point is defined as the lowermost point on the ray 2 ection is horizontal and the incidence angle is 90 . 2 1/2 = u cos he turning point, p⌘ = and✓ ⌘==(u0. p ) . (3.7) = u cos ponent, sx , of the slowness is the ray parameter p.dT/dX= In pan = ray parameter 2 2 1/2 horizontal slowness us 90= (u analogous⌘ way, we ✓=may definepthe slowness = u cos ) vertical . (3.7)⌘ by== constant for given ray sz s X s z A travel for a model with a continuous increase travel with depth. point, Figure p = u3.4:and ⌘ =time 0. curve We At canthe useturning these relationships to derive integral expressions tovelocity compute Each point on the curve results from a di↵erent ray path; the slope of the travel time curve, gives the ray 2parameter for the ray. dT/dX= p = ray parameter dT /dX, = horizontal slowness = constant for given ray and distance along a particular rayto. derive For aintegral surface-to-surface raycompute path, the We can use these relationships expressions to travel into rits andtotal vertical components. The lengthbof 2 .istance For and aX(p) surface-‐to-‐surface ay horizontal path, the d X(p) is given y time distance along a particular ray For a surface-to-surface ray path, the l distance is given by sx s is given by u, the local slowness. The horizontal comp = u sin ponent, sx , ofZthe the ray parameter p. In an zp slowness isdz total distance X(p) is given by me curve for a model with a continuous velocity increase with depth. analogous way, we may define the vertical slowness ⌘ by X(p) = 2p e results from a di↵erent ray path; the slope of the travel time curve, arameter for the ray. 0 Z 2zp 1/2 2 dz 2p ) 2 1/2 . = u cos X u ⌘ = (u u cos(z) ✓ = (u p ) . (3.7) X(p) = 2p . p = u sin d vertical components. The length of 2 (z) 2 )1/2 stime 0= u(u p z re s is the turning point depth. The total surface-to-surface travel s At the turning point, p and ⌘ = 0. p ocal slowness. Thez horizontal x where is the tcomurning point depth. The total surface-‐to-‐surface travel Hme is p (3.8) s is (3.8) cos ✓ = (u2 p2 )1/2 . p = u and ⌘ = 0. (3.7) = u cos Zdepth. u can where s is the turning point total tosurface-to-surface travel time is zp these The We use relationships derive integral expressions to compute travel u2 (z) wness is the ray parameter p. In an p slowness ⌘ by ay define the vertical T (p) 2s distanceZalong dz. (3.9)the 2 . For a surface-to-surface ray path, time= and ray 21/2 zp a particular 2 2 u (z) 0 (u (z) s T (p) =X(p) 2 is given byp ) dz. (3.9) total distance 1/2 z relationships to derive integral expressions to compute travel ee section 4.22 of ITS for details ong a particular ray . For a surface-to-surface ray path, the s given by 2 See section 4.2 of ITS for details 0 (u2 (z) X(p) = 2p 2 p zp) Z 0 dz (u2 (z) p2 )1/2 . (3.8) PROPAGATION OF SEISMIC WAVES IN A SPHERICAL EARTH Wysession, 2003, Fig 1-1-02 Sivu 7 apallon rakenne 2002 a b - Stein& 16 a) Constant velocity; b) Gradient velocity va 7.17. P-tyypin seismisiä säteitä erilaisille hypoteettisille maapalloillle: a) nopeudeltaan homogeeninen Maa; b) jatkuvasti syvyyden mukana kasvava nopeus. 9 Forward Branch Shadow Zone Forward Branch Backward Branch PcP Backward Branch Forward Branch Shadow Zone PKP Forward Branch PcP P Forward Branch Backward Branch Forward Branch ・ 1912 Gutenberg observed shadow zone 105o to 143o ・ 1939 Jeffreys fixed depth of core at 2898 km (using PcP) Shadow Zone secondary waves, mainly phases, which have been reflected or converted at the surface of the Earth or at the core-mantle boundary(CMB). Fig. 2.52 depicts a typical collection of possible primary and Jeffreys-‐Bullen travel-‐Hme curve (1940) secondary ray paths together with a three-component seismic record at a distance of D = 112.5° that relates to the suit of