2002 denali earthquake - sgcobinsus.files.wordpress.com · 02/08/2017 · 2 seismology – what is...

1

Seismology

ieso 2010

2002 Denali Earthquake

From: http://w

ww

.citiesoflight.net/AlaskaQuake.htm

l

Note: Anim

ation only visible in powerpoint

2

Seismology – What is a WaveSeismology is primarily concerned with determining the structure of the earth – on all scales.

To accomplish this it uses the ability of seismic waves to propagate through the earth.

What is a wave?

From: Mussett and Khan, Looking into the earth

Wavelength: The length between two crests or troughs

Amplitude: The maximum height relative to the zero position

Frequency: The number of “wavelenghts” that pass a point in one second

It is important to note that although the waves travels through the earth, the material itself does not – in the same way as water.

3

Frequencyλ*fv =

Velocity = frequency x wavelength

Frequencies in seismology range from less than 10 to perhaps 100s of Hz (cycles per second).

Ancient Seismographs

From

Tar

buck

and

Lutg

ens,

200

5

The Chinese made the first seismographs over 2000 years ago. This seismograph will tell the observer the direction of the first motion of the ground during an earthquake

4

Earthquake

Nowadays, if an earthquake occurs, how do we detect it, and how do the waves travel – in the next few slides I will attempt to answer these questions.

From

Tar

buck

and

Lutg

ens,

200

5

Detecting Seismic WavesIt is easy to visualize the motion of an ocean wave, and how we might measure its wavelength, amplitude, and frequency, but how do we do this with the earth?

On the left are two basic seismometers. When the ground moves the pendulums moves either horizontally (top) or vertically (bottom). The movement is damped so that subsequent motion of the earth is not obscured by the pendulum oscillating.

On the right is a more conventional seismometer. Motion of the magnet through the coil generates a current in the coil which is amplified and recorded.

Three seismometers oriented vertically, N-S and E-W are required to measure the full ground motion.

Figures from: Mussett and Khan, Looking into the earth

5

Building A SeismometerA basic seismometer is actually very simple. Shown here is a simple seismometer that with the addition of some electronics (amplifier etc) will happily record earthquakes. Refer to http://www.iris.edu/edu/AS1.htm for more details.

Snell’s Law

''

2

1

2

1

AABB

vv

==λλ

''sin

''sin

2

1

ABAAi

ABBBi

=

=

The earth is not a uniform sphere. Broadly speaking, it is made up of layers.When wave fronts cross from one rock type into another with a higher velocity they turn.

Wavefront

The time between successive wavefronts remains unchanged, so the wavelength must increase in the second rock in proportion to the increase in velocity.

Trigonometry tells us that:

21 sin'

sin''

iAA

iBBAB ==Rearranging

gives:

From

Kea

rey

et a

l., 2

002

From M

ussesttand Khan, 2000

6

Snell’s Law

2

2

1

1 sinsinv

iv

i=

As BB’ and AA’ are in proportion to the velocities v1 and v2, the equation can be rearranged to

Snell’s Law

o

o

i

i

37sin45sin

5sin

437sin

2

2

=

=So i2 = 48.8o

Answer the following question:A ray traveling in a rock with a seismic velocity of 3 km/s encounters an interface with a rock of 4 km/s at an angle of 45o. At what angle from the normal does it leave the interface?

From M


Snell’s Law – Multiple Horizons

3

3

2

'2

2

'2

1

'1

sinsin

sinsin

vi

vi

vi

vi

=

=

constantsinsinsin

3

3

2

2

1

1 ===v

iv

iv

i

As i’1 = i1, I’2 = i2, and so on

From M


The ratio (sin i/v) thus remains unchanged

7

Snell’s Law – Curved LayersHowever, in the real earth, the layers are curved, so it is not true that i2 = I’2, etc.

From M


The differences between the angles also depends not on the velocities of the layers, but only on the geometry of the triangle ABO. Snell’s law determines how the angle of a ray changes on crossing an interface, while geometry determines the change of angle between interfaces. Now,

pv

irv

ir constant sinsin

2

22

1

11 ===

The parameter p is known as the ray parameter and has the same value all along the path of any given ray provided that v, i, and r are measured at the same place.

