FOCAL
DEPTHS
AND MECHANISMS
OF MID-OCEAN RIDGE EARTHQUAKES
FROM BODY WAVEFORM INVERSION
by
Paul
B.Sc.,
M.S.,
Yi-Fa Huang
Calgary (1974)
Alaska (1979)
University
University
SUBMITTED TO THE DEPARTMEN
EARTH, ATMOSPHERIC
ANn PLANETAR
IN PARTI AL FULFILLMENT
OF THE REQUIRE MENTS
FOR TH E DEGREE OF
DOCTOR
OF
SCIENCES
OF PHILOSOPHY
at
the
MASSACHUSETTS INSTITUTE OF TECHNOLOGY
December 1985
© Massachusetts
Institute
of Technology
1985
7Signature of Author
Department of Earth,
anj PTanetary
Atmosp Ker-ic,
?F-Scien c
rm_
Certified
b
Tean C. Solomon
Thesis Supervisor
/
Accepted
Sha irman,
APT I IRRARIES
MAIT !IRRARIES
Deoartment Committee on
Graduate Students
FOCALS DEPTHS AND MECHANISMS OF MID-OCEAN RIDGE EARTHQUAKES
FROM BODY WAVEFORM INVERSION
by
PAUL YI-FA HUANG
Submitted to the Department of
Atmospheric, and Planetary Sci
on December 30, 1985 in Partial Fulfil
Requirements for the Degree of Doctor
Earth,
ences
Iment of the
of Philosophy
ABSTRACT
Accurate focal depths of mid-ocean ridge earthquakes are
critical to tectonic interpretations of ridge-crest processes,
notably the thickness of the brittle layer beneath the ridge
axes and, implicitly, the depth of hydrothermal circulation.
We determine in this thesis the focal depths and source
characteristics of 29 ridge-axis earthquakes along the MidAtlantic Ridge, the Azores spreading center, the northern Red
Sea, the Gulf of Aden, and the Carlsberg and Central Indian
Ridges. The source parameters are obtained by a formal
inversion of long-period P and SH waveforms [Nabelek, 1984].
Because of the special importance of focal depth, we devote
particular attention to the resolution of centroid depth and to
an assessment of the relationship between the formal error
given by the inversion method and the actual uncertainty. We
also stress by example the advantage of including SH waves as
We examine as well the
well as P waveforms in the inversion.
effects of uncertainties in the source region velocity
structure and of incomplete azimuthal coverage on the inversion
results. The waveform inversion solutions are shown to be
generally insensitive to details of source velocity structure;
as long as azimuthal coverage spans more than. 2 quadrants of
the focal sphere for both P and SH waves, the solution obtained
by waveform inversion is robust with respect to inclusion or
deletion of selected subsets of the data.
The ridge-ax is earthquakes of this study are all
characterized by n ormal faulting on planes dipping at about 4 50
and generally str iking parallel to the trend of local ridge
Seismi c moments range from 2 to 106 x 1024 dyn-cm,
segments.
The P and
and source time f u nctions are all of simple form.
S waveforms for a 1 1 earthquakes can be well matched using
111i
The
conventional values of t* (1 and 4 sec, respectively).
P waves from these earthquakes show strong water-column
reverberations, suggesting that fault rupture generally
extended to the seafloor.
The predominant period of these
reverberations constrains the water depth in the epicentral
On the basis of es timated water depth and the
region.
epicentral location, all of these earthquakes can be shown to
have occurred beneath the i nner floor of the median valley.
We
cannot exclude the possibil ity that some of these events
occurred near the inner wal 1 of the rift
valley (e.g., the
August 15, 1966 and the May 31, 1971 earthquakes) or that they
represent slip on major inw ard-dipping faults that contribute
to the relief of the inner valley walls.
It may be inferred
with reasonable certainty, however, that none of these
earthquakes occurred beneat h the rift
mountains.
The
centroid
depths of the
earthquakes
in thi s study
prov ide important information on the vertical exte nt of brittle
fail ure along slow spreading ridges. A few of the earthquakes
incl uded here, however, may not be representative. Two
eart hquakes in the Azores region may have occurred in response
to the slow rifting of oceanic lithosphere rather than along a
true oceanic spreading center.
Two earthquakes in the northern
Red Sea are likely to have occurred in thinned continental
crust.
Excluding thes e earthquakes, the centroid depths of the
ridge-axis events with adequate azim uthal coverage are all
very
shallow, ranging from 1 to 4.2 km.
If centroid depth marks the
mean depth of fault sl ip, then earth quake faulting extended to
2-8 km beneath the sea floor during these earthquakes.
The two
Azores earthquakes hav e ce ntroi d depths of about 5 km.
The two
northern Red Sea earthquak es have centroid depths of about
6 km, consis tent with the view th at seismic faulting
accompanying continental rifting
extends to greater depths than
does seismic activity alon g ocean ic
ridge axes.
Two
microearthqu ake experiment s condu cted in regions of the
Mid-Atlantic Ridge near ep i center s of one or more of the la rge
earthquakes in this study s how th at microearthquakes occurr ed
deeper than the probable s1 ip zon e of the larger earthquake S.
However, the uncertai nti es both i n centroid depths and in a nd
microearthqu ake focal dept h s also permit
a coincidence in depth
of the two measures
of
activity.
The combined set of centroid depths for well-constrained
source mechanisms of Atlantic and Indian Ocean ridge
earthquakes show that the greatest depth of earthquake faulting
tends to shoal with increasing spreading rate.
This provides
direct evidence for a general decrease with spreading rate in
the maximum thickness of the zone of brittle behavior for
slow-spreading ridges.
The maximum depth of earthquake
faulting probably outlines the depth of the brittle-ductile
transition
beneath the ridge axis.
It
is likely
that the
1V
lithosphere is colder and mechanically stronger beneath the
slowest-spreading ridge segments because there is more time for
the lithosphere to cool between major episodes of crustal magma
injection and volcanism.
Additional evidence for this
interpretation is the fact that the largest earthquakes
(Mo > 9 x 1024 dyn-cm) occur on ridge segments with
half-spreading rates less than 14 mm/yr.
Thesis
Supervisor:
Title:
Sean
C.
Solomon
Professor
of
Geophysics
ACKNOWLEDGEMENTS
I wish to express my deepest gratitude to my advisor
Professor Sean Solomon for suggesting this thesis topic, and
for thoughtful discussions, guidance, encouragement,
inspiration and
financial
support.
Without
his
illuminating
comments and cr itical editing, this thesis would be much less
readable.
I am also indebted to Eric Bergman
body
modelling.
waveform
numerical
thesis
aspe
designed
have
generously
of his
insights.
His
enhanced this
I
Aki,
I
and
want
to
h is
Lynn
going
late
got
John
Nabele
hod used in this thesis
9 who
gave most
expertis e and provided invaluable
c ritiques,
a nd
consultations greatly
his
discussions wit h Professor Kei
constructive crit icism.
I
am also
stimulation and inspiration Provided by
express my gratitude toward a
those
who provided companionship.
particular,
to Kiyoshi
night
f riendship
Hall
in the
Madden.
thanks
exciting
for
him for
associates
special
and
possible
h elpf u 1
some
for t he
Professor Ted
me enormously
me
undertak ing.
thank
grateful
time
comment S,
also ha d
and
been
i nversion met
the
helped
"training"
cts of my re search, and without his help this
n ot
would
He
for
provided
rough.
friends
a
Yomogida for his f riendship and many
discussions.
I also thank
especially whe n I was sick
Stuart Stephens
in the hospital.
encouragement a nd warm friendship when the
I gained much fr om discussions with
Roger
Buck,
Tianqing Cao,
and
Justin Revenaugh
Dan
Scott
Davis,
Steve
Rob
Hickman,
Doug
Phillips,
Toomey.
Jack
McCaffrey,
T hanks
Jemse
ck,
Mark
Murray,
also to Karl
Bob Nowack,
Coyner
Jea n
Baranowski,
Tom Wissler,
Shedlock,
lene Lyon-Caen, Tien When Lo, Sarah Kruse, Jeanne
Sauber,
He
John
her
also to Jean
thanks
with
to
thanks
Feldstein
Gre g Beroza.
1 ogistics
and Janet
and
their
much
support
from
A bove
all
others,
I
wi f e,
self-sacrifi ce.
re
Theresa,
To
her
I owe
My grati tude
figures.
A speci al
Sahlstrom
friend
for
thei r help
hip
my family fo r wh ich I am
wi sh
for
smile.
wa rm
had
to my
This
and
f riendship and constan t
Sh aron
also
thankful.
Grimm,
Joao Rosa, Jose Rosa, Kaye
ni e Kohl who drafted many
necessa ry
I
Okubo,
are due to Jan Nattier-Ba rbaro for her expe rt
Thanks
typing and
Bob
Gof f,
Paul
to
express my
her endless
special
encouragement
so much.
sear ch was suppor ted by the National Science
Foundation under
grants
EAR-8115908
and
EAR-8416192.
and
vi1
TABLE OF CONTENTS
Page
ii
Abstract
Acknowledgements
v
CHAPTER 1.
1
INTRODUCTION
Volcanism and Tectonics of the Axial
Hydrothermal
2
Zone
5
Circulation
Ridge Axis Seismicity
6
Thesis Outline
9
INVERSION PROCEDURE AND RESOLUTION
WAVEFORM
CHAPTER 2.
13
OF SOURCE PARAMETERS
Introduction
13
Source Paramaterization
13
Inversion
General
17
Details
Importance of
Including SH Waves
21
Resolution
22
Estimation of the
The
15
Formalism
Procedural
Depth
15
Procedure
Effect
of the Assumed
The Effect of
Comparison
Uncertainty
With
in
Centroid
Depth
Velocity Structure
Incomplete Azimuthal
Surface Wave Methods
Coverage
24
29
36
39
Tables
41
Figure Captions
43
Figures
48
viii
CHAPTER
3.
RIDGE-AXIS EARTHQUAKES
IN THE NORTH ATLANTIC
OCEAN
64
Introduction
64
The Mid-Ocean Ridge System in the North Atlantic
64
Earthquake Data
Set
69
Inversion Results
70
Waveform
Discussion
90
Tables
94
Figure Captions
101
Figures
106
CHAPTER 4.
RIDGE-AXIS EARTHQUAKES
CENTRAL
IN THE NORTHWEST AND
INDIAN OCEAN
131
Introduction
131
Spreading Centers of the Northwest
Indian Ocean
Region
132
Earthquake Data Set
135
Waveform Inversion Results
136
Discussion
153
Tables
158
Figure
Captions
165
Figures
CHAPTER
5.
169
SOME GENERALIZATIONS ON THE CHARACTERISTICS
OF RIDGE-AXIS EARTHQUAKES
Introduction
Uncertainty in Centroid Depth:
191
191
Another Look
191
Comparison of
Source
Mechanism with
Harvard CMT
Solutions
193
Modeling Water
Column
Reverberations
193
Depth Range of Brittle Behavior
194
Choice of Fault Planes
199
Stress Drop
201
Seismic Moment vs.
Dependence
Centroid Depth
202
of Source Parameters on Spreading Rate
Near-Ridge Earthquakes
202
204
Tables
206
Figure Captions
209
Figures
212
CHAPTER 6.
THESIS SUMMARY AND SUGGESTIONS FOR FURTHER
WORK
227
REFERENCES
230
APPENDIX A
TESTING HYPOTHESES
250
APPENDIX B
RESOLUTION OF CENTROID DEPTHS
256
APPENDIX
COMPARISON OF INVERSION SOLUTIONS WI TH
C
ALTERNATIVE
SOURCE
VELOCITY
STRUCTURES
289
1.
CHAPTER
the
major
exerts
upper mantle
crust
Earth's
most
active
and
contributed
significantly
the
1950's,
Bullard
the
rift
zones
of
studies
Herzen and
Based
on
[1963]
the
that
the
of deep-seated
studies
of magnetic
suggested that
imprinting
genesis
of
magnetization of oceanic
determinations have
[Barazangi
and
Dorman,
have
1969]
of
plate tectonics.
discovered
This
mantle
Von
1959;
1964]
in
is
[Hess,
1962].
the crust
reversals
and
in
related
Matthews
is
recorded
by
the remanent
focal
the boundaries
Von
support
ridges
Vine and
Epicenter
flow
In
subsequent
Herzen,
material
and their
heat
and Maxwell,
lineations,
field
high
and
of mid-ocean
revealed both
are
ridges
ridges
Herzen
crust.
processes
of mid-ocean
the spreading of
geomagnetic
the
from
ridges.
[Von
is
proper
that
the
on the
remote
Vacquier and
1963;
Uyeda,
ascent
flux
which
and
the theory
[1954]
of mid-ocean
ridge heat
the hypothesis
al.
et
in
influence
strong
observation
studies
to
zones
as
the fact
Despite
stations,
seismographical
the Earth's
system, with tectonic
ridge
our eyes.
geophysical
of
ridge axis
The
lithosphere.
inaccessible to direct
largely
regarded
are
an exceptionally
of the
part
evolving before
to the
they
boundaries;
plate
marine
up over half
make
ridges
Mid-ocean
research.
in
important discovery
most
ridge system was
of a worldwide mid-ocean
The discovery
probably
INTRODUCTION
mechanism
of plates
relative motions
[Sykes,
1967].
The mid-ocean
in structure,
ridge
tectonics,
system displays
volcanism,
and
significant
hydrothermal
variations
circulation [Macdonald,
Oceans
ridges
the
1982].
In the Indian and Atlantic
are relati vely narrow
prominent deep
val ley along the crest.
spreading
Pacific Rise is
has
East
no prominent
including
Biehler,
"hydraulic head
1970;
[Deffeyes,
Sleep
1970],
Francheteau,
correlation
None
valley along
and
1978],
of
and
rift
date,
largely
deeper crust
In
istics
and
of
large
study
been
has
lack
the maximum
depth
interpreted
as
1979],
clearly
knowledge
we will
believed to
which
the depth
of
be
the
models,
and
imbalance"
spreading
be
rate.
superior to
of
rheology
the
ridge axis.
an
depths.
of
character-
inversion of
the main
result
earthquake
the
with
from
reliable centroid
and
[Tapponnier and
the source
One of
fast-
explain the apparent
of the
earthquakes
a
rugged
Sleep
shown to
determine
the
"material
beneath the
SH waveforms.
at
1969;
proposed to
been
of
to obtain
are
[Sleep,
less
Several
characteristics
ridge-axis
earthquakes
crest.
and
"steady state necking"
of the mantle
P and
is
loss"
valley
this thesis,
long-period
this
for
its
contrast,
wider
Rosendahl,
have
of these models
much
with
d rugged,
purposes of
Since
brittle failure,
faulting
occurs
can
be
brittle-ductile transition
zone.
VOLCANISM AND TECTONICS
The
axial
region
a central
zone of
boundary"
zone
is
the
OF
[Macdonald,
active volcanic
AXIAL ZONE
of a mid-ocean
"crustal
characterized by
THE
recent
portion
ridge can
accretion
1982].
and
T he
be divided
a flanking
crustal
of
this
region
is
"plate
accretion
and ongoi ng magmatism
and
into
zone
volcanism;
sometimes called
the
"neovolcanic zone."
nonvolcanic zone
volcano chains
1975].
al.,
At
can consist
[Needham and
In
contrast,
essentially continuous
transform faults
and
segments
et
central
[Searle
Macdonald,
al.,
19801.
volcanic
zone.
for
of
range
al.,
1977].
is
view,
and
1980a]
ridges
and
This
varies
and
The plate
zones
of
along
all
ridges
display
and
Fox,
km wide
Moore,
with
time
(e.g.,
FAMOUS
well-developed wall
fissuring and
ridge
axes.
At
intense fissuring occurs
both
in
terraces
zone
faulting.
can
1-2
of
view
he
ridges
accretion zone
median
In
support
ridges
can
floor and narrow
a narrow
(e.g.,
AMAR).
by active
have
ridges,
km wide
of the
1975;
be subdivided
Fissures
types
bands
al.,
to one with
characterized
This
deformation.
Spiess et
1977].
inner
Luyendyk
estimated to
slow-spreading
area)
is
1977;
individual
[Macdonald,
a wide
all
slow-spreading
an
The
anomaly polarity
been
of crustal
valley of
boundary zone
has
km for
structure of
at
1979;
[Klitgord et
1-8
by
1983].
instantaneous
zone
et
spreading center
1-2
an
variation
the median
flanking terraces
faulting and
Macdonald
Ballard,
only
from a configuration with
inner floor
1974;
Bryan and
accretion
that the
segment
this
Andel
discontinuous
From the width of the magnetic
suggests
valley
van
1976;
the
interrupted only
Macdonald
is typically
fast-spreading
[Macdonald,
zones
overlapping
1981;
[Normark,
the crustal
Macdonald et
width
fast-spreading
However, this
transition,
1-3 km
al.,
ridges,
isolated and
Francheteau,
offset
1977;
of
neovolcanic
neovolcanic zone
spreading rates
and
slow-spreading
been
into
observed
the most
flanking the
central
volcanic zone.
across,
extending
Macdonald,
parallel
formed
1977;
to
by
and
At
10 m to 2 km
Ballard
the strike of
the deeper
spreading
slow
ridges,
fissures
by
At
continuous
Outside the fissure
of the
crust
spreading
indicates
200 m or
greater, and
600 m or
higher,
Luyendyk,
1977]1.
evident,
scarps
Using
50 m or
precise bathymetric
axis
Using
of
that
in
the
a similar
faulting
the active
rift
intermediate- and
data,
4 to
valleys
10
axially
slow
dipping
and
of
scarps
and
valleys are
rift
fault
ludie,
1974].
Atwater F19781
up
estimated
ridges.
disruption
throws
to
the Mid-Atlantic
km from the
fast-spreading
no
[Klitgord and
[1980]
1982].
[MacDonald
ridges,
are
but
creates
faults
faulting continues
Shih
At
vertical
Macdonald
mountains of
analysis,
extends to
less
along the
observed
faulting.
creates
1977].
[Lister,
[Macdonald,
have
spreading
faulting
normal
with throws
estimated
the
but
fast
that
likely
vertical
these
of
deep
in
is
to fast-spreading
significant
a series
resulting
At
volcanoes
faults
individual
ridges,
are
due to thermal
It
are also
normal
are
they
occur along the axis
zone,
active
that
crust
intermediate-
central
and
seawater to penetrate
hot young
also
probably
1977].
for cold
of
They
than cracking
fissures
ridges,
spreading axis.
implying
MacDonald,
1-3 m
[Luyendyk
19771.
Andel,
rather
levels
central
obscured
van
provide access
are typically
along strike
the ridge,
and
[Luyendyk
fissures
cool
and
tension failure
contraction
these
These fissures
30
km
from
Ridge.
that
active
spreading axis
of
HYDROTHERMAL
Both
CIRCULATION
high
and
close to the axes
low
heat
of mid-ocean
confused picture of the
and
Uyeda,
Lee,
1969;
1965;
Von
Langseth
addition, all
and Von
of the
heat
al.,
importance
at
isotopes,
the
direct
[Macdonald
is
field
Pacific Rise
first
et
al.,
steady-state.
Ridge
at
1980b].
large,
26 0 N,
smokers,
The
at
1979
and
In
1969;
flow
Sleep,
in young
of
Hydrothermal
of
in
at
a mechanism
bottom water
the
Group,
helium
et
al.,
axis
19801
1977].
of the
allowed
hydrothermal
heat
flux
flux
from
the vents
heat
that
hydrothermal
these
venting
were discovered
providing direct
activity
as
salinity
total
suggesting
Herzen
led to a realization
19721.
Project
of axial
[Lee
19701.
the heat
spreading center [Weiss
210 N [RISE
Recently,
black
hydrothermal
at
Von
Sclater,
measurement
in
ridges
Lister,
circulation
and
discovered
estimate
overwhelmingly
including
temperature
measured
near-ridge thermal
[Lister,
verified by
Galapagos
The hydrothermal
the
hydrothermal
first
potential
samples at
of the
These observations
of
1965;
1970;
overestimate
mid-ocean ridges
was
circulation
Langseth,
McKenzie and
19711
lithosphere.
loss
East
1967;
been
pattern across
Herzen,
conductive models
Sclater et
oceanic
and
have
initially producing a
ridges,
heat flow
Herzen
structure [McKenzie,
1969;
flow values
evidence
vents
not
phenomena,
along the
of
are
Mid-Atlantic
high-temperature
slow-spreading oceanic
ridges
[Rona,
19851.
It is difficult to estimate the depth of penetration
hydrothermal
circulation,
but
of
earthquake depths should provide
some bound.
In
the
East
Pacific Rise
at
area
hydrothermal
21 0 N, microea rthquakes are between 1 and 3 km deep [Macdonald
et
al.,
1980c
;
I
Rie
-
rina al
.
.
I
a
nIlf
I1 L
intraplate ea rthquakes [Chen and
19831
have
regions at
sh own
9-
I
I
.
Mol n ar,
Focal
depth
1983;
studies
Wiens
and
abo ut 600
than
0
-800 0 C.
At
differential
stress hecause of du ctile flow.
represent bri
ttle fail ure, the ma ximum dept h at which
occur
a rid
thickness of
crust at a temperature les
circulation.
result
Oxygen
provide
addit
oceanic
crust may
Taylor,
1981].
AXIS
onal
manifestation
coo
i sotope studies
evidence that
reach
depths
lin
th an
600
0
C-800 0 C,
by hydrothermal
in the Samail ophiolite
hydrothermal
greater than
along mid-ocean
of the
spreading East
faults.
of on-axis
of a significant
5
circulation in
km [Gregory and
SEISMICITY
Earthquakes
distances;
earthquakes
The occur rence of earthquakes
ge axis i ndicates the p re s nce
presumably th e
Since
can provide bou nds on the depth of the
le transi tion zone.
brittle-ducti
RIDGE
higher
the lith osphere can not sustai n significant
temperatures,
beneath
Stein,
that seismic fai lure is generally confined to
te mperature s less
earthquakes
of
process of
Pacific
recorded
Slow-spreading
by earthquakes
Modern
study
of
seafloor
Rise appears
earthquakes
marked
ridges
ridges,
with
are
in
are a primary
spreading.
aseismic
at
The
fast-
teleseismic
confined to transform
contrast,
are commonly
body-wave magnitudes
ridge-crest earthquakes
dates
as
large
from the
as
6.
demonstration
by
Sykes
characterized
by
normal
are
approximatly
direction of
solutions,
spreading.
characteristics of
particularly
and
events
oriented
have since
focal
having
of
earthquakes.
ridge
crests
and
its
an extremely
implications
structure
and
for
injection
and
cooling by
Lister,
activity
it
only
important
as
behavior
thermal
shallow magma
circulation
[e.g.,
was that of Tsai
generated
fo r
when
19681,
well-
1 number of large ridge-crest earthquakes.
sp ect
65
Isacks et al.,
An
to determine the focal depth of ridge-crest
effort
30 to
ridge-crest seismic
foc al depths have been presented in the literature
amplitude
that,
processes
is generally agreed that
a s mal
earthquakes
of
is
of brittle
for local
hydrothermal
qui te shallow [e.g.,
is
constrained
early
controlling
to the
1977].
While
for
such
that
the source
This
quantity for understanding the depth extent
at
T axes
fault-plane
been made
depth,
are
parallel
Aside from simple
ridge-crest
true of
such
faulting mechanisms
horizontal
few studies
that
[19671
ra from several events to synthetic spectra
v arious source depths and obtained focal
phas
depths
Weidner and Aki [1973] demonstrated, however,
km.
are consider ed,
[1969], who fit Rayleigh-wave
e as well as amplitude spectra of Rayleigh waves
the focal
depths of ridge-crest earthquakes are
required to
be very shallow; they obtained depths of 3 + 2 km
beneath the
sea floor for two normal-faulting events on the axis
of the
Mid-Atlantic
reported that
the
Ridge.
Duschenes
short-period
and
P waveforms
Solomon
[19771
from these two
events contain apparent depth phas es consisten t
shallow
with
such
Conventiona 1 fault-plan e solutions
focal depths.
derived from P-wave first motions for
often show non-orthogonal
nodal
ridge-cr est earthquakes
pl anes
Thatcher and Brune, 1971; Solomon
[Sykes,
and Julian,
1967, 1970a;
1974],
in
apparent disagreement with the exp ected double-couple
model .
Hart
[19781
suggested that
this
non-orthogonality
consetquence of the interference between direct
and
This
sugge'stion was confirmed by Trehu et al.
who
for the
source
[19811],
moment
is
a
surface-
refle cted phases for extremely shallow sources.
Rayle igh wave spectra
source
tensor
inverted
and
synthiesized long-period P waveforms for two earthquakes on the
Reykj anes Ridge.
spect ra
and
the
They
showed that
P waveforms
sourc e mechanisms
and
very
were
both
the surface wave
compatible
shallow
focal
with
depths.
model ed the short-period P waves from a small
double-couple
Pearce
[1981
(mb = 4.7)
earth quake on the axis of a spreading center segment in the
Gulf
of
Aden
Beyond
remaining
and
inferred
these
a focal
few studies
publish ed
i
of
depth
large
of
5 + 2 km.
earthquakes,
the
nformation on the depth of seismogenic
ridge
axes comes from mi croearth quake surveys.
While a number
of mic
roearthquak e studi es have been carried out
along
of the
faulting
segments
systems
(OBS's),
an
along
using eit her
the
[Duschenes
et
al.,
Ridg e and ot her ridge
sonobuoys or ocean-bo ttom seismometers
of these exp eriments were conducted with
majo rity
insufficient
Mid-Atlantic
numbe
r of stations to reso lve focal depths
1983].
Two Mid-Atlantic Ridge microearth-
]
quake
experiments with
networks of more
noteworthy exceptions.
of St.
[Francis
et
valley at
faults
al.,
(near
et
[Toomey
Zone
19781,
a site
al.,
near the
One was
Fracture
Paul's
and the
second
230 N) distant
Both
19851.
km below
the
sea floor,
large
earthquakes well
These
data
provide an opportunity
the large
depths
of microearthquakes
was within
ridge
the median
displayed
regions
are
also the
to depths
sites
in the
of
of
teleseismic distances.
to
compare
the centroid
with the distribution of
earthquakes
of
valley of the
the median valley
recorded at
are
from major transform
and both
depths
3 instruments
eastern intersection
and the median
microearthquake activity within
7-8
than
same
focal
region.
THESIS OUTLINE
Accurate
critical
foca
to tecton c
depths would
place
the
ridge axes
thermal
circulation.
reliable
earthquakes.
cannot
phases
arrival
pP and
recordings.
possible
sP
In
impli
main
depths
Generally the
be determined
without
and,
The
focal
the thi
oh
times
at
focal
f
clearly
contrast,
very
go
using waveform modeling
Such
citly,
limits
on
the brittle
are
focal
layer
o n the depth of hydro-
jective of this study is to
depths of shallow earthquakes
rom trav el times,
s tati
nearby
are not
teleseismic distances.
ckness of
earthquakes
a large number of ridge-crest
for
accurately
ridge
bec ause
interpretations,
bounds
beneath
obtain
of mid-ocean
depths
effo
at
least
ons, because the depth
observed on teleseismic
od depth determination is
for even ts well-recorded at
rts are worthwhile,
because,
although OBS's will give increasing numbers of precise depths
the future, microearthquakes
may not
be
similar to larger
rthquakes.
We
det ermine
characteris tics
an d
Atlantic
by
a formal
this
of 29
Indian
inv ersion
oceanic
int
Bergman
and
Oceans.
of
of
focal
of
resolut ion
we
the formal
individual
of
valley tectonics
alon g
procedure.
We
also
velocity structure
We find
solutions
are general ly
velocity structure.
spans
SH
more
waves,
In
than
the
2 qua drants
3, we
obtained
al.
We
applied
[1984]
to
and
the special
assessment
the
generalize
to
to
of the
inversion
the
a discussion
source parameter s for median
technique of
resolution
th e effects
th at
as
of
result
present
to
long as
the
is
of the
Nabelek
in version
of uncertai nties
coverage
on
in
the
the waveform inver sion
insens itive
Also,
inversi on
Chapter
we
an d complet e azimuthal
results.
as
given by
the i nversion
e xamine
inversion
are
After a presentation of
derived
the depth
and discusses
North
slow-sp reading ridge systems
Chapter 2 descri bes
[1984]
et
and to an
inversions,
the
the
particular atten.tion
error
uncertainty.
implications o f
and source
SH waveforms.
Because of
devote
centroid depth
results
in
[1984],
by Bergman
depth,
actual
P and
Nabelek
1985a].
method and
depths
earthquakes
[1984,
relationshi p between
focal
long-period
procedure of
Solomon
the
The source parameters
raplate earthquakes
importance
of the
thesis
ridge-axis
inversion
use the
the
in
details
of s ource
the azimuthal
focal
sphere for
coverage
both
P and
stable.
source
studies
of
14 Mid-Atlantic
Ridge earthquakes and 2 normal-faulting earthquakes on the
slowly spreadi ng portion of the Azores-Gibraltar plate
boundary.
All
of
these
earthquakes
have mechanisms
characterized principally by normal
with dip angle s near 450.
wave-trains
floor
inner
all
Mid-At lantic
1 to
of the
Ridge
3 km.
centroid d epths
In
Ch apter
we
earthquakes
events
he Red
and Centra
maps
beneath
inner
th
earthquakes
Sea
crust.
from
are
likely
Red
ranging
have
Indian
cean
and
of
region,
Ag
to
have
including
ain, a comparison of water depth
tha
from
P waveforms
the median valley.
however
are a.
The
urred
th ree
with
that
from
t these earthquakes occurred
is
in
events,
two
One
of
these
probably an
th e se i smicity does
this
13
Gulf of Aden and on the Carlsberg
the
19 81),
at
of
not
clearly
earthquakes
thinned
define
in
the
continental
the centroid
depths
range
even
(April
22,
1969)
near a ridge-transfor m inters ecti
where
he
itho-
5 km.
The
sphere is probably
the
in the
the se
Excluding
1 to
occurred
boundary
of
depths
large
tudies
thou gh
beneath
region
source
7,
P
n the Azo res
present
floor
t he
5 km.
i ndicates
(October
off-ridge event,
Red
very sha llow
picente r determ ined
bathymetri
from the
occurred
earthquakes
Ridges.
Indian
above the
the plate
Sea
in
The centroid
ar
plates
estimated
median valley.
of about
4,
fault
reverberations
The two earthquakes
ridge-axis
in
on
in dicate that the earthquakes all
the
from
The water depths
of water-column
predominant pe riod
faulting
Sea
deepe st
of
colder than
these
normal.
have greater centroid
The two earthquakes
depths, providing additional
evidence that seismogenic deformation within rifting
continental
crust
a greater depth
extends to
along oceanic
than
spreading centers.