seismic rays shown in red in the upper part of the cross section through the Earth. The phase names are standardized and in detail explained in IS 2.1 http://earthguide.ucsd.edu/mar/dec5/earth.html AK135 1-‐D EARTH MODEL: α is P-wave velocity β is S-wave velocity, ρ is density. Q is a factor that defines elastic and anelastic properties KenneZ, B.L.N. Engdahl, E.R. & Buland R., 1995. Constraints on seismic velociHes in the Earth from travel Hmes, Geophys. J. Int, 122, 108-‐124 1-‐D PRELIMINARY EARTH MODEL (PREM), Dziewonski and Anderson (1981) Fig. 2.79 Radial symmetric reference models of the Earth. Top: AK135 (seismic wave speeds according to Kennett et al., 1995), attenuation parameters and density according to Montagner and Kennett (1996); Bottom: PREM (Dziewonski and Anderson, 1981). - and : P- and S-wave velocity, respectively; - density, Q and Q = Q - “quality factor” Q for P LITOSPHERE AND THE EARTH’S CRUST The lithosphere, which is the rigid outermost shell of a planet (the crust and upper mantle), is broken up into tectonic plates. The Earth's lithosphere is composed of seven or eight major plates (depending on how they are defined) and many minor plates. Where the plates meet, their relaHve moHon determines the type of boundary: convergent, divergent, or transform. Earthquakes, volcanic acHvity, mountain-‐building, and oceanic trench formaHon occur along these plate boundaries. DEEP DRILLING: DIRECT PROBING OF THE EARTH’S CRUST 1970-‐1989: Kola Superdeep Hole (SG-‐3): 12 262m 1987-‐1994: KTB (KonHnentales Tievohrprogramm der Bundesrepublik Deutschland) borehole (Germany): 9101 m 2012: Z-‐44 Chayvo well (Sakhalin, Russia) is 12 376 m long (horizontal drilling, Exxon Newegas Ltd.) InternaHonal ScienHfic ConHnental Drilling Program (hZp://www.icdp-‐ online.org/home/) Outokumpu deep drilling project (2004-‐2005) 2 516 m. Ocean Deep Drilling Program (www.iodp.org) Deepest hole: Hole 504B: 2.1 km JOIDES ResoluHon Riser-‐equipped Chikyu Mission-‐Specific Plaxorms IODP uses mulHple drilling plaxorms to access different subseafloor environments during research expediHons. Three Science Operators in the United States, Japan, and Europe manage these plaxorms. SEISMIC METHODS USED TO STUDY THE EARTH’S CRUST AND UPPER MANTLE: 1) Deep seismic sounding (controlled-‐source deep seismic sounding), in which the acHve source of seismic energy is used: blast, mechanic vibrators, airgun,) a) ReflecHon sounding b) RefracHon sounding NB: in deep seismic sounding both reflecHon and refracHon P-‐ and S-‐waves are used (wide-‐angle reflecHon and refracHon method) 2) Seismic tomography using body from sources at local distances (P-‐ and S-‐ waves and surface waves). Both explosions and local earthquakes can be used as sources of seismic energy 3) Passive seismic methods using seismic waves from teleseismic earthquakes a) Receiver funcHon method that uses converted waves (P-‐ and S-‐receiver funcHons) b) Surface wave tomography c) Ambient noise tomography, in which seismic noise is used as a source of seismic energy Snell’s law sin i V 1 = sin r V 2 (1) Snell’s law: pre-critical incidence If V2>V1, then angle r is larger than i: Pre-Critical incidence Reflection and refraction Snell’s Law: sin iP VP1 sin RP VP1 sin rP VP 2 p where p is the ray parameter and is constant along each ray. Reflection and transmission coefficients for a specific impedance contrast Applied Geophysics – Refraction I Snell’s law If i grows, r approaches 90o Snell’s law Critical Refraction At the Critical Angle of incidence ic, angle of refraction r = 90o sin ic V 1 = sin 90 V 2 V1 sin ic = V2 (2) V1 ic = sin V2 (3) −1 Critical incidence When rP = 90° iP = iC the critical angle sin iC VP1 VP 2 The critically refracted energy travels along