Velocity-Depth StructureUsing an approximate velocity model, travel times can be calculated for the distance to actual seismic receivers and compared with observed times. The difference between the two is then minimized by adjusting the velocity-depth curve. This is repeated for millions of earthquakes and hundreds of seismometers all over the world. From this, a velocity depth model can be derived.

From M


8

Body WavesThere are two types of body wave (waves which travel through the earth).

P-waves – Travel through the earth in a series of dilations and compressions. Akin to sound through air.

S-waves -- Shear wave, do not travel through fluids, travel at about half the speed of P-waves.

From M


P, Primary (Body) Wave

• Deformation parallel to direction of propagation, e.g. like sound wave heard by human ear or pressure wave in a liquid. P waves can travel through solids, liquids or gases.

• Speed 1 km/s (in water) ~ 14 km/s (Lower part of mantle)

Axial Compression

Axial Expansion Direction of propagation

9

P WAVE

Direction of propagation

radial expansion

radial compression

S, Secondary (Body) Wave

• Deformation perpendicular to direction of propagation, shear wave that cannot travel through gases or liquids

• Speed 1 km/s (in unconsolidated sediments) ~ 8 km/s (Lower part of mantle)

10

S WAVE

Body WavesS-Waves cannot travel through water. The passage of an S-wave depends on the medium restoring its shape after initially being sheared. Water does not do this.

The S-wave velocity is always less than the P-wave velocity (vs = 0.55 vp).

S

SP

S

SP

vri

vii

2

2

1P

1

1

1

1P

1

sinv

sin :refraction

sinv

sin :reflection

=

=

As the velocity of the S-wave is different to the P-wave, the angle of reflection of a converted S-wave is not equal to the angle of incidence of the P-wave. Also, an S-wave is refracted at a different angle to a P-wave.

From M


11

Surface WavesWater waves are an example of a surface wave.They are slower than P- and S-waves and often have larger amplitudes.

Particle motion is a vertical elipse. It has both vertical and horizontal motion.

Particle motion is horizontal and transverse. It has only horizontal motion

Surface wave amplitude decreases rapidly with depth, similarly to water waves.

From M


Rayleigh WAVE

12

Love (Surface) Wave

• Deformation (in plane of surface) eg. side to side motion, not recorded on vertical seismometer.

• Speed 1 ~ 7 km/s

S wave front

Love Surface wave

Multiple reflections of (horizontal component)

SH wave trapped by surficial layer creates Love wave

Surface WavesFrom

Mussesttand Khan, 2000

13

Seismic Velocitiesmass eappropriat

force restoring wavesofvelocity =

The velocity depends on two main things – the restoring force (analagous to the strength of a spring), and the mass (analagous to the mass of the spring). As the restoring force increases, the velocity increases. However, as the mass increases, this will slow the spring, reducing the velocity. The mass in the case of a rock is its density (mass per unit volume).

S-waves involve a change in shape – this requires a shear force. The size of the force depends on the shear, or rigidity modulus, μ. A P-wave also involves a change in size, so the compressibility modulus κ is also involved.

ρμ

ρ

μκ

=

+=

s

p

v

v 34 Where ρ = denisty.

In a liquid, μ is zero, so vs is always zero.

From M


14

Seismic Velocities (P-wave)

15

Rock Velocities (m/sec)

Influences on Rock Velocities• In situ versus lab measurements• Frequency differences• Confining pressure• Microcracks• Fluids – dry, wet

16

Wave Propagation on Inhomogeneous Medium

Reflected P wave

Refracted P wave

Reflected SV wave

Refracted SV wave

Reflection & Refraction

• P and SV (vertical component) waves, reflects and refracts at boundary layer between two rock/soil layer: producing both SV and P waves

Incident P wave or SV wave

17

Reflection & Refraction

Incident SH waveReflected SH wave

Refracted SH wave

• SH (horizontal component) waves, reflects and refracts at boundary layer between two rock/soil layer but no P reflected orrefracted waves are produced.

Refraction through stratified layers near surface

• Refraction tends to cause P and S waves to become vertically orientated as they approach the surface.

Surface

P & S

P & S vertically orientated

18

Scattering of P and S waves

• Reflection and refraction, add complexity to seismograph recorded at the city.