In
the
Chapter
ridge-axis
such
earthquakes of
interest
particular
discuss
5, we
earthquakes may
is
the
be
We
spreading
In
also examine
characteristics
this
A question
study.
confidence with which
used to
brittle-ductile transition
ridges.
some general
zone
infer the
the
of focal
of
depths
of the
extent
spreading
beneath the slowly
the issue
of
depth
versus
rate.
Chapter
summarized,
and
6, the major
suggestions
conclusions
for
future
of this
research
thesis
are
are
offered.
of
CHAPTER
2.
INVERSION PROCEDURE AND
WAVEFORM
RESOLUTION OF
PARAMETERS
SOURCE
INTRODUCTION
source that
forms
we describe
the
advantage
length
at
the
obtainable with
resolution
the
inversion
structure and
of
less
of
of
We
depth
centroid
Finally we show the
in
uncertainties
the
azimuthal
optimal
than
Next,
P waveforms.
as
procedure.
results
inversion
on the
velocity
as well
earthquake
by example the
stressing
inversion procedure,
centroid
inversion.
body waveform
for
including SH waves
of
then discuss
effects
basis
the
the general
to describe
used
is
source which
point
of the
the concept
introduces
first
This chapter
coverage.
PARAMETERIZATION
SOURCE
The formulation
primarily
from
distribution
Nabelek
of moment
uk(x,t)
where
uk(x,t)
point
x and
at
the
Fi
the
is
time
t;
tensor
= -mij,j.
integral
which
The
is
source
density
=
V
are
E;
tensor
,t)*gk
in
Green's
and mij
are
are
related
to
* denotes
volume
displacement
The
displacement
symbol
over the
mij(
below
parameterization
[1984].
gki
source point
density
of
V of
can be
i,j(x,t;
the
ijxt;
k
(1)
direction at
components
the body
,j
mij.
the
= 3/3xj
taken
of the moment
force density
convolution,
a
as
,0)dV
functions;
nonzero
due to
field
written
drawn
is
and the
Equation
(1)
by
expresses
the theoret ical
the
unknowns mij(
all
time
and
,t
)
.
have only
parameterization
When the
the
with
series
about
data
source
wav elength
expand
a point
set,
is
is
the
mij
some
sort
uk
and
encompassing
define
[Ward,
1983].
of
necessary.
period
are small
of
the
compared
the motion,
Green's function
in
(Co,To)
set
required to
and duration
and
the data
large data
however,
dimension
we can
between
density distribution
of t he
dominant
infinitely
a finite
source
respectively,
An
po sitions,
space
completely the moment
Since we
relationship
in
source
region:
k i;jMZ
ijX
a Taylor
+
- gki;jM1 ij
= gki;jMoij
uk(x,t)
+ 7 gk i ;jM2 ij
-
+
1 g
ki;j
M ij
+ 7
ki;jxmM 0 ijem + ...
(2)
where
TO0
...
n =
(T-To ) nM i ...
j
f
0
=
f
9ki ;j =
F
n
...
where
are
the
dot
the
and
source
=
9ki
cosines
region.
The
vanish;
time.
S waves,
(5)
with
and
six zero-order
tensor.
o and T o
Dziewonski
et
(4)
9ki,j
respect
of the departing
P and
the moment
and M'ij
centroid
).0...
(n-n 0 n)mij(E,t)dV
Cn
downgoing
components of
Mijt
(0 -
denotes differentiation
the direction
upgoing and
in
n
(3)
n (r)d r
We can
are
al.
rays,
Cn
choose
then the
[19811
n is
an
is the wave
terms,
yjn
to time,
MOij,
centroid
for
velocity
are
Co and
combine
index
To
the
so that
location
MOij,
o
and
To
to form
their
higher-order
centroid
terms
M2ij
duration,
the
source
about
centroid
[1983],
about
difficult,
so
Nabelek
(personal
and therefore
the
it
in
we
The
time
location
is
nor
unknowns
degree zero,
INVERSION
General
It
should be
in
the
its
However,
there
centroid depth
of the
Nabelek
[1984],
called this
noted
is
that
we
group
because the
origin
With
time
the
are the
centroid depth,
of the
centroid
source
above parameterization,
properties
spatial
and the
of the
moment
source.
tensor
source time
of
function.
PROCEDURE
Formalism
We set
up the inverse problem
d
where d contains
seismogram
is
analysis
and
and
by
neither the absolute
the average
our problem
Ward
dropped.
Following
together
of the
argued that
duration
terms
directly.
nonetheless study
has
of the
terms
include the effect
ignored,
The
out by
the CMT
usually
1985)
to
pointed
than
are
inversion.
function.
be determined
can
the
higher-order
absolute travel
can
terms
important
As
the extent
second-order
making
communication,
is
source time
epicentral
thus
and
sm aller
between the source
duration
grouped
the
10,
velo city,
are
solution.
(CMT)
Moijkm give estimates
positio n.
higher-order
some trade-off
source
rupture
second-order terms
a factor of
tensor
and
lijZ
source
the
moment
for
based
on
S
=
the
=
in the
m(P)
(6)
the observations and
model
parameters
minimization
form
P.
m(P)
The
is a synthetic
least
squares problem
of
[d-m(P)TCd '[d-m(P)]
O
+
+ [Po.p]TC
O-l[Po-P]
(7)
are the parameters
Po
where
the covariance matrices
and
T denotes
least
squares
an
for data
initial
and
model,
Since the problem
solution is
found
show that
an
Cd
parameters,
transpose.
[1982a]
Valette
of
is
iteratively.
appropriate
0
and
are
_p0
respectively,
nonlinear,
Tarantola
algorithm to
the
and
compute
P
is
Pk+l
=
Pk
+
(JTCd
J)-
0
where J is the matrix
TCd
of part ial
Ji j
For the
priori
ridge earthquakes
knowledge of
covariance matrix
a limit
we are
[do-m(Pk)] - .P -'(Pk-Po)l
0
=
(8)
derivatives
ami /P
j
in our study we have almost no a
the model
Cp 0 = a 21
parameters,
i n the limit
so we
where
set
a2
the
In such
.
+
minimizing
S'
=
[d-m(P)ITCd
- 1 [d-m(P)]
and the corresponding algorithm can be expressed as
Pk+1 = _k + (-dTCd
which
J)-ITCd - Ido-m(Pk)1
0
is the solutio n of the classical
nonlinear least
squares
problem.
The theoretical body wave Green's functions are calculated
in
the frequency dom ain
using propagator matrices
reciprocity theorem [Bouchon, 1976].
for each
layer is
seismograms.
st ored
for the
We ass ume that
by shear failure, so in
A
set
computation
ridge-axis
our analysis
the
of
of
and
the
these functions
synthetic
earthquakes
inversion
are
caused
solutions
are
obtained
using
double-couple.
communication,
component
is
the
constraint
Trehu
1985]
et
[1981]
al.
have shown
(<
insignificant
that
that
10%)
the
source be
a
and
Nabelek
any
non-double-couple
[personal
for two Mid-Atlantic
Ridge
earthquakes.
Procedural
Details
Long-period
Standardized
Seismic
P and
Seismograph
Network
(GDSN)
800
are
hand
0.5
sec
[Wiggins,
projecting
S waves
at epicentral
1976].
observed
then
The
the
epicenter.
obtain
We try to
each earthquake.
near each
other
When
are
ratio.
about
we
the
same,
The
tion.
a common
for
and are
Additional
based
weights
may
be
invertsed
P waves
first
is
the window
onset
to
and
the
end
of
applied to
coverage
stations
usually extends
first
and
best
magnifica-
to a
Station weights
noise
each
are
level.
for
uneven
the waveform
station;
20 sec
wave.
to
through
spreading,
about
water
the
ratios are
window containing
for
for
are equalized
compensate
a priori
of the
one with
largest
3000,
a
similar waveforms
the
the
background
The time
onto
azimuthal
40 degrees.
be
by
seismograms
with
and
of
to the
geometric
the
300
intervals
signal-to-noise
one with
Digital
the station
select
on
selected
is
from
different
distance of
station distributions.
to
from
attenuation
epicentral
variable
we
When the
select the
seismograms
component
stations
instrument magnification
corrections
common
several
at
are obtained
an optimum
available,
signal-to-noise
distances between
SH waves
azimuth
World-Wide
the Global
interpolated
horizontal
line perpendicular to
by the
(WWSSN) and
Network
digitized and
the
recorded
If
from
for
the
the window
is
too
short,
the
erroneous source
may
result
in
source description may
parameters may
too much
producing a poor
fit
The parameters
which
minimize the
weight
mean
the
M
E
1=1
= j=1
port ion
residual
wj
a time window
wat er waves,
best-fitting
square
N
E
r2
on
incomplete and
Too long
result.
to the earlier
of the
be
thus
of the
point
waveform.
source are those
r 2 , defined as
(sij-oij
(9)
j=1
where
sij
and
observed
weight
oij
are
seismograms
used
number
waveform
station
in
Mj
j,
the
used
in
to the
the
wj
normalized to
the
A
of
in
time
the
for
is
synthetic
sample
station
average of
counted
window.
all
formula
wj
i,
and
j,
both
is
The
weighting
an
(9).
to the
factors
by
the
SH
The
noise
weighting
the
in the
N is
a P and
twice in
inverse proportion
the time
and
number of samples
station with
inversion
are
start
is
of the
and
j
inversion
of stations.
station weights
prior
at
for station
time window used
total
the amplitudes
level
is
the
2
W
)
1
/ (
1 / N
=
(10)
2
where oj
errors
tion
in
an
the estimated
the data
of equation
In
by
is
the
(9)
and
model
yields
inversion,
average
the
mechanism at
couple mechanism
is
noise
variance at
are
its
specified
is
the
If
the
minimiza-
likelihood estimate.
assumed
centroid
by
j.
normally distributed,
the maximum
source
station
to
be
location.
strike,
slip
represented
The double
and dip
angles,
[1980]
specified
and
given
(strike, dip,
history
is
is
with
for
represented
by
of the time
amplitudes
all
in
of
the
angles
and
Richards
degrees.
The source
amplitude.
elements
time
in
and
a priori,
the inversion.
The
length
function element should
be
time
interval,
depends
on the
frequency content
If the time element
length
is
signal.
the
result
is
an
the other hand,
description
other
instability
it
is
of the
too
in
long,
source and
source parameters,
most
function
elements
needed
However,
there
is
generally
the
total
duration
it
important
depth
the
of
and
inversion
the
is
source time
negative
value
of
set to the minimum
the
result
possibly
likely
on
not
to
function become
estimates
let
time
element
centroid
function.
a significant
is
of time
duration.
between the
negative;
on
of
The number
the source
of the source
If,
a poor
biased
some trade-off
time
of the
to be too short,
is
depth.
of each
resolvable
estimated amplitude.
depends
the last
chosen
duration
but the
time
which
function
functions of
The number
are chosen
form
The time
of overlapping triangle
variable
determined
Aki
abbreviated
in
be complicated.
a series
function
are
convention
sources
with
to
duration and
equal
all
slip),
allowed
the
however,
In
portion
a small
allowed.
Since we i gnore source finiteness effects, the synthesized
centroidal
the
seismograms
observed
iteratively
sometimes
first motion.
During the
realign the observed
if a better cross-correlation
allowable
shift
is
0.5
sec,
begin earlier
and
can be
the
same
or
later than
inversion, we
the synthesized
found.
as
our
waveforms
The minimum
digitization
interval.
Inclusion of this
horizontal
mislocation
is
time
not
shift
is
included as
inversion, we have to remove the lateral
important;
a parameter
variations
since
in
the
in travel
The derived focal mech anism, source time function and
time.
centroid depth all depend on the proper alignment of synthetic
and observed waveforms, so these time shifts are necessary to
obtain an optimal
solution.
For all mid-ocean ridge axis earthquakes the waveform
inversions were performed with the same source velocity
a wa ter
structure:
crustal
layer with
layer of th
a = 6.4 km/s
and
ckness
B
3.7
variable thickness,
6 km and P and S-wave velocities
km/s,
respectively,
and a mantle
halfspace with a = E3.1 km/s and B = 4.6 km/s.
the water layer is
a single
The thickness
1first estimated from bathymetric maps,
then
further constrai ned by the predominant period of wate r column
reverberations, seer as pr ominent phases immediately
A cru stal thickness of about 6 km ha s been
the P wave arriv al.
reported for sev eral1 axial
valley segments
Mid-Atlantic Rid ge [ Fowler
and Detrick,
foll owing
hi le
h
198 5].
1976;
Bunch
and
along
the
Kennett,
uniform-velocity crust
and
Pu rdy
1980;
ma nt
le
layers are overs impl ificat ions to the actual structur e, we
demonst rate
bel ow
general ly i nsensit
iat the
,e to d etails
Ot her aspects
fol 1ow the
waveform inversion
procedure
data
of
source
reduction
of Bergman
we empl oyed the conve ntional
et
values
results
structure.
a nd waveform
al.
of
are
synthesis
[1984].
partic ular
1.0
s for
attenuation parameter t* [Futterman, 1962]
and
the
for long-period
P and
21
SH
waves,
and
between t*
cannot
time
source
P waves
are
alone [ Nabelek,
the June
by
examplI e:
M s = 5.5)
in
the no rt hern R e d
mechani
short
of
amplitudes
seismic moment
first
only
fix
the
solution
Although most
amplitudes
MAT,
KBL,
are
waves
and
P waveforms
obtained
a
-p eriod P an d pP waves.
he did
not
Since
parameters we
y
Fig ure 2.1 shows
nvers ion.
give
him
hat
and
use
Pearce's
the obse r ved long-pe riod waveforms well.
polarity
in
wrong.
are definitely
h ave
SH
aves is correct,
Syntheti
P waves at COL,
the P and
amplitudes too sm all and
all
synthetic SH
2.2 sho
ws the solution after 5 iterations using only
and
rce's [1977] solution as the starti ng model.
The mechanism
quite well,
Pea
(21/6 0/345)
different
inversion,
pure
is
from Pearce's
but
first motions.
nearly
Pearce [1977]
Siea.
= 5.5,
too large
Figure
very
1972 earthqua e (mb
28,
relative
f
NDI
ustrate this
of the
of th e
some
We il
1984].
sm (335/ 7 5/2 ) fr om an analysis
the
not
1 mecha nism is obtained than
e couple oriienta tion as given
doubl
does
as P waveforms in a formal
ce ntroid dep th, to obta in thes
and
in
P waves
foca
used
principle
strike-slip
values of t*
f rom our invers ion.
as w ell
constrai ned
a better
inversion
more precise
function,
SH wavefo rms
By including
significant trade-off
is
SH waves
Including
of
there
det ermi ned
independently
be
Importance
if
Because
respectively.
after
SH waves
at
Whe n we then
5 more
normal
Most of the P waves fit
solution.
NIL,
an d
CHG,
include the
the
mechanism
fit
e-slip, althou gh it is
strik
iterations
faulting
wavefo rms are quite well
still
(Figure
MAL have the wrong
SH waveforms in the
s olution
conver ges to a
(303/40/278).
2 .3).
All
This focal mechanism
is
very close
Chapter
to
4 using
the
final
different
faulting mechanism as
From the above
provide
phase
essential
DEPTH
the
stati on
example,
clear that
i n focal
a
pure
SH waves
mechanism
should
information
and
in
normal
model.
is
it
(288/40/260) obtained
weightings
start ing
information
and amplitude
ambiguous
solu tion
be
can
studies.
employed
to
Both
avoid
solutions.
RESOLUTION
the centroid
Since
perhaps
the most
the depth
particular
phases
correct
source
new sou rce
our invers ion
ridge axis
parameters
procedure
For deeper earthquakes,
the free su rface are useful
so urce
dur
of
attention.
fro m
nt
import
resolution
fou nd f or
depths
depth.
However,
for
in
here,
reported
deserves
the
refl ected
the
determining
shallow
are
events
the
earthquakes ,
ation becomes much greater than the travel ti me
between th
e direct and
difficult
to estimate.
centroid d epth
reflected waves,
A major
and the
limitation
in
depth
source
determini ng
is
the
using long-period body waveforms is the trade-
off betwee
n the duratio n of the source time function an d the
best-fitti
ng depth [Ste in
further dr
awback in our
number of
time function elements and their durations ar e chosen
a priori.
total
Kroeger,
inversion
This method, if
predetermi ne
and
and
not
the best fitting
1980;
[Nabelek,
used with
depth
Forsyth,
1984]
care,
th at
is
can
1 9821.
act
since the centroid
the
to
depth
source time function duration are closely rel ated.
To address this
pr oblem,
we
first
fix
at each of a range of preassigned values
the
and
centroid
invert
A
depth
for the
remaining
sou r ce
different
comb inations
parameters.
At
(i.e.,
number and
residual.
the
smallest
the
iteration
At
the same
we constantly
at
solution
solution
a fixed
depth,
in th e neighborhood
Following thes e
steps
to guidelines
a nalogous
is
frequency
doma in)
and
A number
in
best
[1982]
a plot
the
of
t he
next
section.
for the
1981,
have been used
from body
wave data,
He
with
compactness
fo und that
tr ial
(o r
a set
were
of
Using
the
the
from
and
Weidner
Aki,
1982].
to
but
determine a
they
mechanism.
delta functions
P,
pP and
estimated
the complexity
(in
deconvolving the
of three
amplitudes
phases
depth.
simplicity),
the depth by
focal
depth;
according
source depths
earthquake
Romanowicz,
reflected
the
This overall
used
1978;
relative
in
versus
best fitting depth
[e.g.,
seismogram with
constraints
residual
phase spectra
determined
obtain the best
fixed-depth
requ ire prior knowledge of the
representing
varies
the
and
depth
observed
solution.
other methods
f ocal
best-fitting
generally
of
relax all
during
of
estimate
Patton,
noted above,
Once we
to the one commonly
wave amplitude
Aki
depths.
to
as
overall
described in the
procedure
times
of time
usually yields the best
we obtain
we estimate th e uncertainty
windowed
many
the synthetic and
fit.
we then
fur ther iteration
inversion;
Forsyth
duration)
time,
realign
waveforms to improve the overall
1973;
we try
el eme nts and choose the source time function that gives
function
surface
each depth,
sP.
Delay
from trial
of the deconvolved
the criterion
preferred
signal
of maximum
focal
depth
is
identified as that depth which gives the simplest source time
function.
Wiens and
one
nodal
After
the
P waveforms
the
lowest
time
several
function
trial
(defined as
the
and
Rayleigh
mechanism, they
focal
depths.
was
gives
the
as the
chosen
between source
a
with
including models
by
modeled
that
The depth
the squares of
sum of
handled
is
very
motions;
first
using
estimated
problem of trade-off
The
depth
P-wave
is
First the
depth.
focal
seismograms)
the theoretical
for the event.
depth
are
slip angles
a method that
from
determined
determining the
at
error
minus
observed
is
plane
and
then the strike
wave data.
used
technique to determine
similar to our
dip of
[1984]
Stein
range of trial time function durations and depths [Stein and
[1985]
used
and
Kikuchi
Kanamori
of
residual
methods
are
does
not
mechanism
of
is
not
good
depth.
tests
is
assumption
in
shown
of
coverage
[Nabelek,
of
depth.
a fixed
they
of depths
obtained a
all
these
source mechanism.
of
to
bias
the
focal
susceptible
of either
We
because
time function.
have
estimates
less
al.
source time
mentioned before,
prior knowledge
basic
is
Centroid
that
the
Depth
inversion technique
centroidal
adequate
1984].
chapters, the formal error (2
0.6 km for the
As
et
procedure
for
residual
a series
the Uncertainty
station
serious
at
prior
Engeln
to determine the
smallest
our procedure
require
Numerical
azimuthal
versus
or the source
Estimation
provides
the
and
[1984]
the deconvolution
[1982]
process
based on
consider that
it
gives
repeating this
plot
Stein
and
a variation of
function that
By
Wiens
1980].
Kroeger,
As
and
parameters
when
noise contamination
described
in
later
) is centroid depth
earthquakes in ou r study, but becau
0.1
of
bias
introduced
by
factors
a posteriori
parameter
the
duration
of the
and
the trade-off
error
a reported
deviation
Limits
some
risk
of
error,
being
as
a nd the
Strictly
greater.
that
however,
can
by which the
to
centroid depth
centroid depth,
the actual
speaking,
is,
the
magnitude and sign of
u sually
is
generall y
incorrect--from the
reasonable limits
the trade-off between
function
the true solution,
measurement process
from
is
such
i n the estimate of
source mechanism and
estimate,
from
to the
variances,
between
uncertainty
its
reflected
source-time
the true
in
not
be inferred--with
pre cision
of the
value was obtained and
reported
the possible
unknowable.
bias
of the measurement
process.
To
estimate the
uncertainty in
conducted a test of significance
best-fitting
depth
performed with
significance
the null
it
is
is
in
of the
does
not
best
i.e.,
and
whether
to examine
there is
not
is
A test
accept
of
or
reject
whether we would be
a difference
hypothesis
data do
to
between the
inversion
values.
between the
A nonsignificant
other depths.
the null
our
when the
nearby
deciding
We want
depth
the residuals
is
give
correct,
adequate
result
merely that
grounds
for
it.
Since we
problem,
for
for
obtained
fixed at
concluding that
tenable,
in
a rule
prove that
rejecting
used
the depth
hypothesis.
justified
fit
and those
the centroid depth we
are
equation
we
instead of
use
the
dealing
(9)
the
are
with
not
residuals
residuals
time
series,
independent.
at
To
individual
at discrete time
the data
get
around
stations
samples.
points
as
This
we
this
our
data
approach
will
underestimates
yield a
very
the number of
conservative
independent data.
The mean
investigated
rj 2 with
the
residuals
can
generally
distribution.
This
distribution,
and
as
of
goodness
be
is
square
and
residual
number of
Billingsley,
The quantities
all
stations
However,
have the
because
19821,
that
in
R2
the
focal
contained
in
mechanism is
information
the
differ because
in
found
that the
rj 2 should
of
of
residuals
a normal
a x2
follow
freedom
increases
distribution
We
also redefine the mean
2
(12)
(12)
would be
equal
inversion window length
as
the
of X 2 in
a measure
are
For
of
the
seismograms
of the radiation
and
only
depth
much
of
SH waves.
of
pattern.
the
if
weight.
F-test
[Steel
of
and
information differs
example,
P waves while
in
an
not homogeneous
amount
contained
contained
station
follow
degrees
(9) and
rj
seismogram to seismogram.
mostly
(11)
j=i
same
is,
is
as
rj
R2
we cannot use
significance
Torrie,
r2 and
set
a normal
N
1
N
=
and
to
1981].
for the model
2
R2
of
2
individual
considered
the X 2 distribution approaches
[Huntsberger
assured
station j
for
oij)
of fit
because the
the
-
are
and
1=1
distribution
the x 2 test for
residual
(sij
of freedom
but we
Mj
-
Mj
We
result,
square
1
2
rj
degrees
from
information is
the
constraint
Also, the
same kind
of wave
on
If heterogeneity of the rj2 samples
may be wise to consider the individual
than R2 [Steel
and Torrie,
1982].
consider the difference in rj2
focal
depths.
observations
That is,
is significant, th en
station residuals
Following this
it
rather
idea, we
at each station for two diff erent
we consider that our experimental
(residuals) are self-paired.
pair tend to be positively
the
If the members o
a
correlated, that is, if members
pair tend to be large or small together, an increase in the
ability of the experiment to detect a small
difference is
possible.
We now consider the matched station
A and
B and the set
of differences
dj
where
B is
occurs.
dj
We
the
depth
sampling
2
=
rjA
at
whic h the
denote the
by Ud and ad,
- rjB
with
mean
hypothesis
tested
is
is
than
The test
zero.
zero;
OD and
that
j = 1,2,...,N
R2
deviation of the set o
of dj may be regarded a
population of station
standard deviation oD.
th e mean
residual
The
null
of the population of
the alt ernative
is
that the mean is
larger
is
statist ic
t
(13)
smallest average residual
The set
a normally-distribu ted
differences
2
and standard
mean
respectivel y.
differences
residuals at two d epths
(14)
=
ad/N
which follows the t-distribution with N-1 degrees of freedom.
Since we are testing the hypothesis that the solution at
one
depth
is better than the
one-sided)
considering
the
Specifically,
is,
differences
and those
exceeds
we
the
at
allowed
ridge-axis
general,
value
for
is
residuals
(or
both
are
consistent
error
is
r 2 and
of
with
in
R2
the
none
in
as
(Type
in
Figure
and deeper.
the t statistic
level
error
centro id
this
of
there
is
a
see
manner
for
along with
source
about
four
the
depth.
In
by the
dept h estimated
of
That
of 0.1.
I error,
2 .4,
functions
centroid
a formal
the
wh ich
determined
shown
the best-fitting
shall ower
at
by
inversion
to conclude that
of 0.1
ranges
resolution
st andard
the
depth
the
residual s at
a significance
fact there
earthquakes
the
t-test
in
The
square
between
depths,
have a probability
Appendix A).
s
nearby
resolution of
we determined the depths
difference where
mean
we use the one-tailed
test.
investigated the depth
We
depth
other,
10 Gh,
where
depth h determined
in
Gh
the
inversion.
aveform
range
The
symmetric;
in
is
it
uncertainty
usually
about
smaller
at
the
best-fitting
the
shallow end than
is
depth
not
the
at
deeper
end .
earthqu
ake s, p robably because more stations are used for la rger
events,
should
degrees
recalled that we have
of
freedom for all
indepen den t
not
to
be
smaller
for
larger
the
number
of
these tests in orde r to be assured
it should be noted that
th e sole
appreciation
underestimated
It
data.
Fi nal ly,
are
range tends
tha t the focal mechanism is better constrained.
so
be
The uncertainty
of
basis for
the problem
drawi ng
and
statistica 1 considerations
inferences;
experience may
a physical
also be
brought
profitably
into the judgment.
waveform
shapes
which
residual
can
important
particular
and
be
are
solution, we
synthetic
poorly
clues
at
reflected
in judging
at
Ridge earthquake of June
6, 1972.
of mean
versus
very
flat
at
residual
depths
greater
uncertainty
statistical
conclude
that
constraint
waveforms in Fi gure 2.5
m
ki
of 1.8
a larg
produces
of
puls
progressively
al lows
The solu
e enough dil atat
1 a rge
too
the first
observed
the
Mid-Atlanti c
the
pl ot
(Figure
2.4)
and
range
basis
place
On
the
the
alone,
virtually
is
of
we wou Id
no
the other hand,
the
us to distinguish a centroid
to
or
on at CAR.
e at CAR
is
too wide.
The
a ggravate these probl ems,
remains
nearly
ac ceptable ce n troid depths
study
of the
even
is a good,
true
1 km
The solution
a di latatio n at AAE and AQU,
range
statistical
solutions
ion at a depth of 1 km does not
error
estimate
depth
this
data
residual
in
2-5
earthquake, the
1.5 km,
ev ent.
this
as bein g s uper
d e eper.
shallower or
produce
On
long-period waveform
on the depth
for
For this
large.
of a
the success
generated with
5 depths
a bout
square
the mean
Figure
ce ntroid
than
in
of
detai Is
how
of the first half-cycle of motion of the synthetic P
character
depth
is
in
3 sta tions
source parameters
squared
illustrate
have plotted
P waveforms
best-fitting
To
uncertainty
in
solutions
although
const ant.
We
rather
the
at
the
in
3 ki
m
width
4 and
5
overall
conclude that
determined
although
and the
at
the
the
conservative,
depth
estimated
in
the
inversion.
The
Effect
of the
The accuracy
Assumed
of
Velocity
source
Structure
parameters
depends
on the
velocity
structure
th e
of
Since we do n ot have direct
sou rce region.
velocity structure at most earthquake sites,
knowledge about
the
we
ized velocity model
use a genera
events,
as
noted
for al 1 of
our
ridge-axis
earlier.
of the events in thi s study, we f ound that the
For most
versus depth displayed tw0 minima,
mean square residual
one
located in the crust and one just below the cru st-mantle
interface
(Fig u re 2.4).
The question
of depth
therefore invo I ves two separate issues.
The fi
resolution
rst concerns the
choice between the two minima in the relation f or residual
versus depth.
The second
in the depth a
With the
the Central
which
issue
is
the estimati
the preferred minimum
xception
of one
earthquake
occ
(Octo
on of uncertainty
urs.
ber 7, 1981,
on
In, ian Ridge), the solution within the crust has a
smaller residu 1 than the sub-Moho solution and would therefore
be the preferri d solution.
However,
are similar fo
both
and the difference
usually small.
Because the deeper solution
below the crus
depths
-mantle
interface
in
the other
the
lie
source
source parameters
in residuals is
s immediately
velocity
structure for m ost earthquakes and occurs withi n about 2 km of
the Moho for al1 events, we suspect that it is an artifact of
the
assumed
sharp veloci ty
To explore
solutions,
the
crustal
we
the
show
reas
in
and mantle
earthquake of
Fi
ons
for this apparent ambiguity in
gu r e 2.6 several
s olu
September 20,
residual-versus-depth
contrast.
cu rvye
P and SH waveforms for
tions for the Reykjanes Ridge
1969, which has sharp minima in the
at 1.5 and 6.0 km (Figure
2.4).
