Download Seismologia ja maan rakenne 762321A Seismology and

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)