City

Epicenter

19

Identification of seismic phases

2008, May 12, M7.9, Eastern Sichuan, China

Seismology of the EarthAs S-waves do not travel through liquid, they do not travel through the outer core. As μ is reduced to zero in a liquid, the P-wave velocity is reduced.

Mohorovicic discontuity: P-wave velocity jumps to more than 7.6 km/s. This defines the crust mantle boundary. The depth of this boundary varies from 5-6 km under the ocean floor to 70 km or more below major mountains. The average is approximately 40 km.

Low velocity zone: ~100 km depth. Not found everywhere.

400 km discontinuity: Velocity increase abruptly –olivine and pyroxene reorganize to more compact forms.

600 km discontinuity: Additional phase/composition change.

20

Ray Paths in the Earth

A ray is named according to the parts of the earth that it travels through – e.g a P-wave traveling successively through mantle, core, and mantle again is called PKP. A P-wave reflecting of the core is named PcP, etc. These are referred to as phases.

There are no main P-wave arrivals in the interval 98o

to 144o. This is the P-wave shadow zone. There are no S-wave arrivals beyond 98o.There are some weaker arrivals between 98o and 144o, because as well as ones reflected into it such as PP, the inner core reflects some rays into it (hence the discovery of the inner core).

From M

ussesttand Khan, 2000Phases

Receiver gather for station 36 and shot line 8+23 (Gulf of Corinth, Greece). Data are bandpass filtered between 3–8 Hz and reduced at 8 km/s. Horizontal axis is shot-receiver offset. Strong, late phase between 10.6–13.0 s is Ps, interpreted as a reflection from the subducting African slab. Weaker phase around 7 s is probably PmP. Inset shows location of receiver (black dot) and shot line (black line) in relation to other stations and shots.

21

Ray Paths in the Earth

A theoretical travel time plot for an earthquake. Earthquakes that arrive at a distance of greater than 18o are termed teleseismic. These are important as they not only sample deep parts of the earth, but they come back to the surface at a steep angle, spending as little time as possible in the highly variable crust.

From M

ussesttand Khan, 20002002 Denali Earthquake

From: http://w

ww

.citiesoflight.net/AlaskaQuake.htm

l

Note: Anim

ation only visible in powerpoint

24

Earthquakes

As both sides of the fault move, strain builds up across the fault – fences may be bent, etc. Once the strain becomes more than the fault can support the strain is released by elastic rebound. Energy is released as friction, block movement, and seismic waves.

From M


25

Where was that Earthquake?Unless we see rupture at the surface, which is rare, we do not know where the earthquake occurred. Using first arrivals is limited – we can tell which station the earthquake was closest to, but we do not know how long it took to get there.

It is as if we are trying to tell how far away a storm is but we are not seeing the lightning – we have only thunder to judge by.

From M


Where was that Earthquake?What if we use both the P- and S-wave (the thunder and the lightning)?

As S-waves travel more slowly than P-waves, the more distant the earthquake from the receiver, the greater the lag of the S- after the P-arrival. There are standard curves for this purpose: In this case a P-S arrival time difference of 6.5

minutes equates to a epicentral angle of 46o. We also know that the earthquake occurred about 8 minutes before the first P-arrival. If this procedure is repeated for multiple earthquakes we can triangulate the location

From M


26

How Deep was that Earthquake?The depth of the hypocenter below the epicenter can be found by measuring the difference in arrival of the direct P-wave and the wave that reflects from the surface, pP. As the depth increases the pP-P difference increases.

From M


Fault Plane SolutionsSimply put, a fault plane solution tells us the orientation and nature of the fault that caused an earthquake.

In the simple case above, imagine a peg in the ground that is struck by a hammer from the north. Immediately following the impact the ground directly to the north of the peg experiences compression, and that to the south experiences dilation. The magnitude of the compression and dilation decreases off axis (b). S-waves are also generated in the east and west directions.

From M


27

Fault Plane Solutions

In the above example there is a N-S trending right-lateral (dextral) strike slip fault surrounded by a circle of seismometers. When this fault moves, the seismometers will record a “first motion”. In the top left quadrant this is +ve, or up, in the bottom left quadrant this is –ve, or down. By mapping out the first motions on these seismometers, we can derive what sort of earthquake it is.