Other
than the centroid
for the
two
solutions.
that the wavef
solutions
orms
for
teleseismi
Matsushiro
body
res
two
of
After the pP phase
phases
on
ou r
the
ab il ity
f ree
phases
are
the
however, the
in
source tit
b
the
very
o nset
s
e two al ternative depths
When
me
fu
we i ncl ude the effects
nction, and anelastic
roadeni n g an d coales cence of P and p
a si gnif ic ant
For
th
to i dent ify the contribution to
su rface.
det!e rio r ation of depth
hal low
1 arge
earthqu akes, we usually
of the first motion to better than
1 sec on long-period records.
arrival
solutions,
and f or this section of the seismogram the
instrume nt,
result
identify
two
the principal
shing the sol utions a t
from
discrimination.
cannot
ations,
entirely
attenuation,
the
pP
are almost identical
solutions
recording
the two solutions at
t iming appropriate to the
ampli tude and
0.5 and 2.0 s ec for
reverber
signal
so is
km depth is the del ay time between the P and
6.0
Distingui
the
this is
The only not able di fference between the solut ions
(about
hinges
That
ponse, time fu nct ion duration, and anelastic
respectively).
water
remarkable
wave phases, without the smoothing effect of
attenuation).
phases
glance,
similar
Figur e 2.7 sho ws elementary seismograms
r"elative
principal
1.5 and
are
elementary seismogram is a series of delta
(an
of
at
at first
c P waves pre dicted by
functions
instrument
seem,
may
parameters
the l imited ba ndwidth of long-period
gnals.
si
It
source
cen troid depths
of
basically a re sult
teleseismic
all
in Figur e 2.6 are so similar for the
t he t wo
at
depth,
times of P and pP can
A small
difference
between the
therefore be taken up by
of synthet ic
realignment
synthetic waveforms
i ncl udi ng
of the
the effects
inversion
earthquake
at
seismograms
to
used
can
in
centroid
onset
depth
of the
between
To
the
are
first
and
of tests
crustal
sea
thickness
minimum
in
2.8
r2
differs
very
shows
similar,
is
the choice
for t
the effect
crustal
and
there
floor.
a crustal
develops
at
the
about
minimum appears immediately
As
6 km
t identify the
1 sec t hen choosi g
is
of solution s, we
he September 20,
1
our first
two minim a in r 2 .
m antle
on
In
1969
of crustal thicknesses on
t he
For
S velocities
for
at 1.5 and
model s.
crustal
thickness.
assumed
stations.
i f we canno
th an
or
sy nthetic
f ive other
at
magnitude of
When
is less than
depths
difficult
a variety of
crustal
For all
inve rsion wind ows
o f the
and
inversions
we examined
onset of the
to the
comp arative
to better
further
we kept the
is
tions,
this earth quake, in fact,
for
for solutio ns
motion
of
sol
the
by only 0 .5 sec.
centroid
solutions
relative
the
P and
below the
in
and
tests,
varied
two
two solutions
earthquake using
these
2.7)
waveforms
investigate
the depths
correspo nds
inversion
at the
performed a series
series
the
the two
the
km
For
of all transfer functions, the beginning
Figure
for
be seen,
6.0
between the beginning
1 sec.
seismograms
1 5 km and
(Figure
the solutions
equal
the
waveforms.
observed
window which
MAT
the difference
for
of
and
veloci t ies cons tant
a halfspace model
with
ly one minimu m, at 1.4 km
ayer greater than 1.5 km
source
structure, a stable
1.5
depth and th e second
km
For
below the crust-mantl e interface.
With a progressively thicker crustal
layer
this
second minimum
moves
with
the
layer boundary.
These
results
suggest
that
the
sharp velocity contrast at the Moho acts to reflect energy
downwa rd in
a manner similar to
the water-crust
boundary.
For
a
source located immediately below the Moho, the upgoing P wave
reflected
down ward
at
the Moho
direct P but
o pposite
of
P arrival
the
first
polarity.
of a shallower crustal
destructive
just
Moho;
effect diminis hes
To test
to
is
the
maximum,
has
residual
an additional
increase
solution
reduced.
For the
additional
loc al
minimum
between the cr us tal
In
princi pl e,
th
distinguish
inversion
we
of
as
the focus
moves deeper,
is
the
inversions with
in
the
across
becomes
the
just
a
lower crust
the crustsignificantly
velocity contrast
2-layer crustal
also appears
this
below the
model,
at
an
interface
layers.
short-period
records
can
be used
to
e crustal and sub-Moho sol utions indicated by the
1on g-period waveforms.
consider th e June
long-period
layer
solution
that
larger.
in velocity
crustal
prope r is
when
repeated the
worse than
the Moho
the amplitude
largest when
is,
becomes
The sub-Moho
best
that
is
earthquake focus
mantle transit ion.
the
makes
The effect
hypothesis, we
simulate a gradual
effect
time as
source.
as
which
velocity model
This
arrival
and the seismogram mimics
and the
t his
the same
smaller,
in terference
below the
has
2,
depth
an
1965 Mid-Atlanti c
bo dy waveform
residual -versu s
As
inversion
curve
(see
illustrative
Ridge earthquake.
show s two minima
Chapter
example,
3 and
in
Appendix
The
the
B),
similar to the double minima shown
in Figure 2.4.
fitting
a secondary solution occurs
solution is at
3 km depth;
The
best-
immediately
SH
below the Moho.
waveforms
tatistical
are very
The synthetic
similar for
the two
long-period
P and
solutions, and
test cann 0 t distinguish them.
Figure
ou r
2.9
ompares P waveforms f or this earthquake recorded on four
hort-perio d vertical
WWSSN
instruments w ith
synthetic
short-period waveforms generated for the solutions at
and 6 km cent roid dept h.
3 km
Because the abs olute amplitudes
of
the short-per iod wavef orms appear to be c ontrolled by path
effects,
only
comparison.
preferable
the shap es of the waveforms are used for
Figure 2. 9 shows that the 3- km-deep solution is
because
all
reflection pwP are in
waveforms.
For th
phases
good
before
agreement
the
first
water-surface
with the observed
e 6- km-deep sub-Moho so lution, the pwP
phase in the synth etic waveform is late
respect to the obs erved seismog ram.
Thi
about
sec with
discrep ncy
particul arl y clear at stati ons SJG, WIN, and PEL.
is
The
ability of t he short-period waveforms to discriminate between
the two sol u tions is due to greater bandwidth in the recorded
signals; th e beginning of f irst motion can be resolved to
better than a few tenths of a second.
From th e above analysi s, we infer that the solution
immediately
below the Moho
indicated by
the
inversion of
long-period body waveforms from the ridge-axis
this study
earthquakes
of
is an artifact of the sharp velocity discontinuity
the crust-ma ntle transition in the assumed source velocity
model.
It is quite probabl e that
overestimates
our
simple
source
structure
the Moho velocity contrast beneath a slowly-
at
spreading
ridge
Purdy
[e.g.,
that the shallow crustal
crustal
and
Detrick,
solution is
solution for these earthquakes .
at
Tarantola
reasonable
Valette [1982b]
and
ikeliho od
1 ayers of const ant vel ocity.
interfaces between
is because the discontinuiti es
This
any
conclude
rep res en Its the true
also noted the existence of secon dary maxi mum-l
solutions
We
stable for
pro bably
it
thickness and that
1985].
in veloc ty ca
the
probability density functions a(z ) to have d scont
slope, where a(z)
within a small
is the probabil ity that
depth
th
tru
Other source parameters were consistent
for a 11 of the
nclud ing the
different structural models we ha ve tested,
hal fspace mode 1s and the model with 2 crust
varied by
les s than 100,
and
the
with
inversions
Table
less than 20%
e
,str
different
velocity
rs.
dip varied
and
structures
The
The
results
are
shown
in
2.1.
The
second
using our
part
of our
generalized
velocity
structure.
about
any
introduced
bias
A refraction
experiment
along
a section
where
several
June 28,
1977
the
This
by
simplified
the inversion
with
comparison
using
and
[Purdy
that
can
using
give
a simplified
Ridge median
(June 3,
Figure 2.7
shows
the
derived
models
by
Purdy
with
and
a known
an
was
idea
conducted
valley near
study
of
us
results
velocity model.
19851
Detrick,
this
earthquakes
swarm).
compares
structure
of Mid-Atlantic
velocity versus depth
as
test
ridge
ridge
well
laye
varied by less than 200.
slip
ution lies
e
dz at depth z.
interval
sei smic moment
ties in
1962 and
23 0 N
the
compressional
Detrick
a uniform
crust
[1985]
and
3
as
homogeneous crustal
layers.
velocities used
The shear
inversions are estimated by assuming a Poisson's
in
the
ratio of
v = 0.25.
Figure 2.8 shows the residual
velocity
different
977 swarm i.
struct ures
For the mode
learly de fined local
versus depth curves for two
for the largest earthquake in the
1 wi th a unif orm crust, there are two
min
one with in the crust at 1.6 km
ima,
beneath th e sea floor and the other im mediately beneath the
Moho.
The uncertainty in the depth of the shallow solution
T he
local minimum below th e
Moho lies outside the unc ertainty rang e.
For the model with th e
spans a depth range from
1 to
2
3-1 ayer
crustal
structure
the
flatter
There
are three
local
km.
residua
minima
occurs in the crust at 2 km below the
1 curve is generally
:
the global minimum
sea floor, the second
minimum lies just beneath the interfac e between the second and
thi rd c rus tal
the Moh
o.
layers,
Since the
the
third
mi
resi dual
curve
is
and
nimum lies just beneath
flatter, the uncertainty
range f or centroid depth is greater (1 .2 to 3 km and 4.4 km to
5 km de pth ).
global
This
analys is
min imum probably
of the velocity
i s generally
structure.
fitting centroid depth,
indicates
of
Ideally
to be inverted
P and SH waves.
for the source
region.
Coverage
a body waveform inversion study,
sample fully the lower focal
All
to the details
however, is probably underestimated when
Incomplete Azimuthal
in
insensitive
The uncertainty range for the best
we use a simple velocity model
The Effect
that the depth of the
too often,
the waveforms
hemispheres
however, the sampling by
for both
available
seismic stations is incomplete, particularly
azimuth.
It
is
therefore
important to estimate the effects
incomplete cove r "age on the estimation
parameters.
nd
We chose the September 20
resolution
1969
earthquake to in vestigate these effect .
largest ridge a x is
in
earthquake we have
of
Reykjanes
This
tudied,
source
Ridge
earthquake
and
is
the
the
signal-to-noise ratio for all observed seismograms is quite
We vary t he azimuthal
good.
coverage by selective rem oval of
data, and we rec ompute the residual
residual
appears
ve rsu s depth curve .
versus depth curves using all
in
Figur
also shown
focal
th e
in
24 P and SH wav eforms
e 2.9 and is used as t he standard resu lt for
The
comparison.
sphere coverage fo r the full
In
the
t t est,
firs
stations that are close to other stati ons
azimuthal
wave s is
the
cov erage.
shown
in
wave forms.
identic
al
The best
full
solution.
cent
to
ro id
Th e unc
also identical for
2.10;
station coverage and
in
Figure 2.11.
ta set is
All
The
s
dip and
ip
bes
obtai
angl
h P and SH
model
parameters
ned
sing all
24
es d ffer by less than
nges
fo r
the same
centroid
as
in
the
depths
are
sets.
this
test,
we
eliminated
for both P and SH waves.
residual-versus-depth
centroid
bot
t-fi
ertainty
its
for
retain
figure is
is
qua drant
we still
he same
5 km
data
we omitt ed some
in
depth at
both
first
sh own
also
the resu
3-quadrant coverage.
stations in the
but
cov erag
station
th curve.
The st ri ke,
20.
The
Figure
residual-versu s- dep
are almost
da
figure.
4-qu adrant co verage.
full
The
data
from
The
curve are shown
parameters are reasonably well
The depth
determined.
couple
angles
However,
6-9
km)
mostly
is
much
from
P-wave data
retain
residual
to
all
curve
are
parameters
by about
150,
centroid
depth
centroid
The second
The
is
at
is
6
depth
well
double
result.
km
(1-4
This
and
is
constrained,
the
range
220
this
test,
we
angle
is
to 290
for
can be
have
increases
for the
as
coverage,
significantly
as
the
are
the
the
standard
the
model
is
different
for
due to a
focal
1.5
depth.
km,
the
of good
coverage
are
includes
very
stable.
the best-fitting
from the
uncertainty
azimuthal
at
with
solution.
parameters
however,
The best
range
importance
the azimuthal
from the
compared
range
occurs
best-fit
overestimated
same.
Moho)
large
curve
underscores
expected,
angles
but
its
quite different
source mechanism and
the best-fit
azimuthal
and
estimated
The uncertainty
residual
long
The
about
This
eliminated
second quadrants,
coverage
below the
km.
depth
As
are
slip
km.
to 10
analysis
and
2.12.
and
(just
1.5
the
In
coverage
trade-off between
coverage.
we might
km
of
0.5
4 quadrants,
solution
with
the first
strike
minimum in
For lesser
less
Figure
the dip
result
above
azimuthal
is
and
solution.
The station
in
for this
best-fitting centroid
The
in
data.
shown
while
depth
significant
depth
the full
compared
stations
SH wave
result.
As
in
km,
from the standard
for centroid
P wave coverage.
from
the standard
380
70
by 0.2
solution.
standard
3 or
range
larger than
180
standard
model
less than
uncertainty
2-quadrant
we
by
overestimated
because the strike angle
varying
the
the
differ
is
range
actual
source.
for centroid
coverage decreases.
Based
on
our
inference that
of the
velocity
(crustal)
are all
the
minimum beneath
discontinuity,
obtained
under different
quite similar.
confidence
in
then
This
the results
the Moho
is
an artifact
the best-fitting
degrees
internal
presented
of
azimuthal
consistency
in
solutions
coverages
gives
us
the chapters to follow.
COMPARISON WITH SURFACE WAVE METHODS
Two Mid-Atlan tic Ridge earthquakes
16,
November
es give u s
earthqua
indicates
we
For
s oluti
mantl
on immediately
tions,
rthquake are iden tic
For the
Novembe
r 16,
by
Aki [197 3] (3
We consider that
inversion
and th at
resolution
t he
+
for
the
using these two method S.
obt a inable by
the so rce
signal-to-noise
ratios are high.
(± 1 km)
than that
indicated
2 km).
depth
comparable when
should he noted,
If
a sligh tly shallower centroid depth (2.1 km) and
er depth
and
beneath the Moho.
19 65 event, our body waveform inversion
a slightly bett
Weidner
our inve rsion
resol u tion (± 2 km)
and depth
1965 ea
suggests
estimated
for reas 0 ns descr ibed above
2,
result
These two
obtained by a detailed surface w ave
thes e earthquakes,
centroid de
June
in Chapter 3.
rtunity to comp are our
that
of
a pos sible
exclude the
the
both
s ummarized
oppo
with
un ertainty
analysis.
an
using the phase and amplitu de
The body wave inversi on results for
earthqua kes are
tw
depth
[ 1973]
and Aki
f Rayleig h waves.
spectra
these
included in our study were als o examined in
1965)
b, Wei dner
detail
(June 2, 1965 and
resolution
using
body waveform
surface wave analysis
radiation
pattern
Both
is
of these
are extremely time-consuming.
well
are
sampled
methods,
The
and
it
surface
wave
method
requires
processing ,
a great
different source
method
along
usu ally
the
whereas
ray
bo dy
informatio n.
path,
or
waveform
For
sufficiently
robustness
respect
surf ace waves
method
is
su rface
for
to
events when
Kafka
surface wave
betwee n nearby sources,
not ne ed such additional
m ore compact
are the most
inadequate [e.g.,
the
we prefer the body
data
sets
sourc e velocity structure.
wave analysis
smaller
re gion
1 arge events
relatively
computer
requir *es testing of many
In additi on,
inversion does
its
events,
the
and
information on phase velocity
precise
in
wave metho d for
since
method
time functions.
requires
with
of digitizing
body wave
the
and
deal
energetic
so metimes
a zimuthal
and Weidner,
is
phases
for
and
However,
shallow
the on ly feasibl e
c overage
1981].
for
body
waves
Table 2.1
Results of Inversions for the Centroidal
Reykjanes
Model
no.
Crustal
thickness,
km
Centroid
depth
(crustal
solution)
Centroid
depth
(mantle
solution)
Ridge
Parameters of the September 20,
1969
Earthquake
Seismic
moment,
1025
dyn-cm
Strike/
dip/slip
Residual,
mm2
Remarks
13.5
20/42/258
10.9
Crustal
2.77
16.1
16.9
30/41/266
24/38/255
13.3
14.2
1.77 km below Moho
2.82
15.6
17.2
21/41/257
24/38/256
10.8
13.1
0.82 km below Moho
3.15
16.0
16.3
25/39/263
22/39/256
11.4
13.5
0.15 km below Moho
4.02
15.7
15.3
25/37/261
30/43/270
10.7
12.1
0.02 km below Moho
5.01
15.5
14.7
21/43/260
29/45/270
10.1
11.2
0.01 km below Moho
6.00
15.2
13.4
23/42/261
27/46/276
10.3
11.3
0.00 km below Moho
3.09
16.4
24/42/260
11.9
Mantle halfspace
6.02
15.3
13.3
14.8
25/44/263
28/45/273
30/45/270
1.40
0.99
1.41
1.44
1.45
1.45
1.50
1.68
4.07*
half space
0.02 km below Moho
Notes to Table 2.1
1 is a crustal half-space; models 2-7 include a single
Model
layer crust
over a mantle
half-space,
and
model
half-space; model 8 is a mantle
9 includes
a two-layer crust
over a mantle
half space.
For model
1-8,
the seismic vel ocities are
Crust:
a = 6.4
km/sec,
8 = 3.7 km/sec
Mantle
a = 8.1 km/sec,
6 = 4.6 km/sec.
For model
Layer
9,
= 6.4 km/sec,
1:
km/sec,
Layer 2:
= 7.2
Mantle :
= 8.1 km/sec,
* A second
the
local
interface
minimum
occurs
between the two
km/sec,
thickness = 4 km
= 4.2 km/sec,
thickness = 2 km
= 3.7
= 4.6 km/sec.
in the crust, immediately below
crustal
layers.
FIGURE
Figure
CAPTIONS
Compariso n of observed
2.1
no rthern
1972,
focal
momen t,
from
centroid
depth
synthetic
focal
mechanism
plotted on the lower
[1977]
(equal-area
he misphere
and
fo r the June 28,
Red Sea earthquake, with the
obtained from Pearce
solut ion
line)
P and SH waves
ion g-period
(dash ed line)
(solid
projection).
The
seismic
and source time func tion are obtained
the inversio n of P waveforms with the double-couple
orien tat ion
fixed at values given by Pearce
amplitudes are
to an instrument magnification of
normalized
3000; the amplitu de sca les corr
would
spond to the waveforms that
be observed on an origi na
i nst ru me nt.
The two ve rtical
seis
m ogram
1 nes i n the
the po rtion of each tim e series used i n the
Symbol
for such
P waves
an
delimit
inversion.
for both types of waves are o p en ci rcle,
s
All
[1977].
dilatation;
solid circle, compression; cross,
emergent
waves, compression corresponds
positive motion as defined
by Aki
and
obtained
Figure
2.2.
waves
from
for the
Figure
for
2.3.
waves
June
obtained
imposed
2.1
the
Comparison
solution
an
Richards
i nversion
28,
the
is
1972 event,
on
SH
also shown.
observed
f rom the
constrai nt
For
The source time function
[19801].
of
arrival.
and
with
inversion
the
source
of
synthetic
the
P and
focal
mechanism
P waveforms
mechanism.
SH
See
without
Figure
further exp lanation.
Comparison
for
the June
of the
28,
observed
1972 event,
and
with
synthetic
the focal
P and
SH
mechanism
solution
and
Figure
obtained
SH waveforms.
2.4.
Mean
earthquakes
indicate
(12),
square d
in
this
residual
r
2
interval
2.5.
earthquake
generated
at
waveforF
from the
not
6,
bes
whic
h is
shown
The
Vertical
marks
at
the
(see
delimit
equations
id)
dept
and
the
es the 90%
h found in the
an arrow.
by
synt
hetic (dashed)
Ridge
nthetic
seis
mograms are
obtain
ed with the
A dditional
value i ndicated.
3.1 6)
centroid
9) and
id-Atlantic
sol ution
Figure
preferred
in
four
es
dashed li
ndicated
The
best-fitting
and
both
for
centroid depth
fitting
(s
of
explanation.
bar indicat
the
for
1972.
fixed
inversic ins .
t ic
the
3 st ations
of June
centroid d epth
zont
observe d
Comparison of
P waveforms
hori
about
inver sion,
versus
further
R 2 , defined
and
The
for
T h e solid
study.
respectively.
body-waveform
unconstrained inversion
See Figure 2.1
residuals
confidence
Figure
from the
are
depth
port ions
used
is
in the
1. 8
of the
km.
waveforms
included i n the inversions.
Figure 2.6.
Coimparison of the crustal
mantle ( 6. 0 km centroid depth)
solu tions
Ridge ea rt hquake of September 20,
cent roi d Sd epth is 1.5 km.
( 1.5 km
tic
portions o f the waveforms included in
Figure 2.7.
Coi mparison
MAT for
th e
of
synthetic and
best-fitting solutions
The
preferred
marks
delimit
the
and
1.5
the
inversions.
observed
at
depth)
for the Reykjanes
1969.
Vertical
centroid
P waveforms
km and
at
6 km
centroid d epth for the Reykjanes Ri dge earthquake of
Septembe
of
the
r
20,
different
1969.
At
phases
the
top
are
the
predicted by the
relative amplitudes
radiation pattern
and
respective time
delays;
of convolution with
long-period
on
the
instrument
sei smometer,
the bottom t race),
subsequent
with the
and
with
traces
response of
source
Figure 2.8.
Comparison
waveforms
operator
at 2 stati o
P waveforms at 3 stations and
for
he two alternative solutions
rthqu ke of September
shown for each stati on and
Also
rel ati ve amplitudes
times of different
predic ted by the
phases
20,
1969.
sol ution are elementary
seism ograms depicting the
and arrival
radiation pattern
so urce depth.
2.9.
Compariso n of
short- period
Ridge
(3.0
are
k m centroid
centro id depth)
The pr eferred
2.10 .
ohs erved
P wa veforms
earthquake
seismo grams
Figure
(shown
time function
the attenuation
of synthetic
r the Reykjanes Ridge
Figure
a WWSSN
= 1 sec).
(6t*
and
effect
show the
of June
depth)
4
from
a nd the
o btained
and
tations
for the Mid-Atlantic
2, 19 5.
g enerated
cen troid
Models
at
(solid)
from
de pth
The
synthetic
synthetic short-period
the best-fitting
secondary
f or compression al
solution
solution
ong-period
i, 3.0
(dashed)
(6.0 km
waveform inversion.
km.
velocity
versus
the source region of rid ge-cres t earthquakes.
depth
The
in
solid
line
shows the v elocity-depth funct i on obtained from refraction
data along the median va lley
23
0
N [Purdy
model
and
Detrick,
with a homogeneous
of
1985].
the Mid-Atlantic
The
dashed
line
crusta 1 layer overlying
Ridge near
shows
the
a uniform
mantle assumed in the calculation of most synthetic
seismograms.
A model with three crustal
layers,
used
for
testing the bias introd uced by the uniform cru st model ,
also shown.
See text
for a discussion of the
varying the assumed vel ocity model
effects
centroid
on
is
of
depth
determination.
Figure 2.11.
Comparison
depth for
of the mean-squared residual
the June 28,
velocity structures
indicates the
dept
1 aye
2.
Figure
(Fi gure 2.7).
confidence
using two different
earthquake
horizontal bar
The
interval
about
h fou nd in the body-waveform inver sion,
a rro w I
an
90%
1977,
The
vertical
dashed
lines
R2 versus
i ndi
the
best-fitting
a s indicated
cate the assumed
r b0 u ndary.
12.
The mean-squared
R2
residual
versus centroid depth
for
the September 20,
The
station coverage and
are
also shown.
the
mean-squared waveform over the same time window.
196 9 earthquake using a 11
focal
The resi dual
mechanism
at each
for
depth
P and
normalization, a per fect fit
zero
and a source with ze ro seismic moment wo ul
res idual
of unity.
have
a
The horizontal bars indic at
the
0%
about the best-fitting de pt
foun,
body
-waveform inversion, indicated by an arro W.
2.
13.
The mean-squared residual
for the
Septembe r 20,
omitted
but
Figure
Figure 2.14.
for the
1969
full1 azimuthal
2.9 for
of
have
idence interval
R 2 versus
in the
centroid depth
earthquake when some stations
coverage is still
b
With
a residual
conf
Figure
SH waves
i s normalized
this
would
stations.
retained.
are
See
f urther explanation.
The mea n-squared
Septembe r 20,
residual
R 2 versus centroid de pth
1969 earthquake when the data from
47
stations
in
the
eliminated.
Figure
2.15.
See
first
Figure
2.9
The mean-squared
for the September 20,
from
quadrant
stations
eliminated.
1969
in
the first
See
Figure
for
for both
further
residual
R2
P and
SH waves
explanation.
versus
centroid depth
earthquake when the
and
2.9 for
second
quadrants
further
are
P-wave
are
explanation.
data
Figure 2.1
JUNE 28,1972
(P waves only)
SH WAVES
P WAVES
S4
Source Time Function
NUR
KEV
(sec)
\
N
IIL•
H
I
'
ESK
I
C HG
S
:
It
1V
VAL
MAL
V
mm
mm
70
70
sec
sec
Figure 2.2
Figure 2.3
20
20 rSeptember 20, 1969
June 6, 1972
15-
10
t
I
I
I
2
I
4
I
I
6
I I
10
8
i
01
20
1
2
1 I 1l
6
4
1
8
I
i
10
r-
April 22, 1979
June 28, 1979
o-
I
-
..
SI
I
2
I
I
4
6
I
I
8
I
I
I1
10
I
,
Centroid Depth, km
Figure 2.4
2
I
!
4
I
1
6
II
I1
8
10
AQU
AAE
CAR
Centroid
Depth
km
I
1.8 km
-II
I,
4km
-
4 km
-
Af
Ii
5 km
se
sec
Figure 2.5
SEPTEMBER 209 1969
1.5 km
6 km
50 sec.
MAT P
SDB P
SHL P
-NAT
P
NIL P
SJG P
TAB P
BEC P
TRIP
ODG P
NAI P
-
AAM P
c
Figure
2.6a
SEPTEMBER 20, 1969
1.5 km
6 km
50 sec.
GOL P
j-
NAT SH
CORP
COL
P
NIL SH
-
-
SJG
SH
BEC
SH
O-GD SH
TAB SH
GOL SH
TRI SH
COL SH
u1
Ln
Figure 2.6b
6 km
1.5 km
pwP
p5wP
p 3wP
IMPULSE
S
*
p6wP
p4wP
pwwP
PpwP
p7wP
,
I
I
.
I
p3wP
I.
pwwP
p5wP
I
,
p4wP
p7wP
,
p6wP
P ,
ti
+
a
a
INSTRUMENT ----------- P !',
A
S,
4'
*
a.
ISO UMETIM
I
'
.
i
I
*a
I
a
a'
'
aI
o
a
a
a
i
,
,*
/
• L=
;
a
1 a
a
,
aa
a
a
I,
,
a a
a
aI
+ SOURCE TIME
FUNCTION
a
e
°
,
a
aa
a
I
*
I
a:
a"
°
r
o
*
-'.
o
,:,:
:
,
st
°,
a,
a
a
",
a
a
+ ATTENUATION
30
30
sec
sec
Figure 2.7
6 km
1.5 km
**
;°'.,
°o.
,
.
NIL P
..
I
i _
I
r
''
SJG P
,
.'''~
''
-
,,
" ~''"'
....... /
OGD P
. -
TAB SH
.......
,
,
/
....-
COL SH
S3
sec
sec
Figure 2.8
6 km
3 km
NOR
SJG
NOR
,
SJG
,I "
uIN
WIN
,'
I5 '
SI
5/
~
I
I
SI\y
PEL
PEL
III\
'
V
20
i
l
i
.
I
I
sec
Figure 2.9
i
i
'
58
Vp , km/sec
5
10
*
I
*Ii
I
--10
3 - layer crustII
Purdy and Detrick
_
10
Figure 2.10
0.4
2
R 0.2
0
10
0.4
2
R 0.2
0
10
Centroid Depth, km
Figure 2.11
SEPTEMBER 20, 1969
All stations
0.4
2
R
0.2
0
10
Centroid Depth, km
Figure 2.12
SEPTEMBER 20, 1969
Selected 4 quadrant coverage
0.4
2
R
0.2
0
10
Centroid Depth, km
Figure 2.13
SEPTEMBER 20, 1969
3 quadrant coverage
0.4
2
R
0.2
10
Centroid Depth, km
Figure 2.14
SEPTEMBER 20, 1969
2 quadrant P-wave coverage
0.4
2
R
0.2
10
Centroid Depth, km
Figure 2.15
CHAPTER
3.
RIDGE-AXIS
EARTHQUAKES
NORTH ATLANTIC
IN
THE
OCEAN
INTRODUCTION
In this
chapter we determine the
focal depths and source
mechanisms of 1 6 normal-faulting earthquakes in the North
Atlantic Ocean.
Fourteen earthquakes occurre d along the
an d 2 earth quakes occurred along the
Mid-Atlantic Ri dge,
spreading porti on of the Azores -Gibra ltar
source
mechanis ms of t hese 16
The
plate boundary.
ridge-axis earthquakes are
determined by t he body -waveform inversion technique describ ed
in Chapter 2.
All
eve nts
show fault planes dipping about 450
to the local
and striking pa rall el
trend of the ridge axis.
The water depth s estim ated from the
chapter al 1 occur red beneath
valley.
The ce ntroi d depths
shallow, and
of the
i ndi cate that the ridge earthquakes discussed
P wave-trains
this
ringing part
the inner floor of the median
of these earthquakes are all
ca 1culat e d depth uncertainty ranges
events are generally
s mall.
centroi d
in
very
for these
The similarities in source
depths may be a reflection of the
mechanisms
and
relatively
uniform tec tonic
setting along this ocean ridge.
THE MID-OCEAN RIDGE SYSTEM IN THE NORTH ATLANTIC
The present-day structure of the seafloor in
Atlantic is dominated by
major plates
(Eurasian,
African) that meet
(Figure 3.1)
the Mid-Atlantic
North American,
along this
boundary.
includes the Mid-Atlantic
the north
Ridge and the four
South American
and
The area of study
Ridge from the equator
Ridge
to the Mohns
Gibraltar
Eurasian
sei smic
and
as well
zone
i t has
promi nent
of the Azores-
boundary
between
the
Ridge nearly bisects the North Atlantic
rugged,
(10-18 mm/yr
half rate),
b lock -faulted topography with numerous
fracture z ones .
bathy metr y and gravi ty,
By analyzing the relati onship between
McKenzie and Bowin [1976]
have shown
isos t atic compe nsat ion in the Atlantic Ocean begins at a
wavel engt h of about
fl exu ral
s
(7-13
increasese with increasing
upport of topo graphi c relief produced near the ridge
by an elastic
analy sis
100 km and
They in terp reted this relationship in terms of the
wavel engt h
axis
part
marking the western
T h e ridge is slo wly spreading
basin .
that
western
African plates.
The Mid-Atlantic
and
as the
by
plate about
10 km in thickness.
A later
Cochran [197 9] yie Ided a similar effective thickness
km) for the elastic lith
Mi d -A tla ntic
Ridge except
osphere near most segments of the
the Reykjanes Ridge at 600 -620 N,
for
where the effective thickness must
be in the range 2-6 km.