the velocity interface at V2 continually refracting energy back into the upper medium at an angle iC a head wave Reflection and transmission coefficients for a specific impedance contrast Applied Geophysics – Refraction I Post-Critical incidence The angle of incidence > iC No transmission, just reflection Reflection and transmission coefficients for a specific impedance contrast Applied Geophysics – Refraction I 2/4/11 Lecture 2003 5 Horizontal interface Traveltime equations Direct wave: T Head wave x V1 Head wave: T TSB TDD ' TBD T 2h1 V1 cos ic T x V2 x 2h1 tan ic V2 2h1 V22 V12 V2V1 T = ax + b slope: 1/V2 intercept: gives h1 Applied Geophysics – Refraction I Horizontal interface Crossover distance, xco Where the direct and head wave cross. Their travel times are equal: xco V1 xco V2 xco 2h1 2h1 V22 V12 V2V1 V2 V1 V2 V1 Another approach to obtaining layer thickness Applied Geophysics – Refraction I Horizontal interface Reflections The critical reflection is the closest head wave arrival. At shorter offsets there are low amplitude reflections (used in reflection seismology). At greater offsets there are wide-angle reflections. Applied Geophysics – Refraction I Reversing lines …shooting to a line of geophones from both ends For horizontal layers the traveltime curves are symmetrical For dipping layers layer velocities appear different for each end – the dip and true velocity can be determined from the updip and down-dip velocities Applied Geophysics – Refraction I VELOCITY GRADIENT If there are velocity gradients in the layers, so that velocity increases with depth, the refracted arrivals are much stronger and Hme-‐ distance plots are curved Apparent velocity • This is simply the inverse of the slope of the Hme-‐ distance graph. • For a structure in which velocity varies only with depth, the apparent velocity of a refracted arrival is equal to the true velocity at the depth where the ray turns; • More usually it is an approximaHon to the true velocity. REDUCED TRAVEL TIME Wide-‐angle seismic data are normally ploZed against reduced travelHme for compactness. In a reduced plot, signals travelling at the reducHon velocity appear horizontal, while slower arrivals have posiHve slope and faster arrivals negaHve slope Tred = T – X/Vred Fowler FigTHE 9.20 REFLECTED AND REFRACTED WAVES INSIDE EARTH’S CRUST: Pg PmP Pn crust mantle Pg and Sg are the P-‐ and waves refracted inside the crust PmP ja SmS are reflecHon from the crust-‐mantle boundary Pn and Sn are the waves refracted in the mantle (head waves) AN EXAMPLE OF WIDE-‐ ANGLE REFLECTION AND REFRACTION RECORD SECTION DATA IN REDUCED TIME (ReducHon velocity is 7 km/s) REDUCED TRAVEL TIMES RECORD SECTION: tR = t – X/VreducHon, B03304 MUSACCHIO ET AL.: SUPERIOR PROVINCE LITHOSPHERIC STRUCTURE Figure 2. Examples of record sections for crustal phases plotted in an offset distance reference frame. Each trace is normalized to its maximum amplitude, and data have been filtered using a passband of B03304 A larger-‐scale conHnental dataset shows how with large explosive shots, refracHon profiles can be 100s or even 1000s of km long and can send seismic energy many kilometres down into the upper mantle (figure from T. J. Henstock) An ocean boZom hydrophone (OBH) dataset, acquired using an airgun source, shows the phases visible in oceanic crust. Note that oceanic wide-‐angle profiles are normally much shorter because the crust is thinner. P2 and P3 mark energy turning in oceanic Layers 2 and 3, respecHvely (energy turning in Layer 1, the sediments, is owen not clearly seen). R2 is a reflecHon from the top of Layer 2 (i.e., oceanic basement). S2 and S3 are energy that has been converted to shear waves on arrival at the top of the crust (figure from T. J. Henstock)