However, there is ambiguity, as a left lateral strike slip fault trending E-W would also fit these first motions. This is where we might also consider the local geology –are there any dominant trends? Is there a cluster of aftershocks that illuminates a particular plane?

From M


Beachballs

Traditionally an earthquake fault plane solution is displayed as a beach ball (b). Here we are looking down onto the lower hemisphere of a sphere (a). To create this beachball, earthquakes are plotted an an equal area Lambert projection net using the azimuth, take-off angle, and sense of the earthquake.

The above beachball defines a fault plane with a roughly NW-SE strike and a dip of either 25o ~NE or 65o ~SW.

From M


28

Making A Beachball

From M


First plot azimuth, take-off angle and sense of earthquake. Find a plane that splits two areas of compressional and dilational earthquakes. Plot the pole (P) to that plane, and then find another plane that also splits two area of compressional and dilational earthquakes, but also passes through the pole to the previous plane (ensuring that both planes are perpendicular).

Beachballs for Various Faults

From M


29

Earthquakes in PNG

Earthquake Intensity

From

Tar

buck

and

Lutg

ens,

200

5

The Mercalli Intensity Scale is a measure of what people reported. This is useful for areas where there were few seismometers, or in study of ancient earthquakes.

30

Earthquake MagnitudeIn 1935 Richter devised the Richter magnitude scheme for describing the size of earthqauke.

•Measured amplitude in microns of the largest oscillation of a particular type of seismometer 100 km from the source.

•The amplitudes have a very large range, so he took the logarithm (to base 10) to make the numbers more manageable. An increase of 1 in magnitude means the amplitude is 10 times greater (energy release is 30 times greater).

•Magnitude = log10(max amplitude of oscillation, in units of 10-6m).

•-ve values are possible (oscillations < one millionth of a meter). Many –vemagnitude earthquakes have been recorded at the HUGO seismic station half way between Hawai’i and the mainland.

•The scale was originally designed for shallow earthquakes near the receiver and a particular type of seismometer. It has been modified to deal with this.

•It underestimates the biggest earthquakes – many seismometers are not as sensitive to the lowest frequencies.

Earthquake Magnitude

From

Tar

buck

and

Lutg

ens,

200

5

31

Seismic MomentThough the Richter magnitude scale is the most commonly quoted, a later and better measure is the seismic moment, Mo. Just before a fault ruptures, the shear forces on either side of the fault exert a couple, whose size, or moment, equals the product of the shear forces and the perpendicular distance between them. The force is dependant on the strain, the area of rupture, A, and the rigidity modulus, μ. The strain depends on the fault offset and the width of the strained volume.

oMAd

dbF

==

==

=

μ

μ

couple ofmoment

,2b

strain and strain,*A F As

2*coupleofmoment

How do we determine rupture area?•Aftershocks

What is the maximum seismic moment of an earthquake limited by?

From M

ussesttand Khan, 2000Seismic Moment

From

Tar

buck

and

Lutg

ens,

200

5

Here we can see both the along strike extent of the subducting plate (shown by the earthquake distribution), and the down-dip distribution. The area of the plate rupturing in a given earthquake is a limiting factor in the moment magnitude of the earthquake.

32

Seismic Moment

Fortunately, only distinct parts of the subduction zone slip at one time, limiting the size of the earthquake.

From

Tar

buck

and

Lutg

ens,

200

5Risks and Mitigation

What are some of the earthquake risks?What are some of the ways that we can minimize damage?

Build Sensibly

Tsunamis and tsunami warning systems

Using automated seismic triggers to slow trains, etc. (Bullet Train in Japan).

From

Tar

buck

and

Lutg

ens,

200

5

33

References Used1. Basic seismic theory:

• Kearey, P., M. Brooks, and I. Hill, An Introduction to Geophysical Exploration, 3rd edition., pages 21-30, 2002.

2. Basic theory, seismology, and earthquakes:• Mussett, A.E. and M.A. Khan, Looking into the earth: An introduction to

geological geophysics, pages 24-64, 20003. Really basic theory:

• Tarbuck, E.J. and F.K. Lutgens, Earth: An introduction to physical geology, chapter 11, 2005

4. Plus additional html references as listed in this presentation.

2002 denali earthquake - sgcobinsus.files.wordpress.com · 02/08/2017 · 2 seismology – what is...

Documents