The
extent to which these estimate s are measures of the thickness of
the zone
of elastic or
brittle
however, is uncertain.
behavior at the rise axis proper,
Macdon ald
[1977
I estimated that the
width of the crustal accretion zone at 37
Ridge is between 1 and 8 km,
a nd
0
N on
the Mid-Atlantic
he sug gested that the crustal
accretio n zone at slow spreadi ng ridges may vary with time and
space.
The
magnetic
anomalies
f rom the Atlantic are generally
not as c lear and sharp as thos e from th e faster spreading East
Pacific
Rise,
perhaps
slow spreading rates
[Macdonald,
1982].
an
indic
at ion tha t
is not continuous in
crustal
accretion
time and space
at
There
numerous
have been
the Mid-Atlantic
Ridge.
valley,
has
areas
however,
at
results
axial
370,
45
crustal
1976].
exist
ma gma
could
7.2
10
0
and
N [Purdy
not
be
detected
the
an 8.1
axis
1985].
The main
evidence
for
1975;
limited
[Fowler,
19781.
crustal
km/sec
Fowler,
(2)
section,
1978;
do
power of
upper mantle,
1976,
[Fowler,
an
chambers
of the
because
experiment
small
three
narrow magma
6-to-7-km-thick
ridge
No
along
the median
in
Ewing,
studies
found [Whitmarsh,
possible that
basa 1 layer and
only
(1)
is
refraction
km of
resolved
chamber was
of a normal,
km/sec
within
well
it
the conventional
formation
60
refraction
structure betneath
studies were that:
However,
but
been
and
0,
of these
The
seismic
The
with
a
occurs
Fowler
and
Keen, 1979].
Purdy and Ewing [1985] and Macdonald [1982]
that most
segments
In
tha t
contrast,
performed
the
of these experiments
on
oceanic
upper mantle
thickness
cooled
are too short
a
recent
a long
crust
state
ridge
appears
km.
at
ma ture,
with
8
km/sec,
oc ean
between episode s of
well-constra ined
23
crust
0
N,
structu re.
Detrick,
and
it
1985]
was
a well-defined
and a total
was
[Fowler,
was
that
found
Moho,
crustal
interpreted to
major magmatic
refraction studies
spreading
crustal
and
[Purdy
se gment
of ~
The
conducted on
resolve deep
experi ment
velocities
of 6-7
to
were
pointed out
in a
be
Two
injection.
1976;
Bunch
an d
Kennett, 1980] suggest the p resence of low-velocity mantle
beneath the Mid-A tl antic Ridge median valley.
material
no strong evidence exists to
mantle
is
suggest that
How ever,
low-velocity uppe
a ubiquitous characteristic of axial
r
valley segments
of
slow
spreading
Two
ridges
[Purdy
ear thquakes
of the
and
Ewing,
studied
in
along the Azores sp reading center.
1985].
this
chapter
The Azores
occurred
region
is
the
site of the triple junction between the North American,
Eurasian, and Afric an plates.
East
of the Azores,
the
Eurasia-Afri ca plat e boundary is dominantly a transform fault
called the Gloria f ault [Laughton et al.,
Whitmarsh, 1974].
At
its
western
end,
1972;
Laughton
the Eurasia-Africa
and
plate
boundary is a sprea ding center with a very slow rate of opening
[Krause and Watkins , 1970; McKenzie, 1972; Searle, 1980].
There
are several hypothe ses concerning the evolution of the Azores
spreading center.
(1)
The
triple junction
began
as
a
ridge-fault- fault ( RFF) junction and changed to the present
ridge-ridge- ridge
( RRR) junction following a change of spreading
direction from E-W to ESE-WNW [Krause and Watkins, 1970].
(2)
The present triple junction (RRR) has been stable for an
extended per i od [McKenzie,
1972].
(3
) Short segme nts of the
Mid-Atlantic Ridge are displaced to t ie east and a re connected
to the main
et
al.,
197
to Europe
r idge by several ESE-WNW leaky transfo rms [Machado
2] .
(4)
c ha nged
The
several
to WSW duri ng the last
this
view,
around
36 M.y.
direction.
RFF
the
during
gradually,
In
most
relative moti
72
times
M.y.
from
on of Africa with respect
ENE to WSW to SW then back
[Searl e, 1980].
Ac cording to
!
Azores spreadi ng cente r was probabl y initiated
B.P. when the
rel ative motion last
this scenario,
of
moved to
its history
the north.
changed
th e tri )le junction was probably
and then suddenly,
rather than
The Azores plateau appears to have been formed
Azores
Using
the
hot
spot
Rayleigh
crust
of
[Schilling,
wave
the Azores
anomalously
place
in
layer
basaltic
lava.
N120 0 E trend,
Recent
at
Terceira
Rift
to the
Rift consists
The
basins
N142
0
-154
0
faults
[Searle,
with
On
ridge
the other
it may
basins were
even
formed
When the spreading
axial
like
volcanic
hand,
developed
[Searle,
normal
e y.
first
u des.
Sl eep
It is not clear whether the
s uggested
and
a hot
ri
center evol
high.
th at a very
should also
have
a median
Rosendahl
[1980]1
argued
sp ot
It
may not
is
have
a median
poss ible
that the
fting of oceanic
ved to hot spot
flanking the
1980].
fault scarps trending
t he basins or ridges or both
axial
by
The Terceira
The ridges also have normal
[ 1973]
near
1980].
and ridges [Machado, 1959].
the Azores
have
ridges
u lt [Searle,
ections unlike the V-shaped
cross
Lachenbruch
region show a
run from the Mid-Atlantic Ridge
ba sins
of
through
a spreading center
valley;
fa
smaller amplit
center runs
slow-spreading
that
square
in the Azores
Terceira, and then down the
to
Ridge median vall
1980].
valley.
a series
uppermost mantle velocities
ivity suggests that the
center may
Gloria
thicker than the
of the tectonic lineaments.
well-developed
E, and
but
spreading
of
have
Mid-Atlantic
d ata
volcanic ac t
due east
60%
found that
to increased production of
similar to that
almost
[1976]
1977].
the thickening seems to have taken
due
magnetic
spreading
38 0 40'N,
of
probably
seismic and
present-day
that
and
Most
The
Searle
Platea u is about
crust
low.
2,
Schilling et al.,
dispersion data,
surrounding oceanic
are
19 75;
by the
rifted
litho sphere.
ridge structure,
basins may
have
EARTHQUAKE DATA SET
We began our study with a search for earthquakes on the
Mid-Atlantic
Ridge
and S waves
at
large enough to yield clear long-period
teleseismic
within
the
search
included
median
valley of
of the
the years
1964-1982,
and
International
earthquakes
Udias
et
1981;
Dziewonski
al.,
Service
[Sykes,
1967,
1970
et al.,
1 979;
].
1983a,b
for 1962-1965,
[1969]
( NEIS )
p lane
Einarsson,
The
1 Centre (ISC)
th e
for
lis tings o f the National
fault
1976;
segment.
ridge
Seismologica
of
studies
Ridge
R othe
the monthly
Information
published
nces and ha ving epicenters
an a ctiv e
the catalog of
bulletins
Earthquake
dista
for
t
soluti
he years 1982-1983
ons for Mid-Atlant ic
Solom on
Savost
and Julian,
197 4;
in and Karasik,
lar care was taken
Particu
to exclude
near-ridge
intraplate
eve nt s [Bergman and Solomon,
1984].
restricted
the
to the
We
Ridge,
good
between
latitudes
azimuthal
The
16
search
00
distribution
selected
and
800N
of t eleseismi
Figure
are shown in
the table,
we have found at
3.1.
Seismograph
800.
The events
(GDSN)
c stations.
Table
For all
at
Network
at
epicentral
or
th e Global
have epicenters within me dian valley segments
series [Laughton and Monahan, 1978; Heezen
between
a nd 72.2
range between latitudes 0.40
Chart
in
of the Wor Id
distance s
maps of the GEBCO (Genera 1 Bathymetric
their
d SH waves
statio
(WWSSN)
3.1 ;
the events
least 10 clear P
recorded on long-period i nstruments
Seismic Network
, Primarily to ensure a
earthquakes are listed i
epicenters
Wide Standard
northern Mid-Atlantic
Di gital
and
300
0
N a nd
o n bathymetr ic
of the Oceans
and
Tharp,
1978;
)
et
Johnson
al.,
half-spreading
the
plate
as
rates
vary
et
al.,
from
1982].
0.8 to
characterized
exhaustive,
by
below,
normal
given
our
all
have
faulting.
selection
The local
1.8 mm/yr
of Minster and
have body-wave magnitudes
demonstrated
through
Searle
kinematic model
earthquakes
and,
1979;
Jordan
[1978].
mb between
focal
We
according to
5.1
and
5.9,
mechanisms
believe that
criteria,
The
for
the
this
list
period
is
1962
mid-1983.
WAVEFORM INVERSION RESULTS
In
from
this
the
section,
inversion of
earthquakes.
seismic
convention
Aki
and
water
of
the
crust.
in Tabl
3.2.
uple o rien tat ion
that
depths are given
lative
All
wav eform
structure:
km/s
Vs = 3 .7
and
Vs =
km/s,
4.6
we also
to
6
F
examined
structure on inversion solutions;
the discussion of this earthquake.
km an d seismi c( velocities
mantle
and
km/s.
ringing portion of the
er of variable thickness, a
lay
layer o f thickness
and
of
ons we re perfo rmed with the
inversi
a water
The
in Tab le 3.2 i s the average
listed
in the
region earthquake,
in
on,
seismograms
crustal
8.1
ori ent a
and observed
Vp 6.4 km/s
Vp =
dge-axis
obtain ed by matching the
source
single
Also
16
hypocenter,
above
P wave train.
same
Centroid
the
obtained
each
depth
synthetic
[1980 ].
ple
are gi ven
depth
describ ing double-co
Richards
the top
P and SH waveforms for
and ce ntroid
for
the source parameters
best -fitting doub 1 e-cou
The
moment,
we present
the
he
half-
July 4,
influence
.(
space with
1966, Azores
of uncertainty
details are given below in
June
3,
1962
The earthquake
Ridge median
(Figure
many
solution
(Figure
The centroid
3.9
depth
the depth
predates
coverage
is
is
1.2
km
less
represents
is
in
the
the
Mid-Atlantic
Fracture Zone
establishment
rather
nearly
pure
of 6.1
The
dip-slip
x 1024
dyn-cm.
quite well-constrained.
from the predominant
in the P waveforms
is
of
poor and
than ideal.
a seismic moment
e stimated
reverberations
occurred
south of the Kane
mechanism
with
of
km water depth
water
km
( 4/43/274)
3.3)
1962
event
station
of the fo cal
inversion
3,
100
Becau se this
stations,
resolution
motion
valley about
3.2).
WWSSN
of June
period
The
slightly
of
greater than
of the med ian valley inner floor at the nominal
latitude of the epicenter.
Beca use of the poor station coverage,
there is
trade-off between t he centroid depth and the focal
significant
mechanism.
in
t
he residual-depth curve occurs at 6 km
(imediately
below
t
he Hoho discontinuity).
(-4/48/258)
for the
second
minimum
component
than
seismic moment
solution.
about
2.4
that
of
sec
in
du
solution.
For both
well-fit.
Using
th
centroid depth s to
level a of 0.1
The focal mechanism
mantle solution has a larger strike-slip
of the best solution at
6 .1 x 1024
The sour
A
1.2 km,
but the
dyn-cm is the same as the best
ce time function for the mantle solution is
ration compared with 3 sec for the best
solutions, all observed waveforms are
e t-test,
we estimate the range of acceptable
be from 1-2 km to 6-7 km for a significance
(App endix B).
To simulate a more graduated Moho
transition and to test the stability of the
mantle
solution,
we
repeated
2 km
the
inversion with
of crust
has a P wave
the mantle solution
but
it
is
Since the
power
still
(1-B, see
solutions
Chapter
within
for
the
Appendix
test
in
velocity
degraded
this
A)
of
sampling of
the
focal
June 2,
1965
of
with
in
7.1
such
solutions
small
(~
case cannot
modeled
range
are
basis
the
found that
structure,
with
a = 0.1.
implying
0.1),
lowest
similar, the
distinguish
On the
that
We
a source
two
is
which the
km/s.
uncertainty
believe
inadequately
model
the
without more data.
2, however, we
artifact
June
well
residuals
statistical
is
a source
test
that
these two
of the analysis
in
deeper minimum is
crustal
the
structure
and
an
poor
sphere.
2, 1965
The
valley
3.4),
about
has
50 km
been
the
earthquake,
north
of the
subject of
which
15020
'
several
occurred
Fracture
in
the
Zone
median
(Figure
previous studie s.
On the
basis of P-wave first motion polarities, Sykes [1970] found a
norma 1 faulting mechanism with non-orthogonal
Tsai
[19691
estimated the
seismic moment
nodal
to be
2.8
p lanes.
x 1025 dyn-cm
from the Rayleigh wave amplitude spectrum at one stat ion,
but
this value was based on an adopted focal
Using
depth
of 65
km.
both the amplitude and phase spectra of Rayleigh wave s from a
numbe r of stations, Weidner and Aki [1973] demonstrat ed that the
focus is actually quite shallow (3 + 2 km).
a bes t-fitting
double
couple
(0/40/260)
and
7.8 x 1024 dyn-cm.
From the
period
Duschenes and Solomon
focal
P waveforms
depth of 2 + 1 km.
They
als o obtained
a seismic
identification of
(1977)
pP
moment
on short-
estimated a
of
invers ion solut ion
The body-wave
3.5.
Figure
The SH waves
(335/38/244) is
provide exceptional coverage, and all
of the observed waveforms are well1 fit.
nodal
plane has
local
trend of the
motion is
A small
ridge axis.
ord er to
in
fit
of the direct SH phase
momen t
3.0 km.
Both
Aki
[19731].
3.4
km.
is
= 0.1
is
still
1-5
is
immediately
very
3 km depth.
mantle
solu tion
smaller.
to
centroid depth
is
ose
and
e
of Weidner
P waveforms
ce ntroid depth
6 km centroid
quite
at
mantle
solu
the
first
half
depth,
solut ion (331/40/237)
tion; how eve r, some
th e s olution at
cycl e of the synthetic
SH wa veform
at SHA are too
becaus e the predicted am plitud es are
When
is
t he water reverbe ration s is worse for the
somewhat
al grounds, both s ol utio ns are quite
a more gradual
the velocit y structure
deteriorates
This
of t he synthetic
0 n statistic
acceptable.
dilatation
best-fitting
a sli ghtly worse fit th an for
the fit
Also
at
b est-fitting
the
For examp le
wide.
small
6-8 km.
below the Moho.
L PB and
of strike-sli
coverage for th is equation is
solut ion occurs
seismograms give
P wave at
with
range in acceptable
km and
sim ilar to
to th
some trade-off between source mech anism and
The
A seco ndary
is
Th
8.3 x 1024 dyn-cm, and th
values agree quite well
centroid de pth
a
th e
The water depth inferred from
there
good,
component
where
is clearly see n.
Althou gh the azi muthal
parallel
the SH waveforms to the
northeast, pa rticularly GDH and KTG,
seismic
The westward-dipping
a strike slightly east of north,
required
shown in
and can be
Moho tr ansiti
on is assumed for
at the source, the mantle
reject ed at a = 0.1.
solution
Th e estimated
depth
uncertainty
for
study
is
as that
of Weidner and
determine d
by Duschenes
the same
than the
1 km
±
have to
remember that
and
necessarily
not
the
our
the
short-period
waveforms.
November
1965
16,
The
earthquake
of
preferred
shallow
of
and
rupture
No vember
16,
in
[1973]
but
Aki
depth estimate
p oint
solution
Solomon
our
greater
[1977].
is
the centroid
as
seen
using
1965 occurred
We
depth
the
in the median
valley about 100 km north of the Atlan tis Fract ure Zone (Figure
3.6).
From P-wave first motions,
faulting mechanism with
estimated the
[1969]
the Rayleigh
wave
focal
depth
a focal
depth
of
3
of Rayleigh waves.
couple of
From an
and
much
of
to
spectrum
km.
They
also
nodal
be
at
Weidner and
of
of the short-period
obtained
a focal
inversion
solution
(Figure 3.7)
closer to the
and
Aki
trend
[1973].
of
the
The distribution
Aki
1.2
the doubl e couple.
centroid
depth
is
at MDS
[1973] estimated
and phase
spectra
x
1025 dyn-cm.
4 ±
shows
2 km.
nearly
plane wh ose
axis
than
of
Duschenes
pure
strike
is
the mechanism
P and
In
of
SH waves
particular,
and LPB constra in the strike of
The seismic moment
2.1
from
on an
based
pP p hase,
provides a good sampling of the focal spheres.
the nodal SH waveforms
1025 dyn-cm
sta tion,
depth of
ridge
Tsai
a best-f itting double
a seismic moment
(25/46/265) on a fault
a normal-
pl anes.
amplitude
reported
found
5.2 x
one
[1977]
faulting
Weidner
45
moment
2 km using both
±
identification
The
normal
seismic
0/59/234 and
Solomon
non-orthogonal
amplitude
assumed
Sykes [1967]
is 8.1
km and well-constrained
1024
dyn-cm,
and the depth
the
is
uncertainty
A
km.
secondary solution
depth
is
outside the
narrowness
synthetic
at
only +
KON,
located just below the Moho,
rang e
MAL,
u ncertainty
depth
seismograms;
and WIN
of acceptable
fo r
are
3 km depth,
too large
results
of
mentioned
4,
and
surface-wa ve
above,
indicati ng
and
can
the
is
first
waves
P-wave
The
seen
on
motions
do
comparable
short-period
agreement
SH
= 0.1.
readily
P-wave
most
is
but
at a
be
2 km.
not
fit
to the
studies
among these methods.
1966
This
within
the
values
range
Our estimated dept h uncertainty
well.
July
of the
The inf erred water depth is
earthquake occu rred
a small
depression at the
spreading center
Banghar
and
faulting
Sykes
with
(112/70/220).
ightly to th e east of Sao Miguel
sl
(Figure
3.8).
[1969]
found
a large str ike-sli
McKenzie [1972]
P-wave first motions
eastern end of the Azores
Using P-wave first motions,
a
mechanism i nvolving normal
p component of motion
restudied thi s earthquake with
and found a similar sol ution (103/70/230).
Grimison and Chen [1985] used both P and SH waves to infer a
pure normal-faulting mechanism ( 126/56/270)
and suggested that
oblique spreading is occurring at the Azores spreading center.
They estimated the centroid dept h to be abou t 12 ± 5 km and the
seismic moment to he 4.0 x 1024 dyn-cm.
Because the depth for this and another nea rby
(April
20,
1968)
obtained by Gri mison and Chen [198
rthquake
are
significantly deeper than the maximum centroid depth of about
3 km that we have obtained for Mid-Atlantic
decided to examine these two earthquakes
Ridge events , we
in detail
as pa rt
of
our study.
solution (135/49/274)
Our inversio
similar to that o
solution is
shown
Grimison and Chen [1985].
very close to the
spreading center.
The
overall
in
3.9
Figure
The strike
trend
of the
is
of our
Azores
cen t roi d depth is 5.4 km, and the seismic
moment is 4.0 x 1024 dy n-c m.
The difference
between our study and t hat
Grimison
in
depth
centroid
and Chen
[1985]
is
significant since both studies used the inversion routine of
Nabelek
[1984].
discrepancy.
Several factors m ay have contributed to this
Fir
s
we used many
t,
more sta tions
(25)
much better azimu t hal coverage for
both
Grimison and Chen [1985] only used
3 SH wa veforms
P and
almost no constrai nts on source mechanism.
sec.
Grimison and
5 sec in du ration,
Chen [1985]
trade-off between the focal
such short duratio n.
have
SH waves.
We
which
used
and have very good constraints on the noda 1 planes.
source-time functi on is
and
while
10
gave
SH waves
Second,
theirs
was
our
2
arg ued that there is little
depth
and a so urce time
function
of
Third, diffe rent vel ocity structures were
used in the invers ion.
We used a generali zed oceanic crust-
mantle structure while Grimison and Chen [1985] used a mantle
halfspace.
To estimate the depth uncertainty, we followed the
procedures of Chapter 2.
event is quite good, the
centroid depth
Since the azimuthal
source mechanism is
fixed in the range 1 to 20 km.
acceptable centroid depths
in
is 4-5 km.
residual-versus-depth occurs
at
coverage for this
stable for
a
The range of
A secondary local
minimum
10.8 km but may be rejected
at a = 1.
well,
For the
but
the
This
10.8-km deep
SH waves
at
NOR,
solution,
SJG,
BEC,
the P waveforms
and SCP
are not
fit
well
fit.
event occurred in an area that may be anomalous
rel ative t o conventional oceani c ridges, both because of the
ext remely
hot
low rate of ext:ension and the proximity to the Azores
spot.
We therefore
repeate d our inversion with alternative
sou rce vel ocity structur e_s.
(Vp
= 8.0 km/s,
Vs = 4.6 km/s)
and Chen [19851
and
secondary local
min imum
occurred
at
11.2 km ,
very clos e to the centroid depth inferred by
For
km.
a source structure approximated as
Fin a lly,
two-layer crust ove r a mantle
4.0,
4.4
km/s ; layer
Vs = 3.7 km/s)
residua 1 among
we obtained a
we used a model with a
h alfspace (Vp = 4.5, 6.9,
7.6 km/s;
thicknesses = 3.8 and 4.2 km,
respectively), f ol iowing Searl e [1976],
depth of 6.3 km.
A
in res i dual-versus-centroid depth
halfspace (Vp = 6.4 km/s,
centroid depth of 3 .7
Vs = 2.6,
identical to that used by Grimison
obtained a centroid depth of 6.8 km.
Grimison and Chen [ 19851.
a crustal
Fi rst, we used a mantle halfspace
and obtained a centroid
This solution gives the smallest overall
the best-fitting solutions for all of the source
structu res we have examined.
A comparison of inversion
structures is
results using different source
given in Table 4.3.
Secondary local
minima in the
residual-versus-depth curves are also given in this table.
we can see, the
stable with
best-fitting double-couple orientation
respect to changes
As
is quite
in the assumed source structure.
However, the centroid depth varies from 3.7 to 6.8 km depending
on the source
structure used.
We consider the 3 km difference
to
be
the
small
ve ry
length
in view of the shallowness
different
of this
source structures used
in
earthquake
the tests.
and
The
of the source time function ranges from 4 to 7.5 sec,
slightly larg er than f or Mid-Atl ant ic
Ridge earthquakes.
All of
the secondary local mi nima are d eep er and have much shorter
source time f unctions.
From thi s analysi s,
beli
we
eve that this event is deeper
than typical Mid-Atlan tic Ridge earthquakes and has
l
duration of r upture.
Because th
!
reverberation s are muc h smaller
believe that
and did not
a longer
e amplitudes of the water
than the P wave amp litudes, we
earthquak e rupture was
extend to the seaflo or.
confined within
the crust
Comparisons of observed and
synth etic seismograms for
inversion results with different
e structures, as well
as the secondary solutions, are
s ou rc
given in Appendix C.
April
20,
1968
The April
20,
19 68 earthquake occurred about
southeast of Terceira (Figure
3.8).
studied this
P and
event
using
faulting mechani
both
(302/54/270).
15 + 5 km
and th
seismic moment
inversion
soluti
(328/34/293),
slightly
greater
is 2.0 x 1024 dyn-cm.
centroid depth
event.
a re
Because th is
coverage is
not
very
SH waves
They
to
be
shown
com ponent
Th e centroid depth
Chen [1985].
moment
trike-slip
Grimison
is
Bo th
similar
earthquake
as good as that
4.6
the
to thos
is
very
for the
and
50
km
Chen
and found
[1985]
a normal
estimated the depth to be
4.0 x 1024 dyn-cm.
in
Fi gure
than
km,
focal
3.10,
Our
has
a
that of Grimison and
and the
seismic
mechanism and
e of the June 4,
smal 1,
1966
the station
earl ier earthquake.
The
range
in centroid depth
estimated
A secondary
solution occurs
duration of
1 sec.
quite well
except
earthquake
is
reliable
the
3 and
larger
July
3 and
20,
September
character both
north
result
typical
P-wave
20,
3.11),
of
Azores
about
ridge
focal
1973;
influence
of other
event;
the
this
can expect
inversion.
is
At
depths
similar to
centroid depth
Vogt
and
if
the
is
and median
lies
heterogeneous model
axis,
a focal
found
(9/41/263).
as
of the
mechanism
Hart
a forward
and
spot;
Julian
of
orthogonal
[1978] modeled
problem.
He
nodal
the
argued
a
presumably
a
to the south
Ridge.
data
Using
obtained
planes.
onto
a
They
the
a laterally
beneath
planes
the
with
ridge
could be
P waveforms
that
of
are more
through
structure
the
lacks
nodal
first-motion
velocity
To
[1974]
nonorthogonal
the
valley
shallow,
Mid-Atlantic
and
in
1975].
valley morphology
for propagation
with
axial
aseismic,
Iceland hot
the Reykjanes
a transition
Johnson,
nearly
Solomon
projection
of
anomalously
of the
sphere were corrected
south
and
is
of the
data,
occurred on
seismicity
ridge
sections
first-motion
earthquake
size of
for which we
this earthquake
100 km
axis
valley,
seismicity
that
small
body waveform
1969 earthquake
normal-faulting mechanism with
showed
The
function
the waveforms
acceptable centroid
We think
590 N the
median
of the
59 0 N the
in
fits
km.
6 km.
[Francis,
of about
prominent
range
3-16
a source-time
TRN.
limiting size
is
1969
(Figure
morphology
the
with
solution
P wave at
long-period
1966,
The September
Ridge
the
6 km.
4,
km,
9-km deep
probably the
a = 0.15,
between
between
for
results from
the level
is
The
at 9
using the t-test
from
this
a shallow
focal
(2.1 + 0.2 km)
depth
apparently
anomalous
first
double-couple source
2 x
the waveforms,
motions,
can
He
(9/35/260).
including those with
be matched
with
obtained a moment
a
of
1025 dyn-cm.
The observed
ratios
and
P and
SH waves
are well-distributed,
problem is unusually
show good
signal-to-noise
so the waveform
inversion
well-constrained (Figure 3.12).
component at NAT and both horizontal
The E-W
components at NIL were
to have reversed polarity and were corrected prio r to
found
The
inversion.
inversi on sol
ion
in strike from that rep orted
(23/42/26 1)
Hart [1978],
bu
ffe
slightly
the
waveforms
alone allow onl y loose constraints to be placed on the fault
strike; the P waveforms in Figure 11 are exceptional ly well
by the syntheti c seismograms.
We obtained a moment
fit
of
1.5 x 1025 dyn- cm and a centroid depth of 1.5 km,
both
smaller than th e values reported by Hart [1978].
Th e indi cated
2.2 km,
water depth
central
media
3.11).
Becau se
consistent with an epicenter
for
The beh vior
in
Chapter
2,
of good azimuthal coverage,
May
the sour ce
-
The 90% co nfid ence
ce ntroid depth as gi ven by the t-test is 1-3
of
km.
a mantle solution has been extensively studied
where
we suggested that
artifact of the sharp Moho assumed in
in
the
in
n valley of this sha llow ridge system ( Fi gure
mechani m-ver su s-depth trade-off is small.
interva
sli ghtly
the mantle
solution
is
the source structure used
the inversion.
31,
1971
The May 31,
1971 earthquake
an
occurred north of Iceland
on
the Mohns Ridge
(Figure
normal-faulting
mechanism with n onorthogonal
P-wave
first motions.
3.13).
Conant [1972]
and Karasik
Savostin
obtained a
nodal
planes from
[1981] used
a larger
data set to constrain the focal mechanism to a combination
normal
and strike-slip
however, does
not
faulting
(141/59/220).
fit the observ ed
Such
of
a mechanism,
P and SH waveforms.
From
waveform inversion we obtained a normal-faulting mechanism
(Figure 3.14) with only a small
strike-slip
component
(51/51/284).
We have excellent azimuthal
and
and most observed waveforms are well
SH waves,
seismic
2.3
km
moment
is
6.8 x
102
4
and well-constrained
(Table
Two slightly different
water-column
the east
water depth
AAM,
2.8
DUG,
km.
KBL,
of 3.1
and COL)
We
MSH,
km,
interpret
northwestern
location
about
15
3.13).
km
whil
is 2 -3
has a focal
AOU
depth
3.2).
depths
were
indi cated
the P wavetrains.
and
TOL)
are
by t he
Stations
con sistent
e stations to the west
t0
wi th
(TRN,
WES,
a somewhat shallower wa ter depth
of
0
north
0
f
the
median
valley
in ner floor,
a
o r northwest of the ISC epicenter
azimuthal
T he
range
coverage,
in
A secondary
mechan ism (53/50 /290)
mantle solu tion
focal
wall
6-9 km.
best-fittin g solut ion
the
The
p attern as indicating a n epicente r
Due to good
and
and
fit.
P
this
mechanism i s very sta ble.
depths
in
IS T ,
suggest
near the
(Figure
water
reverberations
(SHL,
dyn-cm,
coverage for both
at 2.3 km.
function is 3 sec in duration.
acceptable
solution
similar
at
to that
The seismic
is 6. 8 x 102 4 dyn-cm,
th e source
centroid
8.2
km de pth
of the
mome nt
for
the
and the sour ce time
When we use a crustal structure
a
with
a more
gradual
deteriorates
April
Moho transition
and can
be
rejected
the mantle
for a
solution
= 0.1.
3, 1972
The April
3,
1972 earthquake
(F igu re 3.11).
Zone
normal-faulting mechan ism
P-wave first
the southern
150 km north of its intersection with the
Reykjanes Ridge, about
Gibbs Fracture
occurred on
motions.
with
Einarsson
nonorthogonal
[1979] obtained
nodal
planes
a
from
Fr om a moment-tensor inversion of the
Rayleigh-wave radiatio n p attern and forward modeling of
long-period
Trehu
P waveforms,
normal-faulting
mechanism
et
al.
(4/46/270)
and
apparent
nonorthogonality
is
source.
They matched
P waveforms
length
13
rupture
the
km,
width
a good
significant
TUC.
The
t he center
focal
match
to
discrepancies
The inversion
mechanism
the
solut ion
strike-slip
motion
waveforms.
The centroid
the
suggested
with
7.5
et
of Trehu
depth
of
but
al.
of
with
extending to
[1981]
there are
at
AAE,
JER,
and
slightly more
matches
1.9 km
the
source
and
SH waveforms
) and
of
the
x 1024 dyn-cm,
(8/48/254) has
(Figur
that
a finite
of the fault
P waveforms,
with
a pure
the shallowness
and moment
3 km,
near
initiating
seafloor.
provides
the
due to
found
[1981]
P and
all
obtained
from
SH
the
inversion is somewhat deeper than the 1.1 km indicated by the
finite fault
of 4.4 x 1024
model
of Trehu et
dyn-cm given by
al.
[1981].
The
seismic moment
the body-waveform inversion is
smaller than that obtained by Trehu et al.
[1981]
from the
forward model ing of P waveforms and from the moment-tensor
inversion of Rayleigh waves.
The inferred water depth
of 2.8 km
is consistent with an epicenter in
median valley inner floor
(Figure
a locally deep portion of the
3.11).
ainty for this earthqu ake is small,
The dept h unce
with
the range in accept
le values between 1.5 and 3 km,
focal mechani sm is
ry well constrained by SH waves.
secondary min imum i
residual-versus-depth occurs just below the
Moho at 6 km beneat
the sea floor.
mantle soluti on is
ry similar to the best
rejected at the a =
.1 level.
solution is s lightl
better than the mantle
station, thus making ad small
because the
A
Despit e the fact that the
This is
solution,
it
may
be
bec ause the best-fitting
solution at every
and the statistic
t
large.
June 6, 1972
The June 6, 1972 earthquake occurred in the Mid-Atlantic
Ridge median
about
valley
Zone (Figure 3.6).
100
Station
km
coverage fo
mostly confin ed to the northern
solution
(20/ 51/253)
with
(Figure 3.16)
represents
a small
south
The
depth
is
sei smic moment
1.8
km,
and the
4.1
predoni
nantly normal
x
l
componen t.
1024
indicated
wate
SH waves
1 to 7 km.
at
6 km below the seafloor.
seismic momen t
A secondary solution
(3.5
x
1024
occ
The focal
dyn-cm)
the best-fitt ing solution at
are
1.8 km dep
examination, we noticed that the first
faulting
The strike
at KTG,
KEV,
dyn-cm,
the centroid
r depth
is
The rang e in centroid depth at a = 0.1
is
Th e inversion
ere.
strike-slip
is
r both P and SH waves is
hemisph
the fault pla nes is constrained by the
CAR.
of the Hayes Fracture
for
urs just
of
and
3.2 km.
this earthquake
below the Moho
mechanis m (25/50/260) and
very sim ilar to those of
th.
Upo n close
half
cycles
of most
synthetic
SH
compared with
waves
mantle
for the
solution
the observe d seismograms.
probability v alue P (see
Appendix A)
make a type I error only
3 out
solution can
transition
mantle soluti on
June 28,
At
is
too
0.15;
20 times.
is
not
co rrect,
in centroid
We
inversion.
depth
and
is
that
is,
Also, the
th
we
hould
man
le
Moho
ther efore beli
a more
1 to
wide when
6 km depth,
be rejected at a = 0.1 when a gr adual
is used in the
of the range
of
are
eve the
re asonable statement
6 km.
1977
On June
swarm occurred near the
1977 an earthquake
28,
epicenter of the June 3, 1962 earthquake discussed above
(Fi
fir
re 3.2).
The ISC lists 5 events ove r a 4-hou r pe ri od.
in the series (1538:37.8)
had mb
5.3,
Ms
= 5.6;
fou
h and largest in the series (1918 :36 ) had mb = 5.9,
Ms
6.0.
The
the
The inversion solution for the largest earthquake of the
swarm (Figure 3.17)
represen ts nearly pure normal
(1/44/255), a solution very similar to that
earthquake.
solution.
All
of
the obs
e rved waveforms
Th e seismic mome nt
centroid dep t h is
1.6
is
1.1
1025
x
P -wave coverage
km.
quadrants, an d SH-wave cove rage i s well
nodal planes
so
resolution
of the
focal
faulting
of the June 3, 1962
are well-fit by th e
dyn-cm, and the
spans about 3
constrained about t he
mechanism is
very good.
A secondary s olution occurs just be neath the Moho, but this
solution can
be
estimte that
the
significance
level.
rejected
range
in
This
at a
= 0.1
.
Using the t-t
centroi d depth
e stimate
is
est we
1 to 2 km at
is probably accu rate,
th e 0.1
be cause
the probabil it y of
to
failing
B = 0 .1 ) f or depths
(i.e.,
is 3. 5 km,
water depth
mislocated 6 -7
bes t
The
3.18 ) is
detect the
outside this
indicating that
range.
the
inferred
The
ISC
only 0.1
epice nter
is
km to the west (Figure 3.2).
ove rall
solution
for the
characterized by normal
first
A secondary
solution
swarm
faulti g with
stri ke-s lip component (355/50/241) and
5.5 km.
difference is
e vent
(Figure
a large
centroid depth of
occurs at
2.5 km depth and has a
foca l me chanism (1/46/254) with a small e r strike-slip component
that
all
is nearly identical to that of the large event.
Because
of t he stations for this event are 1 ocated in the northern
hemi sphe re,
n 0 t well-constrained.
the inversion solution is
we const rained the slip angle to be lar g er than 2500,
shal lowe r solution has the smallest ove r all
residual.
If
then the
We prefer
the shallow solut ion because its mechani sm is similar to those
of other
large ea
rthquakes (with better station coverage)
and
microearthquakes
in this area.
short-period
-waveforms for this eve nt indicates that the
shallow
The
body
solution
shallower sol
seismic moment
depth
is
4.0
10
km east
preferable
is
[Bergman
3.0 x 1024 dyn-cm,
and Solomon, 1985b].
and
The
the inferred water
greater than that of the larger event but also
an
of the
epicenter in the cent ral median valley about
ISC location (Figure 3 .2).
The P wavefo rms
mb = 5.5,
a preliminary analysis of
ution (1/46/254) is sho wn in Figure 2-17.
km,
consistent with
is
Also,
Ms = 5.7)
of another event in this swarm (1618:16,
are obscured by
earthquake 40 minutes before.
the surface waves from the
A normal
faulting mechanism
(0/45/255),
with
a seismic moment of 5 x
centroid depth of 1.5
this
km,
th e observed SH waveforms for
fits
earthquake quite well.
January 28,
1979
The January
28,
1979 earthqu ake occurred in
intersection with a small
valley near its
100 km north of the Vema Fracture Zone
mechanism of this earthquake,
pure normal
6.3 x
The
fracture zone
(Figure
as determined
3.4).
in the
about
The
planes di ffers
N10 0 E rather
water depth of 3.7 km is
Figure
about
The SH waves are
is
but the station coverage for SH wlaves
a mechanism striking
by
100
0
E.
is
of
3.19).
from
the
well-fit,
reasonably
inadequate
N20
than
focal
inversion,
1024 dyn-cm and a centroid depth of 2.2 km (Figure
strike of the nodal
to
rule
out
inferred
The
consiste nt with the epicenter shown in
3.4.
Becau se the
some trade -off
Using the
from
fault
t-test,
1. 5 to
we
range
indi cates,
residual -v ersus-depth
the
solution
only 3 out
3-km-deep
solution
is
about
of
20
a 0.1
in
centroid
3 km
differs
0.15;
times.
solution;
that
well-constrained,
depth
range of
depth
reasons.
is,
At
an a
local
depth,
level
we have
of
about
However,
minima
2 and
the
from
would
depths
t 0
we
smaller than this
sharp between
we
there is
mechanism.
acceptable
The
centroid
that
focal
level.
is
significantly
is,
and
significance
curve are
At
not
centroid
for several
7 km.
are
estimate the
uncertainty
6 and
the
9 km at
the
between
planes
between
think
which
the median
faulting (20/46/270), with a seismic moment
strike of the ridge axis.
be
a
1024 dyn-cm and
in
the
3 km and
P value
at
the best-fitting
make
a type
0.1,
B = 0.6
I
error
for the
a 3-in-5 chance
of
being
unable to
detect
The
3-km-deep solutions.
chance of
more
deteriorates
the
Moho
for
waveforms
centroid
km and
3.3
km,
as
is
very
The
centroid
10
km.
shown
improves
our
solution
range at
our best
a = 0.15.
estimate
of
2 to 3 km.
for
f
in
seismi
to those
dept h and
Using
the
t
t-test,
depth
of 0.1
(8.3 x 102
to those
solution deteriorates,
the observed
it by a no rmal-faulti ng
Figure 3.2
c moment
The centroid depth
The indicated
is
9.9 x
the
of the June
for this
chanism
6,
1972
event
which
is
gives
water depth
1024 dyn-cm.
centroid
is
The
and water
depth S
earthquake.
fairly
very
good,
good
constraints
e re is little trade-off between
mechanism for the depth range 1 to
focal
ce ntroid
similar
Th
Station
waves is excellent, an
SH wa ves
the
planes.
fault
(Figure 3. 6).
1972 event
stat ion coverage
momen
1979 occurred very near the
ril 22,
orientation and both
significance level
very
the mantle
well-c 0 nstrained.
very
simi lar
acceptable
seismic
Ap
well
an d the
part i cul a rly
for the
on
P and S H
very
double-couple
are
is
and the
Also, when we use a
considerations,
depth
June 6,
both
are
(17/52/262),
is
= 0.5).
inversion,
statist ical
epicenter of the
1.8
(8
slightly
outside the uncertainty
eart hquake
coverage
= 0.15
the best
1979
22,
The
t he
in
and lies
above
range in
April
use of a
detecting a difference
gradual
From the
the difference between
we
estimate that
is
between
The
1.5-2.5
focal
dyn-cm)
for
the
range
km and
mechanism
of the best solution.
however, and
the
in
6-7
km at
a
(13/54/255) and
mantle
solution are
This mantle
lies outside the uncertainty
inversion.
Based
centroid
correct
June
a gradual
at a = 0.15 when
range
28,
on the
depth
Moho
structure is
above analysis,
li es between
1.5
used
in the
we believe that the
and
2.5
km.
1979
The June 28,
1979 ea
occurred in the median valley
rthquake
about 30 km south of its eastern intersection with the St.
Station coverage for
Paul's Fracture
Zone (Figure
and SH waves
concentrated in the northern hemisphere.
focal
waves
pure
are
km
nor
and
fit by this mechanism,
a centroid
a se
ismic moment of 3.6 x 1024 dyn-cm.
shown in Figure 3.21.
mantle
level
nism (340/44/263) and
for thi
solution are si mila
so
solution
of
0.1
resolution
1.5-4.5
km.
January
29,
ution.
and data
of
The
by
P and
dep th
SH
of
in ferred
of the
minimum in
A secondary
just
the
beneath
seismic moment
to those
11
Moho.
(3.5
x 1024
of the
waveforms
with
th e
How ver, at a significan ce
the mantle solution
centroid depth
can
c a n be reiected.
The
therefore
the
in
lies
range
1982
about
3.23).
Th e fits
are quite good.
The January
valley
occurs at 6 km,
mecha
best-fitting
The
P
3.5 km, consistent wit h the epicentral pos ition
is
residual-versus -depth
dyn-cm)
with
both
characteri zed
The observed
bathymetry
The focal
is
mal faulting (346/40/2 80).
well
water depth
and
of this event (Figure 3.22)
mechanism
nearly
2.5
is
3.21).
Using
from
200
29,
km
1982 earthquake
north
of
the Kane
occurred
in
Fracture
the median
Zone
(Figure
a semi-automated centroid moment tensor inversion
the
GDSN,
Dziewonski
normal-faulting mechanism with
et
al.
[1983a]
a significant
obtained
strike-slip
a
component (353/61 /246).
waves,
is
Station
ra ther poor for this
not yet been dist ributed.
coverage,
event
From the
particularly
because many
body-waveform
for
records
P
have
inversion,
we
obtained a nearly pure normal-faulting mechanism (9/44/264),
with fault p1 anes
striking
subparallel
ridge axis ( Figur e 3.24).
Most
to
the
local
of the waveforms
trend of
the
are well-fit
by
the syntheti c sei smograms, especially the nodal character of the
SH wave at KBS.
The seismic moment
of
5.4
x
1024
dyn-cm
is
somewhat sma ller than the 8.1 x 1024 dyn-cm obtained by
Dziewonski et al.
[1983a].
the water depth 3 .8
The ran ge in
The
centroid
depth
is
2.4
km,
and
km.
centroid
and 6-7 km at a = 0.1.
depth
for
A secondary
this
earthquake
is
1-4
km
minimum in
residual-ver sus -d epth occurs just beneath the Moho at 6 km
The
depth.
(6.3
x 1024
focal
mechanism
(13/43/271)
and
seismic moment
dyn-c m) of the mantle solution are very similar to
those of the
best
solution.
However, the
fits
of most waveforms
are slightly worse than those of the best-fitting solution.
Using
a gradual
outside
the
the uncertainty
centroid depth
May 12,
Moho,
mantle
range.
solution deteriorates
lies
We therefore believe that the
is between 1 and 4 km.
1983
The earthquake of May 12,
1983 occurred in the median
valley about 250 km north of the 15020
(Figure 3.24).
by Dziewonski
motion.
and
'
A centroid moment-tensor
et al.
Frac ture Zone
sol ution
(354/50/259)
[1983b] shows predomina ntly normal-faulting
Station coverage for this event is good,
and the focal
mechanism
S hown
a I.
is
well-resolved.
in
Figure
3.25
One
of the
nodal
o f the
ridge axis.
s imilar to
[ 1983b].
the
is
4.1
There
very
planes
similar
is
The seismic
8.1 x
1024
The centroid
d epth is
Our inversion solution
to
that
to
subparal lel
moment
of
dyn-cm obtaine d by
depth
is
3.1
km,
Dziewonski
the
x
o f 7.1
(341/48/251)
trend
local
1024 dyn-cm
Dziewonski
and the
et
et
inferred
is
al.
water
km.
is
centroid
some trade-off between the
S ource mechanism
for
s econdary minimum
in
fixed depths
in
depth
and the
range 1-1 0 km.
the
A
residual-versus-dep th occurs
at
6
i mmediately below the Moho, but this sol ution can
be
rejected
a
We estimate that
= 0.1.
lies
in
th e range
km,
a
the cent roid depth of this earthquake
2-6 km at a s igni ficance level of 0.1.
DISCUSSION
The
16 earthquakes
occurred
two
along
Azores
spreading
region
segments
earthquakes
of
proper,
we will
study are
All
show
oceanic
14
consideration
different
plate
earthquakes,
of the
all
basins
lithosphere
discuss
Mid-Atlantic
remarkably
them
14 earthqu akes
probably
in
ridge-axis
their
nea rly horizontal
occurred
formed
in
for the
on
Since the Azores
by the
a mid-ocean
ear
ridge
thqua kes of th
source chara
with
and striking
to the local trend of the ridge axis .
all
Ridge.
rather than
normal-faulting mechanisms
plane dipping at about 450
Except
separately.
Ridge
similar
his chapter
in
boundarie s.
Mid-Atlantic
occurred beneath
rifting
The
four
under
at
1 east
cteristic
one nodal
approximately
The
and are approx imately
parallel
inferred T axes are
aligned with th e
spreading direction.
factor of 5,
The
from
3 to
seismic moments
15
x 1024
span
dyn-cm, and
functions are generally of similar durati on.
depths are all
very shallow,
between 1.2
and
a range
of only
a
the source time
The centroid
3.1 km beneath the
seafloor.
The water depth at the epicenter of each of these earthquakes is well -constrai ned by the predomi nant period of the
large water-column
water depth
reverberations in the P wavetrains.
from the P waveform is
estimated
3.2 with the maximum
uncertai nties in the epicentral
from
variable
sampling
contour maps
the two
depth
hundred
meters
from this
density
for all
eart
exclude the
near the
31,
earthquake)
inward-dipping
valley walls
with
reasonable
Mid-Atlantic
inner wal
that
t he agreement between
of the ea rt hquakes in this study
ls of the rift
We
valley (e.g.,
the May
they rep rese nt slip on major
contribute to the relief of the inne r
Luyendyk, 197 7].
however, that
Ridge earthquakes considered
occurred beneath the
We conclude
ity that some of these events
and
certainty,
accurac y,
floor of t he median valley.
or tha t
[Macdonald
om ship-track data of
hquakes in Tab le 3.2.
a 11
inner
faults
c oordinat es and in the depths
surprising ly g ood, to within a few
possibil
occurred
1971
a nd
is
comparison that
Given the
constructed
estimates
occurred beneath the
cannot
compared in Table
1ocal depth of the median valley as
indicated on availabl e bathymetri c maps.
obtained
The
none
It may be inferre d
of the
in this
chapter
rift mountains.
From the centroid
depth,
the source-time function,
and the
seismic moment,
dimensions,
we can
average
obtain
slip,
and
For si mplici ty,
earthquakes.
estimates of fault
average stress
drop for these
we assume that the zone of slip
ho rizont al length L and do wn-dip width w.
with
rectangular,
rough
estimate L fro m the length ts of th e source time function.
the basis
of t he lar ge wate r-column
We
On
reverber ations we assume
that rupture e xtende d to th e surface, and we estimate w from the
uni 1 at eral
centroid depth .
For
rupture veloci ty
Vr is take n to be equal
velocity,
or 3 .0
km/ s.
the centroid d epth;
about
450
fault
slip.
If
this
relatio n fo llows
width
th e faul t len gth,
The average sl ip
1966],
to 0.8 times the shea
W e estimate w from w = 2.8 h,
the fault
extent of faul ting
[Aki,
where the
where h
om a fault dip of
a cent roi d depth th at marks the average depth of
and
larger than
rupture, L = Vrts,
gi ven by this relat ion is
w
we set w = L so that th e vertical
i s no greater tha n the horizontal extent.
D can be estima ed from the relation Mo = VLwD
w here the rigidity u of the crust
4.0 x 1011 dyn /cm
2
from our assumed
crustal
is
st ruct
aken
to
re.
The
stress drop Aa can be estimated from the relat ion
Aa
=
cMo/(Lw 2 )
[Knopoff,
where c is
1958 ; Aki,
a qeometrical
qm#
const
from 30 to
100 km 2
(Table 3.5)
resolvable correlati on with moment
studied here.
The estimated
The a verage fault
stress
range shown by other
Anderson,
19751.
near unity
1966].
For the Mid-Atl antic Ridge earthquakes,
area Lw ranges
ant
drops
the estimated fault
and
shows no
for the limited
slip varies from
lie between 7 and
tectonic earthquakes
set
of events
10 to 80 cm.
70 bars, within the
[Kanamori
and
The two
earthquakes
in the Azores
larger centroid
depths
than
Using
source
structures
different
best-fitting
July
4,
occurred
centroid
1966,
in
a different
Mid-Atlantic
Ridge.
within these two
The
estimates
earthquakes
in
large, and the
small.
the Mid-Atlantic
depth
earthquake.
It
in
varies
from
environment
is
from
did
not
and
fautlt
region
average slip
and
Ridge earthquakes.
km to
those along
fault
area
(Table 3.5)
drop
the
6.8
these two
extend
stress
slightly
inversion,
3.7
possible that
width
the Azores
our
We believe that
earthquakes
of fault
region have
km for
earthquakes
the
rupture
to the seafloor.
for these two
are
the
probably too
are probably
too
Table 3.1
Epicentral Data for North Atlantic Ridge-Axis Earthquakes
Half
Date
Origin
Time
Latitude
oN
June 3, 1962
1502:25
22.51
-45.08
13.5
June 2, 1965
2340:23.1
15.93
-46.69
13.9
5.8
6.0
Nov. 16, 1965
1524:44.0
31.00
-41.53
12.5
5.9
6.3b
July 4, 1966
1215:27.4
37.51
-24.75
1.3
5.3
April 20, 1968
1018:00.0
38.30
-26.77
1.4
5.0
Sept. 20, 1969
0508:57.8
58.35
-32.08
May 31, 1971
0346:50.6
72.21
+1.09
April 3, 1972
2036:20
54.33
-35.20
June 6, 1972
0525:50.7
32.93
June 28, 1977
1538:37.8
June 28, 1977
Longitude spreading
oE
ratea, mm/yr
10.5
mb
5.7b
b
5.6
6.0
5.5
5.7
11.1
5.1
5.5
-39.79
12.2
5.3
5.7
22.64
-45.07
13.4
5.3
5.6
1918:36
22.68
-45.11
13.4
5.9
6.0
Jan. 28, 1979
1945:21
11.92
-43.70
16.2
5.7
5.6
April 22, 1979
0950:13
33.00
-39.72
12.2
5.6
5.8
June 28, 1979
0414:44.9
00.40
-24.96
18.2
5.5
5.2
Jan. 29, 1982
2232:06.1
25.52
-45.30
13.2
5.7
5.8
May 12, 1983
1051:50.0
17.61
-46.55
13.8
5.7
5.8
7.9
Epicentral and magnitude data are taken from the International
Seismological Summary for 1962 and the ISC for 1964-83.
a Half spreading rate calculated from model RM2 of Minster and Jordan
[1978]
b Magnitude M from Rothe [1969]
Table 3.2
Source Mechanisms Obtained from Body Waveform
Inversionsa
Dip
Slip
Centroid
Depthd
Water
Depthe
Depth of
Median Valleyf
42.8_+1.7
273.9±3.1
1.2±0.11
3.9
3.4-4.2
335.0± 1.9
37.5±0.8
243.5±2.1
3.0+0.06
3.4
3.5-4.0
3.0±1.0
24.6±2.5
45.9±1.0
265.3+2.8
2.1±0.14
3.2
3.0-3.5
4.0±0.3
5.0±1.0
135.4± 1.0
49.0+0.4
274.33±1.1
5.4+0.10
3.2
2.0-3.0
2.0+± 0.1
4.2+± 1.4
327.8± 2.0
34.0+0.9
292.6+2.2
4.6±0.14
2.4
2.0-2.5
15.2+1.0
4.0± 1.0
22.7± 2.4
41.8±1.1
260.5±2.7
1.5±0.08
2.2
1.5-2.0
May 31, 1971
6.8±0.5
5.0± 1.0
50.9±1.2
52.1_+0.5
283.8±1.2
2.3±0.06
3.1/2.8
2.5-3.0
April 3, 1972
4.4±0.8
3.2±0.8
8.2±3.0
48.3±1.2
253.6±3.4
1.9±0.09
2.8
2.0-2.5
June 6, 1972
4.33±0.4
4.0± 1.0
19.7±2.3
51.9± 1.2
252.5±2. 5
1.8+0.07
3.2
3.0-3.5
June 28, 1977
3.0+± 0.3
2.7+0.9
1.4_+2.0
46.1±0.8
253.6±2.3
2.5±0.09
4.0
3.4-4.2
June 28, 1977
10.7± 1.2
5.0± 1.0
0.5±1.6
44.3±0.5
254.7_+1.5
1.6+0.06
3.5
3.4-4.2
Jan. 28, 1979
6.3_±0.5
2.0_+1.0
19.6±2.1
45.9± 1.1
270.3+2.3
2.2±0.11
3.7
3.5-4.0
9.9±0.5
4.0_+1.0
16.6±1.3
51.6±0.6
261.7+1.5
1.8±+0.06
3.3
3.0-3.5
June 28, 1979
3.6±0.3
3.6± 0.9
345.5±1.5
40.0±0.9
279.7± 2.1
2.5+0.09
3.5
3.5-4.0
Jan. 29, 1982
5.4±0.8
3.0± 1.0
9.0± 3.3
43.88±0.6
264.1±3.3
2.4±0.08
3.8
3.5-4.0
May 12, 1983
7. 1±0.5
2.4± 1.2
341.0± 2.0
48.2+0.7
251.3±1.8
3.1±+0.08
4.1
4.0-4.5
Date
Mob
tsc
June 3, 1962
6.1±0.5
3.0+± 1.0
4.4+0.5
June 2, 1965
8.3+0.7
3.0± 1.0
Nov. 16, 1965
8.1+0.8
July 4, 1966
April
20, 1968
Sept. 20, 1969
April
22, 1979
Strike
Notes to Table 3.2
a) The range indicated for source parameters (Mo, strike, dip, slip,
and centroid depth) is one standard deviation (formal error).
b) x 1024 dyne-cm (1017 N-m)
c) Total duration of the source time function,
sec
d) Relative to the seafloor, km
e) Inferred from modeling the water-column reverberations in the later
portions of the P wavetrains.
f) Local depth of the median valley inner floor indicated by
bathymetric maps [Laughton and Monahan, 1978; Heezen and Tharp,
1978; Johnson et al., 1979; Searle et al., 1982; Toomey et al.,
1985].
4W
0
64
Table 3.3
4
Inversion results for the July 4, 1966
Azores spreading center earthquake
Source structure
Seismi c
moment,
1024
Double
couple
(strike/dip/slip)
Centroid
depth,
km
dyn-cm
Duration of
source time
function,
RMS
residual,
mm
sec
A.
Overall best solution
1.
Crustal halfspace
4.0
131/49/270
3.7
1.90
2.
Uniform crust over
mantle halfspace
4.0
135/49/274
5.4
1.89
3.
Two-layer crust over
mantle halfspace
4.7
133/50/271
6.3
1.80
4.
Mantle halfspace
4.8
130/50/267
6.8
2.05
B.
Secondary local minimum
1.
Crustal halfspace
2.6
150/48/286
7.9
2.27
2.
Uniform crust over
mantle halfspace
3.6
141/49/281
10.8
2.34
3.
Two-layer crust over
mantle halfspace
3.8
143/47/281
10.8
2.13
4.
Mantle halfspace
3.6
146/47/285
11.2
2.37
4
Table 3.4a
Date
Summary of Source Parameter Uncertainty Ranges for Atlantic Ocean Ridge-Axis Earthquakes
Range in centroid
given
depth, in kmin,
by t-test, a = 0.1
Secondary solution
with sharp lloho
Within
uncertainty
Depth,
range
km
for a = 0.1?
Secondary solution
with gradual Moho
Within
uncertainty
Depth,
range
for a = 0.1?
km
Adopted
uncertainty
ranges
Depth, km
a
June 3, 1962
1-2 and 6-7
6
yes
yes
1-2
0.1
June 2, 1965
1-5 and 6-8
6
yes
no
1-5
0.1
Nov. 16, 1965
1-3
6
no
m--
1-3
0.1
July 4, 1966
5-6
10.8
no
5-6
0.1
April 20, 1968
3-16
9
yes
3-6
0.15
Sept. 20, 1969
1-3 and 6-7
6
yes
6
no
1-3
0.1
May 31, 1971
2-3 and 6-9
8.2
yes
7.5
no
2-3
0.1
6
no
1.5-3
0.1
1-6
0.1
April 3, 1972
1.5-3
yes
6
no
June 6, 1972
1-7
June 28, 1977
2-9
5.5
yes
2-9
0.1
June 28, 1977
1-2
6
no
1-2
0.1
Jan. 28, 1979
1-9
6
yes
yes
2-3
0.15
6
yes
yes
1.5-2.5
0.15
1.5-5
0.1
1-4
0.1
2-6
0.1
April 22, 1979
1.5-2.5 and 6-7
June 28, 1979
1.5-5
6
no
Jan. 29, 1982
1-4 and 6-7
6
yes
May 12, 1983
2-6
6
no
no
Table 3.4b
Summary of Source Parameter Uncertainty Ranges for Atlantic Ocean Ridge-Axis Earthquakes
Date
Strike,
deg
Dip,
deg
Slip,
deg
Seismic Moment,
1024 dyn-cm
Source time function
duration, sec
June 3, 1962
0-8
38-45
265-285
5.6-7.2
3-4
June 2, 1965
332-338
33-46
235-255
7.0-9.2
3-4
Nov. 16, 1965
24-26
45-46
263-268
7.1-9.3
2-3
July 4, 1966
130-142
44-51
269-280
3.8-5.1
4-6
April 20, 1968
324-330
30-42
285-296
1.6-2.3
4-5
Sept. 20, 1969
22-300
39-47
256-277
12-16
3-4
May 31, 1971
45-520
50-55
275-284
5.6-8.0
4-5
April 3, 1972
6-15
42-50
250-258
3.2-5.0
3-4
June 6, 1972
5-307
46-53
240-265
3.5-5.1
3-5
June 28, 1977
-6-+4
44-51
234-260
2.8-4.1
1.5-3.5
June 28, 1977
-5-+10
44-47
247-256
9-11
4-5
Jan. 28, 1979
16-245
44-51
268-272
5.8-6.3
2
April 22, 1979
11-180
50-54
254-263
8.3-10.2
3-4
June 28, 1979
340-355
35-48
265-284
2.5-3.9
3-4
Jan. 29, 1982
6-17
42-49
260-280
5.0-6.2
3-4
May 12, 1983
335-346
48-51
248-260
6.0-7.9
2-3.5
Table 3.5.
Date
Estimated Source Dimensions and Stress Drop for North Atlantic Ridge Earthquakes
Fault length,a
km
Fault width,b
km
Fault area,
km2
June 3, 1962
9
3
30
50
60
June 2, 1965
9
8
70
30
10
Nov. 16, 1965
9
6
50
July 4, 1966
15
15
220
1
April 20, 1968
12
12
150
1
Sept. 20, 1969
12
4
May 31, 1971
15
6
9
5
12
5
20
10
June 28, 1977
8
7
10
8
June 28, 1977
15
4
40
40
Jan. 28, 1979
6
6
40
30
April 22, 1979
12
5
40
30
June 28, 1979
11
7
10
7
Jan. 29, 1982
9
7
20
10
May 12, 1983
7
7
40
20
April 3, 1972
June 6, 1972
a Estimated from duration of source time function.
shear wave velocity.
b Calculated from the assumptions:
Average
slip, cm
30
50
70
100
10
50
50
Stress drop,
bars
20
Rup ture velocity is assumed to be 0.8 times
w < x and w < 2.8h.
101
F IGURE
CAPTIONS
Figure
3.1.
e
Location of Mid-Atlantic Ridge and Azores
region
rthquakes stud ied in this chapter, along with the overall
d stribution of seismicity [Tarr, 1974].
Filled circles
s
ow earthquakes with mb > 4.5 during 1963-1972; open
c
rcles denote e vents
Figure
3.2.
in
this century with
Ms > 8.
Detailed bathymetry of the Mid-Atlantic Ridge
median valley near the epicenters of the earthquakes of
June 3, 1962,
and June 28,
m, are from Toomey
region
et al.
1977.
Bathymetric contours,
[1985].
The location of this
relative to neighboring elements
Ridge system is shown
solutions are equal
in Figure 3.23.
area
of microearthquakes
ocean-bottom seismometers
triangles,
Figure
3.3.
(dashed
Mechanisms
1982 are composite fault
by Toomey et al.
located at the positions
Fault plane
quadrants are shaded.
labeled A and B for February
clusters
of the Mid-Atlantic
projections of the lower focal
hemisphere; compressional
solutions determined
in
plane
[1985] for two
recorded by a network
of
and ocean-bottom hydrophones
shown
by numbered
circles and
respectively.
Comparison
line)
of observed (solid line)
long-period
1962, earthquake,
P and SH waves
with the
and synthetic
for the June 3,
focal mechanism
solution
obtained from the inversion plotted on the lower focal
hemisphere (equal-area projection).
normalized to an instrument
All
amplitudes are
magnification of 3000;
amplitude scales correspond to the waveforms
the
that would be
102
observed
on an
The two
vertical
series,
digitized
inversi on.
1 ines
deli mit
0.5
interval
circle, dilatation; sol id
emergent
the portion
for b th types
Symbol
arrival.
For
circle,
SH waves,
waveform
inversion.
the inversion is
Figure
3.4.
used i
the
waves
e open
and
1979;
Bathymetric
contours,
not
to
[1980].
used in the
obtained from
seismicity of the Mid-A tlantic Ridge
the earthquakes of June
and M ay
12,
i n km,
1983
epicenters are from the
are from Searle
study
are
hemisphere;
elt al.
[1982];
Earth quake
ISC for 1964-1979;
denote events with mb > 5.4.
2, 1965;
(Mercator p rojection).
regions shallower than 3 km are shaded.
lower focal
cross,
and Richards
shown for stations
epicenters of
of this
ime
compression ciorresponds
Th e source time function
January 28,
events
rument.
also s hown.
Bathymetry
near the
each
of
compression;
positive motion as defi ned by Aki
Some first motions are
an in
sei smogram from such
ori ginal
lar ger symbols
Fault plane solutions for
equal
area
projections
compressional
quadrants
of the
are
shaded.
Figure
3.5.
waves
Comparison
for the
June
of
the
observed
2, 1965,
and
event.
syntheti
c P and SH
3 .3 for
See Figure
further explanation.
Figure
3.6.
Ridge
16,
from
Bathymetry
near
1965;
and
seismicity
the epicenters
June
Searle
explanation.
et
6, 1972;
al.
of
and
[1982];
the
April
of
the
Mid-At 1 antic
earthquakes
22,
see Figure
1979.
3.4
for
of
November
Ba t hymetry
f
u rther
is
103
Figure 3.7.
Comparison of the observed and synthetic P and SH
waves for the November 16, 1965, event.
See Figure
3.3 for
furthe r explanation.
Figure 3.8.
Bathymetry and
near the Azores triple junction,
and mechanisms of the
April 20,
Searle
1968.
et al.
shaded.
Figure 3.9.
includ ing
earthquakes
Ju ly
of
Bathy metric contours,
[1982];
regions
centers
of th e spreading
seismicity
the epicenters
4,
1966,
in km,
shallower
and
are from
than 2 km are
See Figure 3.4 for further explanat ion
Comparison of the observed and synth et i
waves for the July 4, 1966, event.
and SH
See Figu re
for
further explanation.
Figure 3.10.
Comparison of the observed and synthetic P and SH
waves for the April
20
1968,
See Figure
event
3.3 for
further explanation.
Figure 3.11.
Bathymetry and sei smicity of the Reykjanes and
Mid-Atlantic Ridges near th e epicenters of the earthquakes
of September
contours,
20,
in km,
1969,
and
April
3,
1972.
are from Laughton and Monahan [ 1978];
regions
shallower than 2 km are shaded.
further
explanation.
Figure
3.12.
waves
Bathymetric
Comparison
of the observed
for the September 20,
1969,
See Fig ure
and syntheti
event.
See Fi
3.4 for
P and SH
re
3.3
for further explanation.
Figure
3.13.
Bathymetry
and seismicity
of the Mohns
the epicenter of the earthquake of May
stereographic projection).
Bathymetric
31,
1971
contours
dge near
olar
in km,
104
are
,,,,,
om Johnson
rr
are sh aded.
Figure
3.14 *
et al.
[1979];
regions
2 km
shallower than
See Figure 3.4 for further explanation.
Comparison of the obse rved and synthetic P and
waves for the May 31,
1971,
eve nt.
See
SH
Figure 3.3 for
furthe r explanation.
Figure
3.15 .
waves
Comparison of the observed and synthet ic
for the April
See Figure
3, 1972, event.
P and
SH
3.3 for
furthe r explanation.
Figure
3.16 .
Comparison of the observed and synthet ic P and
waves for the June 6, 1972,
event.
SH
See Figure 3.3 for
further explanation.
Figure 3.17 .
Compari son of the observed and synthetic P and
waves for the mb
5.9 event of Jun e 28 , 1977.
See
Figu
3.3 fo r further exp lanation.
Figure 3.18 .
Compari son of the observed and
waves for the mb
syntheti c P a nd SH
5.3 event of June 28 , 1977.
This
earthq uake was the first in a series of 5 events with
hours , including the earthquake of Figu re 3.17.
See
4
gure
3.3 fo r further exp lanation.
Figure 3.19 .
Compari son of the observed and
waves for the Janua ry 28,
1979,
event.
synthetic
See
Figure
P and
3.3
SH
for
furthe r explanat ion
Figure 3.20 .
Compari son of the observed and
waves for the Ap ril1 22,
1979,
event.
synthetic P and
See Figure 3.3 for
furthe r explanat ion a
Figure 3.21.
Bathymetry and seismicity of the Mid-Atlantic
Ridge near the epicenter of the earthquake of June 28,
SH
105
1979.
Bathymetric contours,
Tharp [1978];
Figure
Figure
3.22.
in km,
regions shallower than
are from Heezen
and
3 km are shaded.
See
3.4 for further explanation.
Comparison of the observed and synthetic P and SH
waves for the June 28,
1979,
event.
See Figure 3.3 for
further explanation.
Figure
3.23.
Bathymetry and seismicity of the Mid-Atlantic
Ridge near the epicenter of the earthquake
1982.
[1982];
Bathymetric contour
regions shallower
box centered on the median
area shown in
in km,
3,
of January 29,
are from Searle et al.
tha n 3 km are shaded.
va
The small
iley near 230 N outlines the
greater deta ii in Figure 3.2.
See Figure 6
for further explanation.
Figure 3.24.
Comparison of the
waves for the January 29,
observed and synthetic
1982, event.
and SH
See Figure 3.3
for
further explanation.
Figure
3.25.
Comparison of the observed and synthetic
waves for the May 12, 1983 , event.
further explanation.
See Figure 3.3
and SH
for
106
0
75
0
60
40
20
0
0
0
60
0
W
40
0
W
Figure 3.1
20
0
W
0
0
N
N
N
N
107
23°00'
220 30' l/.
44050'
45010'W
Figure 3.2
Figure 3.3
109
20N
150N
50W
45"W
Figure
3.4
J 10N
40W
JUNE 2,
1965
P WAVES
SH WAVES
4
GD H
K G
Source Time Function
WES
GDAH
(sec)
i
mm.
60
sec
sec
Figure 3.5
111
35°N
J 29°N
37"W
0
44 W
Figure 3.6
NOVEMBER
P WAVES
16,
1965
SH WAVES
KON
Source Time Function
'
NOR
(sec)
GDH
KON
MAL
C
CAR
s
P
LPA
LPB
WIN
CAR
LPB
mmLPB
mm
sec
sec
Figure 3.7
4j
S
Q
40'N
35'N
25W
30'W
Figure 3.8
Figure 3.9
APRIL 20, 1968
SH WAVES
P W,
6
GU11
Source
(Fvc)
NOl
"/~iC~
ifit
T ime FiiO(, lion
KEV
KTG
iv~SCl
UME
A
-
-op
I IIN
i--
SJG
3
TRN
mm
2
mm
50
50
Sec
sec
A
Figure
3.10
116
60*N
55 0 N
J 50°N
40'W
34'W
Figure 3.11
28W
SEPTEMBER
20, 1969
S
SH WAVES
SHSource
unclion
Time
COL
',
(sec)
S
NIL --
NIL
L
TAB
TAB
I'
--
-OGD
BEG
NAI
SJG
10
NAT
S
20
lmm
mm
50
5O
sec
sec
Figure
3. 12
Figure
3.13
Figure
3.14
Figure 3.14
APRIL 3, 1972
SH WAVES
P WAVES
4
Source
Time
KEV
riinction
(sec)
C
-
TU
AAE
BEC
7
r
NAT
.I" lf
mm
61
mmS"C
sec
Figure 3.15
i
il
i
6
i
6
4
4uDi
4
Figure
3.16
~a.
JUNE
28.
1977
P WAVES
SH WAVES
6
, e
CIC
L
I
KC
I
Sorce Time FuInc tion
SCP
pO
(9Pc)
ESK
LPS--
IU
mm
mm
61s
61
Sec
Sec
Figure
3.17
JUNE 28,
1977
SH WAVES
P WAVES
4
urce Time
AAM
Function
KEV
(s e c)
S
BL
,
s H
3
2
mm
mm
71
sec
sec
i
Figure 3.18
Figure
3.19
Figure
3.20
3" N
°
0
J 3S
30°W
26 0
W
Figure 3.21
22 0 W
Figure 3.22
27oN
24°N
48 W
45
w
210N
0
42 W
JANUARY 29, 1982
SH WAVES
P WAVES
4
KBS
Source Time Function
KEV
(sec)
PTO
BOG
LPB
10
mm
51
5(
sec
sec
Figure 3.24
Figure 3.25
131
CHAPTER
4.
RIDGE-AXIS EARTHQUAKES
NORTHWEST AND CENTRAL
IN THE
INDIAN OCEAN
INTRODUCTION
we determine the focal depths
In this chapter,
mechanisms
and
rift earthqu akes in the north ern Red Sea,
Sou rce mechanism
earthquake.
Indian Ocean,
in the
of 11 ridge-axis earthquakes
studies of
19841
ridge
mid-ocea n
Oceans
have
1981; Chapter
s hown
that
they
ridge-axis earthq uakes have s everal consistent fe atures:
are extremely
sha llow, locate d within the median valley,
characterized by nearly pure normal
dipping at about
450
of the ridge axis .
and
stri king
faulting,
parallel
wit
h nodal
and
planes
overall trend
to the
However, both the Mid-Arctic
and northern
simple tecton ic
histories.
Mid-Atlantic
Ridg es have rela tively
In contrast,
whil e mid-ocean ridges in the Indian Ocean have
generally compara b e spreadin g rates
2
near-ridge
one
earthquakes in the North Atla ntic [Trehu et al.,
31 and Arctic [Je msek et al.,
and source
(half-spread
ing rates
between 6-22 mm/y r ,they are the products of a mo re complex
tectonic evolutio n
Thus,
a synthesis
of the
of ridge events in the Indian Ocean will
sou
provide
rce mechanisms
)
us a chance to
compare the focal depths and source mechanisms of earthquakes
over a range of both spreading
settings.
rates
and different
tectonic
As in the previous chapter, we pay speci al attention
to focal depth, aIs it
can
be
an
extremely
important
quantity for
unders tanding the depth extent of brittle behavior beneath the
ridge axis.
132
SPREADING CENTERS OF
The
the
includes
Carlsberg
Indian
To the west,
Ridge.
Red
end of the
of
account
The Red
km
1800
Sea
Red
the
from
Peninsula,
the basement
that
to
the
[Tramontini
inner half
suggesting
There
the crust
and
that
are
several
trough
schools
the Red
studies,
and
trough
is
that
Sea.
that
most
is
of
of
between
between
oceanic
the main
shown
are
similar
higher velocity for
220
underlain
thought
with
refraction
later seismic
found
150
1964].
have
trough
and
by
23 0 N,
oceanic
about the
One hypothesis,
the
Sinai
Girdler,
1964]
the main
A
1969]
main trough
the main
beneath
geologic
the axial
of the
and
of
of
associated
is
Girdler,
rocks.
Davies,
and
and
beneath
velocities
continental
those of the
study
[Drake
data
[Drake
anomalies
Straits
the
trough
trough
some
extends
15 0 N to
axial
axial
The
1983].
large-amplitude magnetic
refraction
from
separating the
consists
Sea
Red
and narrower
a deeper
24 0 N [Cochran,
Seismic
[1982].
to the
Peninsula
trough
a main
It
Shield.
The
1982].
[Shazly,
and
African
the Sinai
of
tip
shelves,
continental
and
Shazly
primarily
drawn
a NNW-SSE-trending depression
is
Peninsula and the
Mandab
El
The
Sea
The
Arabian
and
[1982]
is
below
presented
Schlich
of
Central
system.
Rift
the
the southern
joins
Ridge
Sheba
East African
the
tectonics
regional
the syntheses
from
Bab
and
Sea
the
Ocean
of Aden,
the Gulf
in
Ridge
and the N-S-trending
Ridge,
NW-SE-trending
Indian
northwestern
the
Sheba
trending
E-W
in
system
ridge
active
REGION
INDIAN OCEAN
NORTHWEST
THE
crust
nature
based
is
trough
crust.
on
seismic
limited
is
of
to
underlain
133
by cont inental
Genik,
1972;
crust
Ross
and
Schlee,
competi ng hypothesis,
anomaly
is
studies,
underla in
by
and
[Huchinson
oceanic
1973;
based on
that
the
crust
Engel,
Owell
1972;
et
al.,
plate kinematic
Red
Sea
[McKenzie
Lowell
is
almost
et
al.,
and
19751.
and
A
magnetic
entirely
1970;
Styl es , 1974, 1976; Roeser, 1975; Styles and Hall,
Girdler a nd
1980].
Cochran [1983] has made a convincing case for an alternative
model
i n which (1) sea floor spreading is
souther n Red Sea,
(2)
mid-oce an ridge is
continental
present
in
occurring
extension
the
northern
is
Red
in
the
ongoing
Sea,
but
and
n
(3)
transit ion area separates the two tectonic regimes.
The She ba Ridge
Ridge
The Sheba
Carlsberg
Ridge
the west,
the
ridge jo
western
portion
valley,
while
the
east.
west
it
from
Zone.
The Sheba
E-W
zones.
accretionary
[Sykes
et
al.,
and
at
1985]
1982].
abo ut
is
plates.
The
560E
to
Fractu re
the
Zone.
African Rift
To
to the
center to the north.
The
ridges to
increase s progressively from
trend of the ridge axis
SE -NIW
offs et by
bou ndary
Landisman,
from
system.
Ridge consists of a single deep
the va lleys
Earth quake
plate
st
km
into a succ ession of parallel
of
Ridge
Owen
sp readi
Sheba
break s
[Schlich,
varies
ins with the
the
The depth
to east
fracture
of
300
by about
-SW-trending
axial
Sea
of the Mid- Indian Ridge
offset
by the
and the Red
south
part
ridge
the
east,
To the
i
at th e Owen Fracture
several NE-SW-trending
ep icent ers
clear ly delineate the
be tween the Afri can and Arabian
1964; Sykes,
1970b] or Indo-Arabian [Wiens
Seismic refraction da ta [Laughton and
134
Tramontini,
1969] and
and
Girdler,
the
Gulf of Aden
the
half-spreading
greater
The
1964;
to the
is
Owen
Somali
al.,
1965;
and
is
Ridge
Zone
1968;
1969] and
rate,
12
[Le
The
13 mm/yr
The
at
Indian
the
the
al.
of
Pichon
(9
slightly
mm/yr).
and
and
Sea
Sclater,
axial
1971].
epicenters
magnetic
Heirtzler,
and
the
[Matthews et
Heirtzler,
magnetic anomalies.
on
equator and
The
ridge
[Barazangi
The present
anomalies,
1968;
1968;
McKenzie
is
about
and
Ridge
Central
southern end
east
by
[19711
a series
are the
Celeste
west
between the
Le Pichon
Indian
of the
triple
the
of
Vema
shown
short
Fracture
Trench,
Zone.
extends
junction
the
25 0 S.
Ridge
Langseth
and
that
Central
the
ridge segments
zones.
Argo Fracture
Earthquake
from the
Ridge to
near
Chagos-Laccadive
Plateau.
have
Ridge
Carlsberg
northeast-southwest-trending fracture
zones
calculated
perhaps
been well-studied
earthquake
based
the Mascarene
Fisher et
consists
1966;
axial
N-S-trending
bounded on
by
[1970]
separates the Arabian
by
African-Indian-Antarctic
west
al.
in
1971].
Central
equator
Drake
that the crust
10 mm/yr,
NW-SE
McKenzie and
clearly defined
Sclater,
et
than to the
ridge has
Vine,
half-spreading
to
about
trends
and
The
Cann and
Dorman,
Laughton
(11 mm/yr)
Basin.
Fisher et al.,
axis
have shown
1963;
Ridge
Fracture
East
oceanic.
rate to be
The Carlsberg
the
lineations [Girdler,
1966]
Laughton,
east
Carlsberg
magnetic
epicenters
cut
It
and
Taylor
the
on
is
the
[1967]
Indian
and
Ridge
by many
The major fracture
Zone,
and the Marie
and axial
magnetic
135
anomalies clearly delineate the plate boundary.
rate varies between
half-spreading
south [Fisher et al.,
N40
0
E.
Fisher et al.
18 and 24 mm/yr from north to
1971]; the spreading direction is about
[1971]
suggest that the bathymetric
complexity was caused by a change in
N-S
in
the
The present
spreading direction
late Cretaceous and early Tertiary
to NWI-SE
from
in
The Chagos Fracture Zone was then broken up into
Miocene time.
a series of ridge segments joined by transform faults.
EARTHQUAKE DATA SET
assembled a list
We first
of mid-ocean
ridge earthquakes
that occurred on the ridge systems of the Indian Ocean between
latitudes 250 S and 30 °N and over the
search was based primarily
Seismological
Center (ISC)
listings of the National
interval
1964-1984.
The
on the Bulletins of the International
1964-1982, the monthly
for the years
Earthquake
Information Service (NEIS)
for the years 1982-1984, and published studies of fault plane
solutions for the study area
et al.,
1969;
Ben-Menahem
Dziewonski
1983;
and Aboodi,
and Woodhouse,
We chose only the earthquakes
long-period
known
1983a;
Dziewonski
Dziewonski
normal-faulting earthquake
data
(October 18,
for the
et al
et al.,
1984].
We excluded
[Bergman and Solomon,
on transform faults,
unable to obtain a satisfactory
Epicentral
McKenzie
large enough to yield clear
intraplate earthquakes
strike-slip events
obtained
1971;
1969;
P and S waves at teleseismic distances.
near-ridge
1984],
[Banghar and Sykes,
and a complex
1964) for which we were
solution.
14 earthquakes
reliable solutions are listed in
for which we
Table 4.1;
their
have
136
epicenters
are shown in
the table,
we have used P and SH waves
at
instruments
(General
4.1.
For all
and 800.
All
of the events
at epicentral
these events have epicenters
valley segments on bathymetric maps of the GEBCO
Bathymetric Chart of the Ocean)
series [Laughton,
Fisher et al.,
19821.
6 to 22 mm/yr,
according to plate kinematic model
and Jordan
between
normal
in
recorded on long-period
stations of the WWSSN or GDSN
distances between 300
within median
Figure
[1978].
The local
half-spreading rates vary from
The earthquakes
5.4 and 6.1,
and
all
1975;
RM2 of Minster
have body wave magnitudes
but the one
near-ridge event
have
faulting mechanisms.
WAVEFORM INVERSION RESULTS
We present here the details of the source mechanisms
obtained
from the inversion of P and SH waveform s
Indian Ocean ridge-axis,
earthquakes.
2 northern Red Sea,
and
for the
1 near-ridge
The best-fitting double-couple ori entations,
seismic moments,
and centroid depths obtained fr om the
inversions are given in Table 4.2.
Richards [1980].
Centroid depths are
The average water depth
above each
is
that o f Aki
hypocenter is inferred by
A discussion of the
is given for each earthquake.
ranges are calculated at a significance level
cases,
we also give P and B values
discussion of these quantities).
and
relative t o the seafloor.
matching the synthetic and observed seismograms in
portion of the P wavetrain.
formal
The conventi on for
describing a double-couple orientation
centroid depth
11
the
ringing
uncertainty
in
Th e uncertainty
of 0.1.
(see Appendix
A for
In
a
some
137
Except
for two
inversions
waveform
a water
structure:
and
Vs
km/s
and
Vs =
probably
et al.,
=
1970;
6.0
km/s
earthquake,
on
models
layer
we
4.6
For
in
occurred
Vs
have
inversion
=
thinned
3.46
also
19,
This
ridge
the
the trend
earthquake
axis
chang
of the
Station coverage
SH waves.
is
The
well
rid
is
for
by
crust [McKenzie
March
1969
31,
influence of
of the
with
halfspace
structure
study are
given
di
the East Sheba Ridge where
es direction (Figure 4.2).
ge axis is
SE-NW.
To the east,
To the west,
the trend
ESE-WNW and the ridge becomes shallower.
t his event is
poor,
particularly for the
solutio n for this event
normal
(282/38/24 1) .
and the
fit,
earthquakes,
earthquake.
occurred on
inver si on
characterized
component
Sea
= 8.1
1964
ridge axis
of the
the
Details
later in the discussion of this
March
with Vp
a crustal
For
examined the
solutions.
Red
continental
we used
km/s.
Vp = 6.4
seismic velocities
the northern
1983],
Cochran,
a single
variable thickness,
of
all
Sea,
source
the same
and a mantle half-space
km/s.
and
Red
northern
performed with
were
= 3.7 km/s,
km/s
Vp
the
layer of thickness 6 km and
crustal
which
in
earthquakes
fault ing with a small
(Figure 4.3)
strike-slip
The obs erved P and SH waveforms
p of one nod al
plane is
are
constrained by the
P wave at
BUL which
shows a clea r upward first motion.
strike
one
nodal
plane agrees well with the trend of the
the we
st ; the stri ke of the second
ridge
of
axis
to
(3170)
paral l el to th e trend of the rid ge axis to the east.
The
is
The
seismi c moment of this earthquak e is 5.0 x 1024 dyn-cm, the
138
centroid
depth
epicenter is
is
1.8 kmin,
and the average water depth
about 3.3 km,
at the
which is consistent with the
bathymetry.
Because of the poor station coverage, there is a
sign fi cant
trad e-of
mech nism.
For
trike
the
2300
the
the d
3400
a fi
d ce ntroid depth varying from 0 to 15 km,
st-fi t ting mechanism varies from 2800
var
to 315
betw een the centroid depth and the focal
from
Us
250
to
and the
410,
slip
t -test we estimate the
the
to
varies
range
from
of
ccept able depths to be from 0.5 to 13 km at an 0.1 significance
evel.
However, w e
think that the depth uncertainty is smaller
han t his range wo ul d indi cate because the P waveform at ATU is
ot very well-mode led when the centroid depth is greater than
km .
Also,
at
de
p th s bet ween 4 and 6 km and greater than 8 km,
he sy nthetic SH w a ve s at SHL, CHG, BAG, and NHA show a small
pward motion whic h i s not
he P value is 0.1 5
out
of
That
that
is,
we
would
From
range
in
December
This
Argo
in
and
the
the
centroid
3,
a type
B = 0.6
depth,
I
error
for
the
only
4
km
The use of a = 0.15
our cha nce of detecting the difference
improves
slightly.
make
centroid
we hav e 3 chances in 5 of being unable to
is
etect the differe n ce in s olutions.
B = 0 .5)
4 km
At an a level of 0.1,
20 times.
oluti on.
At
observed.
above
depth
analysis,
is
1 to 4
our
best
estimate
of the
km.
1964
event
occurred
on
the Central
Marie Celeste Fracture
epicentral
region
is
Zones
marked
by
Indian
Ridge
(Figure 4.4).
two
parallel
between
The
the
ridge
valleys,
and
139
the earthq uake epicenters are scatt ered over both valleys and
the
dividi ng
ridge.
east ern va lley.
well1 as
earthquak e occurred near a bend in the
This
Using
motion s of the phases
first
th e polarization
of S waves ,
mechanis m with
a normal-faulting
prop osed
Banghar and
strike-slip motion for this
event
estimated t he seismic moment
to be
3.4 x 1025
of
component
Tsai
(314/60/216).
[1969]
Sykes
a large
as
PKP,
P and
[1969]
using the
dyn-cm,
vertical Rayleigh -wave ampli tude spectra at two stations, but
this estimate was based on an assumed focal depth of 65 km.
inversion solutio n (Figure
Banghar and Sykes [1969],
(302/46 /249).
component
4 .5) is similar to the solution of
bu t with a smaller strike-slip
The
centroid
seismic moment is 6.8 x 1024 dyn-cm,
is
depth
and
synthetics.
The
the trend of the
nodal
plane striking
ridge axis,
is
at
probably the
by
the
subparallel
fault
and SH waves have almost thr ee-quadrant coverage.
The
mechanism for thi s earthquak e is
3
less than
10 km.
100 for
stable:
all
both
P
focal
angles
vary
centroid depths held fixed between 1 and
Using the t-test,
acceptable centro id depths
we estimate that the range in
is 1 to 6 km at a = 0.1.
that this estimate is accurate because the
failing to detect the difference (8)
We
believe
probability
is smaller than 0.4 for
depths outside this range.
August 15,
to
plane.
The station coverage for this event is quite good:
very
is
depth
the water
3300,
The
km.
1.6
All obse rved wavefo rms are very well fit
3.4 km.
The
1966
This earthquake occurred in the Carlsberg Ridge median
140
valley
(Figure
4.6).
Using P-wave first motions,
Banghar and
Sykes [1969 ] found a normal-faulting mechani sm (272/28/270)
this event.
One of
well-determ ined.
the nodal
Tsai
planes
[1969]
in
1.26 x 1025 dyn-cm for this event,
r solution is not
thei
a se ismic
estimated
using
for
moment of
vertical
the
Rayleigh-
wave amplit ude spectrum at one station and a n assumed focal
depth of 45 km.
The
best
inversion
solution
similar to that of Banghar and Sykes [1969]
(Figure 4.7) is
but has a small
strike-slip component (294/51/260), a centro id depth of 3.4 km,
and a seismic
moment
of 4.4
x
1024
The emergent
dyn-cm.
character of the SH wave at WIN is revealed by the clear arrival
of SV energ y on the radial component.
waveforms are very well-fit
by
our
However, us ing a higher value of t*
better over all
fit.
The
inferred
The s hapes
mechanism.
inversion
for
of P and Si
S wa ves
water dept h
would
of
cons iste nt with the ISC epicenter and the lo cal
a
produce
3.8 km is
bathymetry.
The
maxi mum wat er depth in the median valley is bet ween 4.0 and
4.5 km.
Th is
earthquake may
inne r walls of the rift
The
and
is
occurred
n ear
P waves.
The focal
quite stable for cent roid depths
10 km.
lie between
one of the the
mountains.
st ation coverage for this event is
particularly for the
event
have
fairly
mechanism for this
held fixed between
Using the t-test,
we estimate the
2 and 6 km at the
0.1
1
centroid depth t
significance level.
estimate is probably accurate; 8 < 0.3 for depths
range.
good,
This
outside this
0
141
March 31,
1969
This earthquake occurred at the mouth of the Gulf of Suez
at the northern end of the Red Sea (Figure 4.8
al.
[19691
used P wave fir
..
McKenzie et
st motion polarities to infer a pure
with t he T axis roughly
normal-faulting mechanism (310/34/270),
Ben Menaham and Aboodi
perpendicular to the ridge axis.
[1971]
found that the su rface-wav e radiation pattern fits
a normal-
faulting solution with sma 11 strike-slip compo nent
(327/44/310).
They estimated the centroi d depth to be about 15±+5 km and the
seismic moment to be 1.6 x 1026 dyn-cm.
The best-fit ting doub le-couple mechanism
indicated by body waveform inversion (Figure
that of McKenzie
et
al.
[1 969].
shallower than th at of Ben
The centroid
Menaham and Aboodi
seismic moment (1.1 x 1026 dyn-cm)
their value.
is
somewhat
(294/37/271)
4 .9)
is
close to
depth of 6.2 km is
[19711, and the
smaller than
The source structure used for th is inversion is a
half space with ve 1ocities appropriate to conti nental crust
(Vp = 6.0 km/sec, Vs =
3.4
6 km/sec).
The station coverage for both P and SH waves
and the observed waveforms are all
strike is
control led mostl y by the SH waves.
motions of the P waves at
of the nodal plan es.
small,
very well
f it.
is excellent,
The fault
Sma 1 downward
SDB and BUL constrai n t e dip of one
Beca use
the water wave for
his event is
we could not estimate the water depth usin g the P
wavetrain.
The water depth
used in the inversion
(0.2
km)
estimated from the bathymetric map.
There is some trade-off between the centroid
depth
and
the
142
strike
and
(which varies
slip
in the
(2600-2800)
range
centroid
1-20
depth
is
between
are
km.
2850
quite
to
3100),
stable
for
Using the t-test,
between
5.5 and
8 km
but
the dip
(350-400)
centroid depths
we
estimate
at a = 0.1.
fixed
that
For
the
this
earthquake, the waveform fits appear to be good between centroid
depths of 5 and 12 km, with the source time function va rying from
5 to 10 sec
in du ration.
Since the velocity structure in the ep icentral
area is not
well-known, we have repeated the waveform i nversion for this
2 other velocity structures.
event with
Fi rst,
we used our
standard oceanic structure and obtained a c entroid depth of 12 km
and a more complex time function.
Second, we used a model
sisting of a 15-km-thick continental crust
(Vp = 6.0 km/s,
3.46 km/s ) overlying a mantle halfspace (Vp = 8.1 km/s,
and we obtained a centroid depth of 9.2 km.
4.6 km/s),
conVs
Vs =
A weak
motivation for such a velocity structure is given by the proposal
of McKenzie
et al.
[19701
that the crust in the Gulf
of Suez
might have been thinned to half its origina 1 thickness.
4.3 summari zes
Table
the inversion results of this earthquake with
three different velocity struc tures ; see also Appendi x C.
we are
not certain
cannot
rule out
April
22,
This
about the v elocit y structure in th is
a centroid dep th as great as
Si nce
area,
we
9 or 12 km.
1969
earthquake occurred near the intersection between the
southern end of the East Sheba
(Figure 4.2).
A very good fit
waveforms from this earthquake
Ridge and the Owen Fracture Zone
to the observed P and SH
(Figure
4.10)
is
obtained with a
143
faulti ng mechani sm with some strike-slip
normal
stat ion distribution
(350/63/305).
The
event is
but the
poor,
The
an excellent cons traint for one of
is 5.1
km
Although the st ation coverage for
solutions in which we fixed the focal
other source
parameter s,
and strike,
this mechanism
accurate.
quite
dip,
For
depth and inverted for
the seismic moment
than
varied by less
and sl ip varied by less than 100.
well -const rained, the focal
is
this
motions at
hanism is very stable.
is poor, the source meec
P waves
20%,
for
seismi c moment is 2.9 x 1024 dyn-cm, and
The
the centroid depth
P waves
c le ar upward first
stations SHL and CHG p rovide
planes.
of
P waveforms clearly require some
strike-slip component.
the nodal
component
depth
Since
is probably
We determi n ed that the centroid depth of this
earthquake is in the
range 3 -8 km at a significance level
of
A secondary mini mum in the residual-versus-depth curve
0.1.
occurs
at 7 km,
1 km
below t he Moho.
The focal
mechanism
(344/63/302) and seismic mom ent (2.5 x 1024 dyn-cm) for the
mantle soluti on are very sim ilar
to those of the best solution.
The t-statist ic calculated f or this depth
smaller than the threshhold ta.
This
is only
slightly
solution becomes degraded
and lies outs ide the uncerta inty range, however, when a gradual
Moho structur •e i s used for t he inversion.
analysis,
we bel ieve the cen troid depth is
Based on the above
between
3 and 6 km.
The wate r depth determi ned from the period of the
water-column
reverberations
in the P wavetrains
is about 4.9
The water depth
in
the median valley indicated on the
bathymetric map
is
between
4.5 and 5 .0
km.
Unlike pure
km.
144
normal-faulting
for
this
earthquake are
the P waves,
component.
fault
probably
If
plane,
We
axis.
ridge-axis
think
this
south
lies
an
December
14,
Owen
1969
3 parallel
the
patt
valleys
the
off-ridge
Carlsberg
cycle
of
the
ri dge
probably occurred in a
region where the crust and
in the normal
ridge environment.
next
the
about
300 km
(Figure
e rn suggests that the easternmost of
The
South
bends
located
Carlsberg Ridge
ISC
epicenter for
valle Y to the west,
event.
Ridge
Zone on
is
t he vicinity of the epicenter probably
in
ridge axis.
within
first
ees with the trend of
earthquake
Fracture
The seismicity
marks
the
1969
of the
4.2).
the
14,
with
of water waves
plane striking at 293
tran sition
colder th an
amplitudes
significant strike-sl
nodal
earth quake
fault
upper mantle are
The
to its
the
the
compared
then the stri ke agr
ridge-transform
December
small
due
we take
events,
and
of the
the
this
and thus
epicentral
median valley
earthquake
it
region,
shifts
is
probably
the
to
the
southwest.
The focal
characterized
component.
mechanism
317/ 63/ 61,
by thrust-fatulti ng with
mornent
The seismic
centroid
depth
3.0
km.
The station
all
observed
is
by
P axis is
is
Figure
4.11)
is
some strike-slip
10.3 x 1024 dyn-cm,
the
km,
and
the inferred water depth is
cover
age
for this event is very g ood and
3.1
waveforms
well-constrained
inferred
(
are
wel 1 f it.
P wavei s
One nodal
plane
at NI L, SHL, CHG, and SNG.
nearly horizontal
aligned with the spreading direction.
and is
is very
The
approximately
One nodal plane is
145
subparallel
to
is
N-S
striking
trend
ridge,
of the
parallel
to
other fault
north
a linear depression
presumed median valley
of the
and the
of the
within the African
and thus
We e stimate that the age of the lit hosphere in the
plate.
epicentral
region
half-spreading rate
about 3 M.y. from the
is
predicted by model
1978].
RM2 [Minster and Jordan,
This event
of the closest known thrust-fau Iting earthquake to
is thus one
ridge.
an actively spreading
There i s very little
trade-off
focal mechanism and
between
Using the t -test,
we estimat e
centroid dep th for this
event.
that the cen troid depth
is between 1 and 7 km at a = 0.1.
are quite good fo r
this earthqu ake, the fits
between
2 an d 6 km.
At depths
P waveforms deteriorates.
greater than
We believe that
of 1-7 km fo r the depth uncertainty range
at
10.3 80 N and 56.830E)
occurred
Zone
and the
long-period
June
28,
The
S waveforms
earthquake
km northwest
earthquake
is
P and
Ridge,
but we
the fit of the
the est imate
i s proba bly
accura te.
1969,
mb
of
were
the Owe n Fracture
unable to match
the
in detail.
1972
This
15
Carlsberg
7 km,
yo ung lit hosphere of
in
the African plate near the intersection
F or
fixed depths
(March 29,
Another thrust-faulting earthquake
5.6,
plane
The epicenter of this event is about 30 km
r egion.
epicentral
west
the
station
of the
(Figure
characterized
occurred
by
in
the northern
epicenter of
4.8).
March 31,
The inversion
nearly
distribution
the
pure
for
normal
Indian
Red
solution
faulting
this
event
Sea
about
1969
(Figure 4.12)
(288/40/260).
is
fairly
poor,
146
however,
6.1
especially
km,
and the
structure used
earthquake
event.
are
the
very
between
from
varies
the slip
between
shows
for the
2600
varies
1 and
is
5-13
that
at
half
cycles
wide
and too
3000,
from
2300
km.
The
is
NDI,
This
Ridge
The
km
PTO,
11
the strike
500,
depths
of
the
B
that
by
the
seismograms
and
=
and
fixed
range estimated
VAL,
km,
1969
a significant
to
the first
We think
at
dilatational
COP are
0.6.
The
the true
too
P value
centroid
11 km.
the southern
(Figure
mechanism
subparalle 1 to
is
300
for this
end of
the
Carlsberg
north of the Mabahis s Fractu re Zone (F igure 4. 13).
normal -fau Iting
depth
0.13.
occurred
invers ion solu tion
The
depth
of
31,
there is
for centroid
source
crustal
March
from
is
The
centroid depth;
varies
MAL,
a depth
dyn-cm.
centroid depth
of the
11
depth
1972
event
just
fit.
than
betwee n 5 and
probably
and
examination
SNG,
At
large.
dip
centroid
greater
CHG,
102'
coverage,
3000
to
A closer
depths
September 7,
the
centroid
a continental
mechanism and
for a dep th of 11 km is
depth
is
mechanism
station
focal
km.
at
3.7 x
similar to those
to
15
is
inversion
focal
The
SH waves.
moment
Due to the poor
trade-off
t-test
seismic
Both
halfspace.
for the
the
km,
is
(333/ 55/276).
ridge axis.
seismic moment
1 .6
4.14)
is
All
3.5
x
a nearly
The
noda 1 planes
obs erved wave
1024
dyn-cm,
th
a nd
the
inferr ed water
secondary minimum in
the
residu al -versu s-depth cu
6 km, just beneath the Moho.
(319/46/257)
depth
pure
is
are
wel 1
forms ar
e centro d
3.4 km.
A
rve occu s at
The focal mechanism at this depth
has slightly more strike-sl ip component than the
best solution.
This deeper sol ution is worsened when a more
147
gradual
Moho is
used in
the
i
at a
range
within the uncertainty
nversion; however, it is still
= 0. 1.
probability of making a type II error, B,
Tha t
the 6 km solution.
not detect a difference
in 3 out
is,
of
At a = 0.1,
is
equal
0.6 for
to
t-test will
5 times the
we use
If
bet ween the two solutions.
a = 0.15, however, then the
the
difference between the mantle
solution and the best-fitti ng solution is significant.
reason, we suggest
that cen troid depth lies
at a significance level
January 3,
This
of
0.15.
event occurred ju st north of the equator on the
earthquake is
(Figure
The station coverage for this
4.1 3).
excellent for
both P and SH waves.
waveform inversion we obtai ned a nearly pure
(Figu re 4.15).
by the synthetics
8.8 x 1024 dyn-cm, the cent roid depth
inferred water depth
shallowest in
From the
normal-faulting
Al 1 P and SH waveforms are very well
mechanism (322/42/254).
epicenter
in the range 1-2 km
1980
Carlsberg Ridge
fit
For this
is
3.5
this study.
km.
The seismic moment is
is
1.1 km,
and the
This earthquake is
the
The water depth suggests that the
of this earthquak e is a little to the south
of the
ISC
location.
The shallowness of this earthquake is reflected in the lack
of dilatational first motions
PRE, and WIN.
in the P waves at SHL,
The focal mechanism is
GRM,
quite stable for centroid
depths fixed in the range 1-10 km; strike, dip,
varied less than 150.
HKC,
and slip all
The range in acceptable centroid depth at
a = 0.1 is 0.5-4 km and 6-9 km.
A secondary solution occurs at
a depth of 7 km; however, this solution becomes degraded and can
148
be rejected when a gradual Moho transi tion is
We therefore believe that
inversion.
used
the depth
in the
range
0.5
to
4 km reflects the true unce rtainty in centroid depth.
October 7, 1981
October 7, 1981 earthquake occurred on the Central
The
Indian Ri dge about 100 km south of the Vema Fracture Zone
(Figure
Whether the epicenter lies along a short
4.4).
median-va lley segment or is within the Indian plate is ambiguous
on the ba sis of seismicity and bathymetry data alone (Figure
4.4).
Using a semi-automated centroid tensor inversion and data
from the
GDSN, Dziewonski and Woodhouse [1983] obtained a
mechanism combining normal
and strike-slip faulting
Station coverage for P waves is rather poor for
(356/44/3 18).
this even t.
We obtained a normal-faulting mechanism
(311/50/2 57)
with the T axis roughly perpendicular to the ridge
axis.
Most of the waveforms are well fit by the synthetic
seismogra ms
(Figure 4.16).
The seismic moment of 4.5 x 1024
dyn-cm es timated in our inversion is
in excellent agreement
with
the momen t of 4.7 x 1024 dyn-cm obtained by Dziewonski and
Woodhouse [1983].
The centroid depth is 8.2 km, and the
inferred water depth is 3.4 km.
This
earthquake is the deepest we have studied.
There is a
significa nt trade-off, however, between focal mechanism and
centroid
depth because of the poor P-wave coverage.
A secondary
minimum in the residual-versus-depth curve occurs at 1.9 km with
a normal- faulting focal
the best
solution and is
and Woodhouse
[1983];
mechanism (336/56/282)
that differs from
closer to the solution of Dziewonski
the seismic moment for this second
149
Most
solution is 4.8 x 1024 dyn-cm.
quite well
(Figure
source
fit
of
the waveforms
by the synthetic seismograms
4.17).
With a gradual
region structure,
mantle solution is
for this
M oho transition
also
solution
included
in
the
both so lutions become degraded but the
still the bes t solution.
Because of
station coverage, however, we do not have confidence
centroid depth.
are
The centroid de pth is
regarded
as
in
the poor
our
lying
between
1 and 9 km.
October
25,
This
1982
earthquake occurred on the Carlsberg Ridge just
of a fracture zone (Figure 4.6).
Direc t ly south of this
south
area,
the Carlsberg Ridge changes its directi 0 n from NW-SE to almost
N-S.
A centroid moment-tensor solution (293/73/260)
Dziewonski et al.
[1983a] shows normal
steeply-dipping nodal
(323/60/292)
shown
plane.
faulting with
Our inver sion solution
in Figure 4.18 has a greater strike-slip
component than that of Dziewonski et al
Station
this event is good, and all waveforms a re well fit
synthetic seismograms, especially the nodal
waveform at SHL.
one
coverage
for
by the
character of
the P
The seismic moment of 4.1 x 1024 dyn-cm is
somewhat smaller than the estimate of 7.4 x 1024 dyn-cm obtained
by Dziewonski et al.
[1983a].
The cent roid depth is 2.9 km,
the inferred water depth is 3.2 km.
Neither
epicenters nor the bathymetric map clea rly
valley in this area.
(October 18,
earthquake
define the median
However, the wate r depth suggests that
this earthquake may have occurred a lit tle
ISC location.
the
and
to
the north
of
We also investigated the nearby earthquake
1964) inferred by Banghar and Sykes [19691 to
the
150
involve
primarily
normal
faulting, but
because
it
double event we were
unable to obtain
a reliable
using the procedures
we
have
followed
here
acceptable
centroid
depth
for
is
a complex
centroid depth
other
ridge
axis
earthquakes.
The
range
earthquake
the sea
are
(324/57/297)
6-km-deep
2.9
be
rejected
km.
for the
occurs
just
residuals
approximately
December 8,
western
the
of Aden
of centroid
that -of
similar
(Figure
Dziewonski
The
to the
[1983a].
is
depth
2.2
et
al.
seismic
3.2 x
km.
on
to
2.9
focal
Moho
this
is
those
beneath
km
and
mechanism
for the
of the solution
becomes degraded
transition
earthquake
between
the West
4.19).
moment-tensor
1 and
moment
Dziewonski
is
lies
and can
assumed
in
the
6 km.
Figure
of
is
but
4.20
has
3.4 x
[1981]
4.2
km,
found
ridge-axis
Ridge in
et
al.
(288/51/272),
the
[1983a]
using
the
Our inversion
is
a
very
small
1024
dyn-cm obtained
depth
Pearce
(mb = 4.7)
in
Sheba
inversion.
[1983a]
1024
The centroid
for a small
the east.
a gradual
faulting mechanism
(290/56/254) shown
component.
at
1982
residual-
6 km
(3.9 x 1024 dyn-cm)
We believe that
centroid
the
at
the
October
1982
a normal
solution
Moho,
Also,
identical
earthquake occurred
Gulf
obtained
equal.
almost
at a = 0.1 when
in
for the solutions
seismic moment
inversion.
This
below the
However, the mantle solution
and that
method
A secondary minimum
The
and
for the
7 km.
solution are
at
depth
curve
floor.
6 km depth
crust
1 to
is
versus-depth
in
by
earthquake
strike-slip
dyn-cm
is
very
Dziewonski
and the
a focal
similar to
et
inferred
depth of
about
300
al.
water
5 + 2 km
km to
151
Although the number of stations used for the inversion
small,
is
the stations are well-located to give adequate coverage
for both P and SH waves.
depths
is 3 to
residual
The range in acceptable centroid
6 km at a = 0.1.
The plot
versus centroid depth is
very flat,
of mean-squarebut the t-statistic
has a clear minimum in the depth range 3 to 6 km.
This
because the standard deviation of the difference, ad,
is
is very
smal l.
April
1983
11,
This earthquake occurred
Ridge (Fig ure
Carlsberg
is
poor
because many
in
4.6).
the median valley of the
Sta tion coverage for this event
records have not yet been distributed.
From the body wavefo rm inversion, we o)btained a normal
mechanism
4. 21)
(Figure
3040,
subparallel
to the
The seismic
plane.
depth is 2.2 km,
with a sig nifi cant component of
(304/51/239).
strike-slip motion
and
The
nodal plane striking at
ridge axi s, is probably the fault
moment is 3.6 x 1 024 dyn-cm,
the centroid
the inferred wat er depth is 3.5 km.
There is signif icant trade-of f between focal
centroid deoth.
0.5
15 km
km
varies
range
om
i
350
accept
minimum
in
the sea
floor.
momen
t (3.4
similar
the
For
faulting
the centroid
dep
mechanism and
th fixed i n the range
strike varies fr om 2900-3100 , the dip
and the slip var ies from 230
centroid depth s is 0.5-10 km.
residual-versus-depth
0
to
2800.
A secondary
curve occurs at 8 km below
The focal mechanism (303/50/243) and seismic
x 1024 dyn-cm)
The
of the mantle solution are very
to those of the best-fitting solution.
All waveforms
152
are also qu ite well fit by the synthetics for the 8-km-deep
solution.
Using
a
increases slightly
quite small
gradual
Moho
for the
mantle
From the present
transition,
solution,
the
residual
but
the
depth mus t
the centroid
data,
regarded as lying between 0.5 and 10 km.
data
More
is
change
may
be
red uce
the uncerta inty range.
December 8, 1983
This event occurred in
Ridge about 110
1983, event
normal
km southeast
(Figure 4.6).
of the Carlsbe rg
valley
the median
et al.
Dziewonski
[1984]
1 1,
April
of the
of the epicenter
a
obtained
faul ting mechanism with a small strike-slip component
(287/62/249).
The available data
give
coverage, but SH coverage is poor.
reasonably good
P
u sed
of stations
The number
in the inversio n is small because many records have not yet been
distributed.
(Figure
4.22)
The observed waveforms are very well fit
by a normal faulting mechanism (307/45/268) at a
depth of 1.5 km .
The
moment
seismic
of
2.9
x
dyn-cm
1024
is
smaller than the 3.9 x 1024 figure dyn-cm obtained by Dziewo nski
et al.
[1984].
The
inferred
water depth
The residu al-versus-depth plot
at 7 km beneath the sea floor.
However,
show too much dilatational
3.5
km.
shows another
Using the t-test,
range in accept able centroid depths is
between 6 and 9 km.
is
the
between
solutions
first motion
at
local
between
BUL.
m
estima ted
the
1 and
mini mu
2 km an d
6 and
Also,
9
using
km
a
gradual Moho transition the inversion solutions at these depths
become worse than the crustal solution,
on the basis of the t-test.
and they can be rejected
153
DISCUSSION
Indian Ocean under consideration in this
The portion of the
chapter
is quite complicated, and the tectonic history is not
well-understood.
by the
At
least three different plates
ridge system in
this area.
to 22 mm/yr [Minster and Jordan,
of the 11
are separated
rates vary
Spreading
from 6
The source mechanisms
1978].
Indian Ocean ridge earthquakes show
great
similarity.
The seismic moments span only a factor of 3, from 3 to 9 x 1024
dyn-cm.
As with ridge-axis earthquakes
in the north Atlantic,
nearly horizontal
the inferred T axes are all
and are
approximately aligned with t he spreading direction.
the range of centroid depths
is
from 1.1 to 8.2 km,
However,
is
which
three times the range for Mid-Atlantic Ridge earthquakes
(1.5
to
3.0 km).
For most of these earth quakes, th ere are two local mini ma
in the residual-versus-centr oid depth curve,
crust and one near the top of the mant le.
usually degrades and falls outside the
solutions when a gradual
adopted in the inversion.
one within the
The mantle soluti on
range of acceptable
rat her than a sharp Moho transition is
We believe that these mantle
solutions are artifacts of t he simple source velocity model
in our inversion.
used
A summary of the de pth uncertainty for th e
earthquakes examined in this
chapter i s given in Table 4.4.
For only one earthquake (October
7, 1981)
is the best
solution within the mantle, a result t hat holds even for a
gradual
poor
Moho
for this
transition.
Because the station coverage is quite
earthquake, however, we do not have much
154
confidence
occurred
in
in
the stress
centroid
drop
would
that
event
the
bar,
ridge-axis
or that
shallow
this
an
order
to the seafloor,
of magnitude
W e suspect
events.
rupture
actually
earthquake
rupture extended
only 0.7
be
of other
is
If
depth.
the mantle and
smaller than
that this
the
not
did
either
extend to the
seafloor.
In
this
trade-off
between
adequate.
coverage
thesis,
we
mechanism
Generally,
have
have
the
smaller intervals
factor in
long-period
body waves.
similar
Indian
Ocean
to that
depth
if
of
good
permissible centroid
pr obably
Excluding the earthquak es
depths
range
for events
from
1.1
to
along the
4.2
for
km.
i
s not
azimuthal
earthquake centr oid
centroid
consider able
station coverage
station coverage is
determining
inadequate coverage, the
in the
and
earthquakes with
We therefore believe that
important
there is
noticed that
mo
the
depth
u si
with
rid ge-axis
T his
dep th
eve
range
nts
i s
northern Mi d-Atl antic
Ridge.
Applyi ng the same statistical test as for the Atlantic
events,
we have determined the uncertainty ranges of all of our
Indian Ocea n ridge earthquakes.
From the statistical analys is,
we conclude that the range of uncertainty in centroid depth for
earthquakes on Indian Ocean ridges is about ± 4 km.
This va lue
is somewhat larger than that for the Mid-Atlantic Ridge
earthquakes and is mainly the result of poorer station cover age.
It
should be recalled that we have underestimated the number of
degrees of freedom in our analysis in order to be assured of
independent data.
155
The water depths e stimated from the P wa veforms are
comp ared in Table 4.2 w ith the maximum local depth of the median
vall ey as indicated on bathymetric maps.
the two depth estimates is fairly good.
The
Give
agreement between
n the uncertainties
in the epicentral coordinates and in the depths obtained from
the contour map,
most of these earthquakes
beneath the inner f loor
of the med ian valley.
rule out the possi b ility th at some
the rift
the inner walls of
However,
we are su r e that
valley
none of
th e upper
occurred beneath
fault
in this chapter (Ta ble 4.5)
earthquakes
faul t
1966).
these ridge earthquakes
fa ult
width,
fault
earthquakes
two
Excluding the
inadequate
Red
treated
Sea
azimuthal
coverage,
area Lw ranges from 30 t o 170 km 2 .
average fault slip varies
f rom 5 to 60 cm.
20 and 80 bars.
drop lies between
(e.g., August 15,
length,
and ea r thquakes with
the estimated
of these events occurred near
and str ess drop for the
average slip
Again, we may not
mo untains.
rift
We also calcu 1 ated the
area,
seem to have occurred
All
The
these
The
estimated
values
stress
are similar
to those of the Mi d -Atlanti c Ridge earthquak es in Chapter 3.
The March 31,
1969 Red Sea earthquake i s the largest
earthquake in this
study.
nature of the seaf
oor
We have performed
bod
velocity structure
and
(Table 4.3).
best overall
The
fit
At
present,
the
i nterpretation
of
the
ene ath the Red Sea is still disputed.
wa ve inversion with 3 different
fou nd very similar focal mechanisms
inversion solution with oceanic crust has
a nd the deepest
signal-to-noise ra ti os are very
cent roid depth .
high for all
the
The
P waves.
The other
156
Red Sea earthquake (June 28,
of the larger
north
Although station coverage for
1969 event.
its focal
is poor,
the 1972 event
slightly to the
1972) occurred
mechanism and centroid depth
These Red Sea
are similar to those of the 1969 earthquake.
events
are probably the deepest events we have studied.
The half-spreading rate
large depth uncertainty.)
northern Red Sea,
model
[1978],
calculated
is about
in the Gulf of Suez.
important in
Freund
dominated
Picard
by vertical
occurred along the Arava-Dead
how this missing motion
notion,
tectonic movements
McKenzie
et al.
[1970]
slip is
However, Freund
imply
[1969]
and
taken up in
suggested
that the
the Gulf
the Red Sea, and that
the Gulf by normal
[1965] suggested that
is
Combining Picard's
faulting.
folds in the Suez area
some left-lateral motion along the Gulf.
Abdel-Gawad
suggested that the missing motion can be taken up
entirely by
[1976]
being
hypothesized that
of Suez is opening in the same sense as
the missing
is
Sea
[1966] observed that the Gulf of Suez
Gulf is the expression of graben tectonics.
[1966]
190 km.
[1965] has shown from field evidence
that only 110 km of slip has
accommodated.
Girdler [1966]
opening of the Red Sea is about
suggested that the total
The question is
kinematic
this study.
in
The two northern Red Sea earthquakes are
fault.
for the
for the Azores the
slowest spreading rate we have encountered
On the other hand,
a very
from Minster and Jordan's
5.9 mm/yr, except
determining the motion
it has
nominally deeper, but
October 7, 1981 event is
(The
left-lateral
suggested that
motion
in
the Gulf.
Ben-Menahem et
earthquake focal mechanisms
are
al.
157
consistent with the interpretation
Freund
[19651].
Later
focal
of folds
mechan ism
i n the Suez area
studies [Pearce, 1977,
1980] also supported the hypothes is o f left-lateral motion
the Gulf o f Suez.
in
However, among the 4 earthq uakes used to
support th e left-lateral
motion,
one earthquak e occurr ed outside
the Gulf near Cairo, about 300 km awa y from the other
The other 3 earthqu akes all occu rred at the mou th
earthquake s.
of the Gul f
of these
of Suez very close to
(March
31,
eac h other.
1969 and June 28,
1972)
but found that th ey
have nearl y pure normal-faulting mech anisms.
earthquake
(January 12,
body-wave analysis.
with
the
et al.
suggestion
[1970].
1972)
We have studi e d
is too small for
The rema ini ng
long-pe riod
Our results are in qualit ative ag reement
of opening of the Gulf of Suez by McKenzie
However, we cannot
motion also occurring in the Gulf.
rul e out some strike-slip
158
Table 4.1
Epicentral Data for Indian Ocean and Red Sea Ridge-Axis Earthquakes
Date
Origin
Time
March 19, 1964
0942:35.8
14.42
56.36
Dec. 3, 1964
0350:01.7
-14.84
Aug. 15, 1966
1020:47.0
March 31, 1969
Latitude, Longitude,
oN
oE
Half
spreading
ratea, mm/yr
mb
Ms
10.8
5.8
5.7b
66.76
22.1
5.7
6.3 b
3.73
64.00
13.6
5.2
0715:54.4
27.61
33.91
April 22, 1969
2234:40.0
12.80
58.25
Dec. 14,
1837:09.0
8.20
58.49
June 28, 1972
0949:35.0
27.70
33.80
Sept. 7, 1972
0255:00.0
-1.99
68.07
Jan. 3, 1980
1811:56.1
0.02
Oct. 7, 1981
1319:23.0
Oct. 25, 1982
6.1
6.8
5.6
5.2
9.4
5.9
5.6
5.9
5.5
5.5
17.4
5.6
5.7
67.17
16.3
5.7
5.6
-11.40
66.30
20.5
5.9
5.6
1708:29.2
2.96
65.93
14.7
5.3
5.5
Dec. 8, 1982
0619:36.4
12.08
46.08
5.6
5.3
April 11, 1983
1716:30.3
5.25
61.62
11.9
5.7
5.5
Dec. 8, 1983
0126:22.0
4.40
62.51
12.6
5.5
5.1
196 9c
5.9
11.2
9.5
Epicentral and magnitude data are taken from the ISC for 1964-April 1983 and
the NEIS for December 1983.
a Half spreading rate calculated from model RM2 of Minster and Jordan [1978]
b Magnitude M from Rothe [19691
c Near-ridge event.
Table 4.2
Source Mechanisms Obtained from Body Waveform Inversionsa
Centroid
Depthd
Water
Depthe
Depth of
Median Valleyf
240.6±4.3
1.8±0.24
3.3
3.0-3.5
46.4±2.1
248.5±4.1
1.6±0.12
3.4
3.5-4.0
293.6±1.8
50.7+0.7
260.3_+1.8
3.4±0.12
3.8
3.5-4.0
294.0±1.7
36.9±0.4
271.3±1.0
6.2+0.08
0.2*
0.2
4.0± 1.0 349.9±+1.5
63.4±0.9
304.7±2.1
5.1±0.07
4.9
4.5-5.0
Moa
tsc
Strike
Dip
March 19, 1964
4.6+0.5
4.0±1.0
281.9±3.0
38.1±4.7
Dec. 3, 1964
6.8±+0.1
2.4±0.8
301.6_+3. 7
Aug. 15, 1966
4.4+0.3
3.0± 1.0
9.6±+1.6
Date
1969
106.3±3.4
April 22, 1969
2.9±0.3
March 31,
Slip
Dec. 14, 1969
10.3±0.9
3.0± 1.0
317.1+± 1.0
63.1±0.8
61.1±1.0
3.1±0.07
3.0
3.0-3.5
June 28, 1972
3.8_+1.1
4. 5±1.5
288.4±1.2
40.0±0.4
260.0±1.3
6.1±0.04
0.2*
0.2
Sept. 7, 1972
3. 5±0.3
3.0±1.0
332.8±1.8
54.6_+1.3
275.9±2.2
1.6+0.07
3.4
3.5-4.0
Jan. 3, 1980
8.8± 1. 1
4.0±1.0
322.4± 3.8
42.1±2.6
254.3±4.8
1.1±0.22
3.5
3.0-3.5
Oct. 7, 1981
4. 5±0.3
4.0± 1.0
311.2±1.5
50.2±0.4
256.9±1.4
8.2±0.15
3.4
3.5-4.0
Oct. 25, 1982
4.1±0.6
4.0_±1.0
323.1_+1.2
60.4_+1.1
292.2±1.7
2.9+0.10
3.2
3.0-3.5
Dec. 8, 1982
3.4+1.7
5.0+1.0
290.2±3.1.
55.7±1.3
254.1±3.6
4.2+0.32
2.2
2.0-2.5
April 11,
3.4±0.5
5.0± 1.0 305.2+4.0
47.9±1.3
243.9+5.2
2.8±0.17
3.5
3.0-3.5
2.9±0.4
4.0± 1.0 306.9±4.4
44.9+4.1
267.7±4.3
1. 5±0.20
3.5
3.5-4.0
1983
Dec. 8, 1983
160
Notes to Table 4.2
a) The range indicated for source parameters (rMo,
strike, dip, slip,
and centroid depth) is one standard deviation (formal error).
b) x 1024 dyne-cm (1017
N-m)
c) Total duration of the source time function, sec
d) Relative to the seafloor, km
e) Inferred from modeling the water-column reverberations in the later
portions of the P wavetrains.
f) Local depth of the median valley inner floor indicated by
bathymetric maps [Fisher et al., 1982; Laughton, 19751
*
Taken from bathymetric map.
4
Table 4.3
Inversion results for the March 31,
4
1969
northern Red Sea earthquake
Seismic
Source structure
moment, 1024
dyn-cm
Centroid
Double couple
depth,
orientation
km
(strike/dip/slip)
Diuration of
source time
function,
RMS
residual,
mm
sec
1.
Continental crustal
halfspace
106.3
294/37/271
2.
Single oceanic
crustal layer over
mantle halfspace
144.3
297/38/272
3.
15 km continental
crustal layer over
mantle halfspace
89.2
301/39/274
6.2
12.0
9.3
5.36
5.28
5.34
Table 4.4a
Summary of Source Parameter Uncertainty Ranges for Indian Ocean Ridge Earthquakes
Range in centroid
depth, in km, given
by t-test, a = 0.1
Date
March 19, 1964
0.5-13
Secondary solution
with sharp Moho
Within
uncertainty
Depth,
range
for a = 0.1?
km
6.5
yes
Secondary solution
with gradual Moho
Within
uncertainty
Depth,
range
km
for a = 0.1?
6
yes
Adopted
uncertainty
ranges
Depth, km
a
1-4
0.15
Dec. 3, 1964
1-6
6
no
--
1-3
0.15
Aug. 15, 1966
1-6
6
no
4--
2-6
0.10
March 31, 1969
April 22,
1969
Dec. 14, 1969
5.5-8
5.5-8
3-8
7
yes
7
no
2-7
0.10
3-6
0.10
2-7
0.10
5-11
0.15
1-2
0.10
4--
June 28, 1972
5-13
Sept. 7, 1972
1-2 and 6-7
Jan. 3, 1980
0.5-4 and 6-9
6
yes
6.0
no
6
yes
6.0
no
0.5-4
0.10
Oct. 7, 1981
1-9
1.9
yes
1.9
yes
1-9
0.10
Oct. 25, 1982
1-7
6
yes
6
no
1-6
0.10
Dec. 8, 1982
3-6
6
no
3-6
0.10
0.5-10
8
yes
8
yes
0.5-10
0.10
1-2 and 6-9
6
yes
6
no
1-2
0.10
April 11, 1983
Dec. 8, 1983
4--
Table 4.4b
Date
Summary of Source Parameter Uncertainty Ranges for Indian Ocean Ridge Earthquakes
Strike,
deg
Dip,
deg
Slip,
deg
Seismic Moment,
1024 dyn-cm
Source time function
duration, sec
March 19, 1964
280-315
25-45
235-290
4.0-6.0
3-4.5
Dec. 3, 1964
296-307
45-50
240-252
5.0-7.4
2-3.5
Aug. 15, 1966
288-298
49-53
259-264
3.8-4.6
March 31, 1969
294-309
35-40
270-280
81-114
April 22, 1969
340-355
61-65
295-305
2.6-3.3
4
Dec. 14, 1969
313-320
58-64
59-64
9.2-12.6
3
June 28, 1972
280-300
39-45
250-270
3.1-4.9
2-7
Sept. 7, 1972
315-335
47-56
255-276
2.8-3.5
3
Jan. 3, 1980
302-327
34-49
241-255
7.2-10.7
3-4
Oct. 7, 1981
310-336
42-55
256-281
4.3-5.1
3-6
Oct. 25, 1982
320-330
58-64
285-295
3.6-4.2
4-5
Dec. 8, 1982
285-295
50-60
247-255
3.4-3.8
4-5
April 11, 1983
292-310
37-52
235-255
2.7-4.4
1.5-5
Dec. 8, 1983
305-330
35-45
265-305
2.8-3.1
3.5-4.5
3
7.5-10
~
£lr
Table 4.5
Date
Estimated Source Dimensions and Stress Drop for Indian Ocean Ridge Earthquakes
Fault length,a
km
1
March 19, 1964
2
&r
Fault width,b
km
Fault area,
km2
Average
slip, cm
Stress drop,
bars
11
5
60
20
20
Dec. 3, 1964
7
4
30
50
50
3
Aug. 15, 1966
9
9
80
10
6
4
March 31, 1969c
27
17
460
80
13
5
April 22, 1969
12
12
140
5
2
6
Dec. 14, 1969
9
9
80
30
15
7
June 28, 1972c
12
12
160
8
2
8
Sept. 7, 1972
9
5
40
20
20
9
Jan. 3, 1980
12
3
40
60
80
10
Oct. 7,
1 9 8 1d
15
15
220
3
1
11
Oct. 25, 1982
12
8
100
10
5
12
Dec. 8, 1982
15
12
170
5
2
13
April 11, 1983
15
8
120
7
4
14
Dec.
12
4
50
10
10
8,
1983
a
Estimated from duration of source time function.
shear wave velocity.
b
Calculated from the assumptions:
c
Calculated with rigidity and shear velocity corresponding to continental crustal material.
d
Calculated with rigidity and shear velocity corresponding to mantle material.
Rupture velocity is assumed to be 0.8 times
w < X and w < 2.8h.
0(
165
FIGURE
Figure 4.1.
Location of ridge-axis earthquakes in the northwest
Indian
and
Red Sea studied in this chapter
overall
distribution of seismicity (ISC
Ocean
with the
epicenters,
1964
and Pakistan
Figure
Detailed
near the
epicent
and
metric
from the
mb
study
than
ISC
Figure
4.3.
(dashed
qu
Comparis
Bathy-
are from Laughton [1975];
regions
Earthqua ke epicenters are
area projec tions of the 1ower hemisphere;
adra nts are shaded.
on of obser ved (solid lin e)
lor
inversi(on
(equal-area
km,
(Mercator projec tion).
196 4-1982; large symbols denote events
1964 earthquake,
from the
1969
Faul t plane solutions for events of this
equal-
line)
Iran
the East Sheba Ri dge
km are sha ded.
3
for
compressional
regi ons of
ers of the earthquakes of March 19,
in
> 5.4.
are
the
along
,
shown bec ause they are too numerous.
bat hymet ry of
contours,
shallower
with
not
22,
April
Earth quakes in
-198 2).
are
4.2.
1964,
CAPTIONS
and
riod
P and
SH waves
the
focal
mechanis m so lution
plotted
proj( ection).
on the
All
ampl
lower
for
the
syn thetic
Ma rch
19,
obtained
foca 1 hemis phere
itudes are normal ized
an instrument, magnification of 3000; the am plitude scales
correspond to the waveforms that would be o bserved on an
original
seismogram for suc h an instrument.
vertical
lines delimit the portion of each time series,
The two
digitized at 0 .5 second int ervals, used in the inversion.
Symbols for both types of w aves are open ci rcle,
dilatation; solid circle,
c ompression;
cros s,
emergent
166
arrival
For SH waves,
compression corresponds to positive
motion as defined by Aki
Richards [1980].
and
Some first
motions are shown for stations not used in the waveform
inversi on.
The source time function obtained from the
inversi on is also shown.
Bathymetry and seismicity of the Central
Figure 4.4.
Ridge near
1964,
the epicenters
of the earthquakes
and October 7, 1981.
are from Fisher et
al.
[19821];
regions
Comparison of t he obse rved and
Figure 4. 5.
of December 3,
Bathymetric contours,
in km,
shallower than 3 km
See Figure 4 .2 fo r fu rther
are shaded.
Indian
explanat ion
synthetic
wave s for the December 3, 1964 , event.
P and
SH
See Figu re 4.3 for
furt her explanation.
Figure 4. 6.
Bathymetry and
seismicity of the Carlsberg Ridge
median valley near the epice nters of the earthquake s of
August 15, 1966; October 25,
1982; April
y is
Bathym et r:
December 8, 1983.
11,
1983;
and
f rom Laughton [197 5];
regions shallower than 3 km are sh aded.
See
Figure 4.2 for
further explanation.
Figure 4.7.
Comparison
of the observed and synthetic P and SH
waves for the August 15,
1966,
eve nt.
See Figure 4.3
for
further explanation.
Figure 4.8.
Rathymetry and
no rt hern Red Sea.
earthquakes
Gulf
Laughton
Figure
The March
occurred
of Suez.
seismicity
in
31,
of the
and
1969 and June 28,
the Red Sea
Bathymetric contours,
[19751;
central
regions deeper than
4.2 for further explanation.
1972
t the mouth of the
in km, are from
1 km are shaded.
See
167
Figure
4.9.
Comparison
of the
observed
and
waves for the March 31 , 1969, event.
synthetic
P and SH
See Figure 4.3 for
further explanation.
Figure 4.10.
Comparison of the observed and synthetic P and SH
waves for the April
22 , 1969, event.
S ee Figure
4.3 for
further explanation.
Figure 4.11.
Comparison of the observed and synthetic P and SH
waves for the December 14,
1969,
See Figure 4.3 for
event.
further explanation.
Figure 4.12.
Comparison of the observed and synthetic P and
waves for the June 28, 1972, event.
See
Figure 4.3 for
further explanation.
Figure 4.13.
Bathymetry and seismicity of the Carlsberg Ridge
near the epicenters of the earthquakes of September 7,
1972, and January 3, 1980.
[1975]
and Fisher et al.
3 km are shaded.
Figure 4.14.
Bathymetry is from Laughton
[1982];
regions shallower than
See Figure 4.2 for further explanatio n.
Comparison of the observed and synthetic P and
waves for the September 7, 1972, event.
SH
See Figure 4.3 for
further explanation.
Figure 4.15.
Comparison of the observed and synthetic P and
waves for the January 3, 1980, event.
SH
See Figure 4.3 f or
further explanation.
Figure 4.16.
Comparison of observed and synthetic P and SH
waves for the October 7, 1981, event.
for a centroid depth of 8.2 km.
Station NWAO is in the
GDSN; see separate amplitude scale.
further explanation.
The synthetics a re
See Figure 4.3 for
168
Figure 4.17.
waves
Comparison of observed and synthetic
for the
October 7, 1981, event.
for a centroid depth
of 1.9 km.
The
See Figure
P and SH
synthetics are
4.3 for further
explanation.
Figure 4.18.
waves
Comparison of
observed and synthetic
for the October 25,
further
Figure 4.19.
1982, event.
P and SH
See Figure 4.3 for
explanation.
Bathymetry and seismicity of the West
Sheba Ridge
near the epicenter of the earthquake of December 8, 1982.
Bathymetric
regions
contours,
shallower than
in
km,
are from Laughton
3 km are shaded.
(1975);
See Figure 4.2 for
further explanation.
Figure 4.20.
waves
Comparison of the observed and
for the December 8,
1982, event.
synthetic P and
SH
See Figure 4.3 for
further explanation.
Figure 4.21.
waves
Comparison of the observed and synthetic P and SH
for the April
11,
1983, event.
See Figure 4.3 for
further explanation.
Figure 4.22.
waves
Comparison of the observed and synthetic P and SH
for the December 8,
further
explanation.
1983, event.
See Figure 4.3 for
169
30'N
20'N
S 10'N
0*
10'S
-
30'E
40'E
50'E
Figure 4.1
60'E
70'E
20'S
80'E
170
1 5N
i o10,4
55'E
60'E
Figure 4.2
MARCH 19, 1964
P WAVES
SH WAVES
5
SHL
Z
SHL
-
Souce Time Function
p
(sec)
It
4
v-
C HG
ATTRI
CHG
BAG
ATU
BAG
NHA
'
NHA
BUL
5
2
mm
mm
60
60
SOC
Sec
Figure 4.3
172
loS
10
K150S
70 0 E
65E
Figure 4.4
Figure 4.5
8'N
5N
2'N
60E
65E
~_
AUGUST
WAVES
MSII
15,
1966
sllVzt
Sorc
TlI( e
I
ut
i
n
I AB
,
I
B
I
N
NUIl
L I 1A
WIN
PIlE
3
m in
8
/0
s5ec
sec
SOC
Figure 4.7
30'N
25'N
-j
30'E
35'E
20'N
40'E
Figure 4.9
APRIL 22, 1969
SH WAVES
4
Source
KRK
Time )Functon
SHL
(sec)
NUR
CHG
SNG
SDB
BUL
6
mm
60
60
Sec
Figure 4.10
OO
Figure 4.11
Figure 4.12
C),
181
00
5S
65E
70'E
Figure 4.13
SEPTEMBER 7, 1972
P WAVES
OUE
KBL
Source
I\
SHI
.
0I
N
furnct
I1mo
V
(sec)
I
JERill
JER
-JER
NAI
_ -o
PRE
4
mm
mm
6(
40s
sec
Sec
I
Figure 4.14
Figure 4.15
00
Co
4W
Figure 4.16
a
co
Figure 4.17
OCTOBER 25, 1982
ES
SH WAVES
5
CHGG
BAG
SNG
BCAO
SN
AO
CTA
WIN
6
rm
5
WWSSN
mm
2
WWSSN
60
sec
sec
Figure 4.18
mm
GDSN
60
130
sec
co
C0)
15N
45E
50E
Figure 4.19
10'N
a
Figure 4.20
4
APRIL 11, 1983
.
SH WAVES
P WAVES
5
CHG
SHL
Source Time Function
ATU
(sec)
I
SNG
RUL
BU
PRE
2
I"
mm
60
mm
6S
sec
sec
Figure 4.21
DECEMBER 8, 1983
P WAVES
SH WAVES
Source
C HG
Time
Function
(sec)
IST
BAG
BUL
BUL
60
mm
sec
mm
60
sec
Figure 4.22
191
CHAPTER
5.
SOME GENERALIZATIONS ON THE CHARACTERISTICS OF
RIDGE-AXIS
EARTHOIUAKES
INTRODUCTION
In previous chapters, we presented the source mechanisms
of 29 large ridge-a xis
earthqua ke s that occurre,d in the North
Indian Oceans.
Atlantic and
In t his chapter we consider the
implication for ridge tectonics o f the earthqua ke source
mechanism s and focal
depths deter mined
and we in terlpret our results in
t erms
in
previ ous
of what
i
chapters,
s known about
rifting at s low-spreading ridges.
We b egiI n with a retrospectiv e view of the uncertainty in
centroid dep th from long-period b o dy waveform inversion as a
function of azimuthal coverage an d seismic moment.
apply the most
We then
d epths for ridge-axis
reliable centroid
earthquak es to a discussion of th e depth extent of brittle
behavior ben eath slow-spreading m i d-ocean ridges.
In
particula r, we consider the distr i bution of centroid depth and
parameters
other source
of ridge earthquakes versus spreading
rate.
UNCERTAINTY
IN CENTROID
For each
studied with
in
of the 29
body-wave
uncertainty
discussed
Tables
3.4
in
in
4.4.
2.
As
ANOTHER
inversion,
depth
These
LOOK
normal-faulting events
ridge-axis
centroid
Chapter
and
DEPTH:
we
by
have
the
statistical
uncertainty
indicated by
estimated
ranges
the numerical
the
range
methods
are
listed
tests
in
192
Chapter 2, we expect that the uncertainty
should
centroid
in
depth
be generally greatest for events wi th the poorest
azimuthal coverage of P and SH waveforms.
That
this
whi
in
is displayed in F igure 5.1,
expectation is borne out
ch
the range i n acceptab le centr oid depths (Tab les 3.4a and 4.4a)
is
between
plotted versus the greates t azimuthal gap
For all
neighboring stations in the c overage of P wa veforms.
events with a P-wave gap of 1ess than 1200, the range in
acceptable centroid depths is less than 4 km (i.e.,
uncertainty is
about + 2 km).
of less than 1800,
For all
event
the
s with a P-wave gap
the range in acceptable d epths is less than
Because the uncertaint y estimated fro m our statistical
5 km.
generally asymmetri cal about the best-fitting
method is
centroid depth,
the 5 km rang e translates to uncertainty limits
If the range of acce ptable centroid
of about -2 and +4 km.
plotted against the average gap fo r P- and SH-wave
depths is
coverage, a similar relations hip is shown (F igure 5.2).
corresponding but weaker rela tionship
alone.
coverage
of
The
above analysis
reliably
in order to estimate
SH-wave
underscores the
coverage, especially
azimuthal
adequate
for
exists
the centroid
importance
P-wave
for
A
depth o f
coverage,
shallow
earthquakes.
In
is
Figure
plotted
ranges
in
against
the
of this
larger
than
range
are
study.
all
in
uncertainty
in
seismic moment.
uncertainty
earthquakes
moment
5.3,
As
expected, t he
associated with
All
earthquakes
5 x 1024 dyn-cm have
ce ntroid
the
with
ranges
depth
largest
smal 1 er
a seismic
in acceptable
193
centroid
+
depths
2 km).
le ss than
The uncerta inty
earthquakes,
the
of
focal
probably
4 km
range
(or
tends
because more
are
mechanisms
better
uncert
ainties less than
to
smal 1er
be
stations
for
larger
are used so that
constrained.
COMPARISON OF SOURCE MECHANISM WITH HARVARD CMT SOLUTIONS
For the most recent earthquakes of this study, centroid
moment tensor solutions have also been obtained by Dziewonski
and Woodhouse [1983] and Dziewonski et al.
may
be
compared with
the
results
of
comparisons are shown in Table 5.1.
the double-c oup le orientations
the two diff ere nt methods
considering the
in
the CMT
our
[1983a,b, 1984] and
inversion.
For most of these event
and seismic moments
agree quite well,
very long periods
i nve rsion [Dziewonski
These
(> 45 sec)
et al.,
obtained from
particular
of the waves us ed
1981].
One event for
which the re sul 1ts of the two methods disagree significantly
the October 7, 1981
Indian Ocean earthquake.
s,
is
However, a
solution cor res ponding to a secondary minimum in the residua 1versus-depth cu rve has
to that
of the Harvard
a double-couple orientation more simi lar
solution (see Chapter 4).
From the a bove
comparisons, we believe that the Harvard CMT solutions are
generally qu ite good and
more
detailed
are an excellent
starting
point for
study with body waveform modeling.
MODELING WATER COLUMN REVERBERATIONS
The water depth above an earthquake epicenter can be
constrained by the predominant period of water column
reverberations, seen as prominent phases immediately following
194
the P-wave
arrivals
Ward,
[e.g.,
estimated from the P waveform is
maximum local
1979].
The water depth
gene ral
in
agreement
with the
depth of the median valley as indicated
available bathymetric maps.
However
,
the amplitudes of observed
water waves are usually muc h larger than the ampli tudes of
Given the very shallow centroid depth s,
synthetics.
is
qui
possible that fault rupture broke the surface and thus gen
rat
it
larger water waves than pre d icted by the synthetic s.
It
i
al
possible that the median va 1 ley can act to focus water
reverberation s,
the first
producin g la rger amp litudes than p redicted after
W
cyc le of such mot i on.
Anoth e r problem is
that the first
obs erved water
reverberat i on can arrive later than in the synthetic wa veform
(e.g.,
the June 2, 1965 Mid-Atlantic Ri dge earthquake).
earthquake occurs beneath the center of the median vall ey,
If
the
then
the first water reverberation samples the deeper part o f the
median val ley,
while later water wave reve rberations ma y sample
a shallowe r "average" median valley depth.
the first
with
full
layered model
needed
in
adjusted
to provide
a
best overall
fit
P waveform.
Numerical
located
sce nario,
water wave arrives later than the synthetic generated
a flat
to the
In this
modeling
structures
to test
these
of
with
P waves
from
non-pla nar
synthetic
sources
seafloor topography
are
suggestions.
DEPTH RANGE OF BRITTLE BEHAVIOR
The best-determined centroid depths constitute an
important constraint on the depth extent of brittle
behavior in
195
the
axial
regions
of
slow-spreading
implications
of our
necessary to
exclude the
and
the two events
Azores
events
the
rifting
Sea
events
earthquakes
with
of the tests
more
to
1970;
less
in
described
in
data
1980].
continental
is
it
northern
[Searle,
Sea
Red
The
set.
formed
by
Re d
The
crust
also exclude al l
We
depths.
On
2, we include only
The
final
data
set
with which we consider
of faulting at mid-oce!an ridges
12 in the Atlantic and 6 in t,he
All
the
basins
centroid
the
segments,
in
19831.
Chapter
discuss
from our
thinned
Cochran,
reliable
in
occurred
lithosphere
occurred
ridge
region
have
To
the bas
earthq
is
uakes
than two-quadrant azimuthal coverage for both P and
SH waves.
extent
2 earthquakes
likely
probably
al.,
for typical
in the Azores
of Atlantic
[McKenzie et
with
are
results
ridges.
18 earthquakes have
Indian
two
1966 and December 8, 1982)
Indian Ocean.
statistical
The
On the basis oif
the
less
deepest
both
than
or
earthqu
occurred
numerical
tests described e;arlier,
18 earth quakes,
Ocean.
centroid depths
to 4.2 km beneath the seafloo r.
(August 15,
includes
the depth
in
experiments
these centroid
dept
estimated to have uncertainti es of + 3 km for the smalle
equal
akes
the
and
hs are
events
and somewhat lesser uncertain ties for the larger events.
In the interpretation
of these centroid depths,
it
is
important to note that
the ce ntroid depth does not mark the
maximum depth of fault
slip.
If
circular in shape and the sli p
the centroid should lie near
is
the
the
fault
is
rectangula r or
approximately
geometric
Since prominent water-column reverberations
uniform,
center of the
are
evident
then
fault.
in the
196
P waves from these events,
it is reasonable to infer that
rupture extends to the seafloor.
fault
Under these conditions, the
maximum depth of fault slip is twice the centroid depth, or 2 to
8 km below the sea
floor for the earthquakes
Given the uncertainty
in
centroid depth
and the
irregular fault shape or nonuniform slip,
depth
of fault
slip may
of this
study.
possibility of
however, the maximum
have exceeded 8 km for
some of these
events.
Two microearthquake experiments conducted with a
s
ufficient
number of ocean-bottom se nsors to resolve focal depth s [Francis
et al.,
1978; Toomey et ai.,
19851 were
conducted
regions of
in
the Mid-Atlantic Ridge ne ar epicenters of one or more
large earthquakes of this study.
It
is
of
interest
o
to
the distribution of micro earthquake focal depths with
f the
compare
t
he
centroid depths obtained for the large earthquakes.
Francis et al.
[1978 ]
reported
the
locations
of
a number of
microearthquakes recorded during December 1974 by a n etwork of 4
ocean-bottom seismometers near the eastern intersecti on of the
St.
Paul's Fracture Zone
and the Mid-Atlantic
Ridge m edian
valley.
Focal depths con centrated in the ranges 0-1 km and
6-8 km.
Francis et al.
[ 19781 commented particularly o n the
absence of median valley events with focal depths bet ween 1 and
5 km,
which they attribut ed to the presence of
layer highly cracked and
circulation.
weakened by pervas ive
The earthqu ake of June 28, 19 79
intracrustal
rothermal
ure 3.21)
occurred beneath the median valley at the southern end of the
zone of microearthquakes
described by Francis
et al.
[19781.
197
The centroid
depth
of
2.5
km
and the large water-column
reverbe rations shown in the P waveforms
value,
at face
would
i
( Figure
3.22),
If correct,
suggest s that the micr oea rt hquakes in
197 4 may
bottom of the
fol lowe d 4.5 years
lat
the
and
centroid de pth
permit
a
The
in
this
have
er.
Of
course,
uncertainties
the
the microearthquak e focal
of
in
depths
both
also
the two measures of activity.
experimen t
second microea rthquake
approximately 90 km south of the
was
conducted
0
early
N,
Kane Fra cture Zone [Toomey et
Beca use of the larg e network (10 stations) and
knowledge of the local
Detrick,
the
sli p zone of the la rge earthquake that
coinci dence i n depth
1985].
result
outlined
1982 in the Mid-Atlantic Ridg e median valley near 23
al.,
taken
ndi ca te that seismi c slip was concentrated
between the seafloor a nd 5 km depth.
top and
if
1985],
crustal
v elocity
t he focal de pths
structure
of many
were determined t o within + 1 km formal
[Purdy
and
of the microearthquakes
e rror at 95% confidence.
Most of the well- resolved hy poce nters wer e located b etween
8 km beneath the seafloor.
and
Thre e large earthquakes in our
tudy
occurred in the same region in 1962 and 1977 (Figure 3.2).
Their fault plane solutions are quite simi
ar
plane solutions obtained for
f microearthquakes
[Toomey et al.,
probably
1985],
atin g that
bot
responding to the same tecton ic
depths of the large
readily
indic
two clust ers
earthquakes are
i nterpreted as
seafloor and about
sets
p ocess.
1.2 to 2.5 kmin,
indicatin g that
5 km depth.
to
sl i
composite
of
fault
events
are
The centroid
values
extended
most
between the
Most of the 1982 microearth-
quakes thus occurred deeper than this nominal slip zone of the
198
large earthquakes
5 and
20 years
earlier,
but the
uncertainties
in bot h sets of quantities also permit the large earthquake slip
zones to have extended well
earthq uake activity.
into
the depth
Unfortunately,
experi ment has been conducted in
range
of micro-
no microearthquake
the
Indian
Ocean that
would
allow us to compare the focal depths of large earthquakes and
mi croe a rthquakes.
If
th e difference in depth between the s lip zones of
the
large eart hquakes and the regions of most int ense micr oea rthquake acti vity along these two segments of the Mid-Atl ant ic
Ridge is
then two possible explanations present
r eal,
themselves
One possibility
is
that microearthquakes
in
the
median val ley outline a stron g portion of the crust [Kanamori,
19811 between 1-2 and 5-6 km depth,
during large earthquakes
and that
fault
init iates near the base
and propagate s upward to the seafloor.
If
fault area or slip wer e conce ntrated in
the
rupture
of the crust
a large
proportion of
upper crust
for
these large events, th e depth of the centroid would be shallower
than half the maximum depth o f
scenario,
fault
slip.
According
to
this
the microear thquake focal depths provide the most
accurate meas ure of th e maxim um depth of brittle
behavior.
A
second possib ility is based o n the premise that the microearth
quakes at 5-8 km depth mark t he strongest portion of the media
valley faults
barrier [Aki,
In thi s scena rio,
19771 to
the
the do wnward
strong
propagation
during the la rge earthquakes of this study.
microearthquake
activity is
zone
of
acts
fault
as
a
slip
Continued
the result of the large stress
199
differences
remaining
larger events.
very
and
large
has
This
of
of this
are unable
breaks
effects
shortly
unusually
earthquake along
moment
than
Possible
include
a large
1026
barrier
favor the first
short-period
(Mo ~
ridge
and
between them
in
of a rare,
through the
second one.
following
large
zone of the
depth
While we
distinguish
monitoring aftershocks
faulting
study.
slip
the possibility
to exclude the
searching for directivity
an
the
greater centroid
research to
discovering
of
earthquake that
a significantly
scenario, we
the base
scenario admits
ridge-axis
any of the events
avenues
near
waveforms,
event,
dyne-cm)
and
normal-
segments.
CHOICE OF FAIULT PLANES
Generally,
the pref erred fault plane may be determined f rom
a fault-plane solution
aftershocks,
available.
ground
F or
if
independent
breakage,
or geological
ridge-axis earthquakes,
plane can oft en be made
on the
be similar to
trend of
the
local
quakes with
nearly pure
this method
f ails
nodal
data
normal
because the
planes are nearly
criterion
the
such
ridge.
are
choice of
f ault
the
For
f ault
identical.
directions
In
this
for
the
two
we try
section
This ap proach gives us a chance both to test
course,
of
another appro ach based on the predicted plate motio ns
Rfi2.
strike
earth-
ri dge
faulting mechanisms,
strike
trends of
the
structu res
that
the
as
the
f rom
model
relation
between RM2 spreading directions and the geometry of ridge-axis
faulting and to quanti fy the selection of fault plane and
associated parameters
(dip angle, slip
direction).
200
For each
studied, we
the
local
chose
as
first determined
small
the
difference
circle
fault
this
Tables
Atlantic
and
events
5.4 and
angles
The
supporting the
the
observed
for the
rms
small
differences
Azores
and
are,
suggest
strike
For most
center.
The
Indian
that
has
of
Figures
strike
a large
strike
of
directions,
oblique
spreading
average differences
strike
Ocean
respectively,
the
earthquakes,
event
1980]
and
smallest
very clear.
predicted
and theoretical
Atlantic
differences
Ocean
angles
events,
11.80
are
-1.1'
and
8.90.
These
provides
fairly
reliable
estimates
of the strike directions
of major
normal
faults
divergent
plate
RM2
are the
the basis
planes
chosen on
as the fault
agreement with
August
15,
In the
epicentral
to
1966,
and
January
regions
available bathymetric
intersect
Histograms
at
boundaries.
same
of
and
respectively;
RM2
The fault
to
We then
of observed-versus-predicted
[Searle,
spreading
normal
calculations for
respectively.
Indian
observed
the
observed
plane is
1966,
suggestion
the Azores
the
and
July 4,
discrepancy between
+4.10
of nodal
the Atlantic
the
have
pole.
plane with
and
events,
5.5 show a histogram
for
between
nodal
we
direction
spreading
5.3 show these
Ocean
the choice
respectively.
along
the RM2
the
and
earthquakes
horizontal
direction
5.2
Indian
the
about
plane
between
direction.
these
of the normal-faulting
the basis
planes that
local
3,
1980
the
nearest
transform
of
the dip
angles
of the
be
two
chosen
except
Indian Ocean
at
for
on
the
earthquakes.
earthquakes,
the local
fault
agreement with
would
ridge trend
of these
maps,
of
according
ridge axis
appears to
an oblique
angle.
preferred
fault
planes
201
are also shown
+ 5.50
for
in
Figures
Mid-Atlantic
ridge events
in the
5.4
and
5.5.
The average
Ridge earthquakes
Indian
and
45.3
dip is
+ 7.30 for
45.40
Ocean.
STRESS DROP
St ress
drops were estimated in Chapters 3 and 4 for
These stress drops are very sensitive to assumed
earthqu akes.
fault widths.
magni fied.
all
shallow eve nts,
For very
this
is
problem
when we ha ve a cent roid depth of 3 +
For example
U
2 km and assume that fa ulting exten ds to the surface, the
calcu lated stress drop f or a 1-km-w ide fault
that
for a 5-km-wide
fa u lt width.
Figure
width
for the stress dro p are estima ted from
P
centr 0 id
It
is
de pth
clear
and
in
that th e
the erro r
depth;
the errors
in
the
of the soLirce time function
the duration
large error
25 times
5. 6 shows the
calcu lated stress drop plotted vers us centro id
bars
is
ba rs
from b eing drawn fr om this data set .
preventt any conclusions
The bi ggest factors
in t he
large uncer tainties are the assumpt ions that the centroid dept h
is
ha 1 f the maximum depth of fault
and that
the
earthquake
ruptu r ed to the seafloor.
I n an effort
to
minimize the
e ffect
width on th e calculated stress drop
assum ptions concerning
fault
widths
fault s
width
of 2.8
have
a constant
we
of the
tested
First,
km.
assumed fault
two
other
we assumed
Figure
5.7
that
all
shows
infer r ed st ress drop versus centroi d depth under this
ass u mp tion.
No clear trend
next assumed that
the
fault width
all
faults
is hal f
can
be
drawn
from this
plot.
We
have the same aspect ratio and that
the fault length inferred from t he
202
duration
of source
versus centroid
assumption
for
cannot
be
depth
function.
does
(Figure 5.8).
that stress
depth
time
drop
does
ridge-crest
Again,
not yield
a plot
stress
drop
under
this
any clear trend
From these analyses,
not vary
of
we believe either
systematically with centroid
earthquakes
or that
any
such variation
resolved.
SEISMIC MOMENT VS. CENTROID DEPTH
We also tested for a relationship between centroid depth
and
seismic
moment.
adequate azimuthal
We
again
cove rage.
use
only the
For the
lid-Atlantic
earthquakes, these quan tities show no
moment
(Figure 5.10).
and
Combini ng the Atlanti
the inverse correlation appears somew nat
larger ridge-axis earthquakes
centroid depths of 1-2 km,
the centroid depth rang e 1-5 km.
and
Indian Ocean
stronger
on
to s how
centroid
depth
dat
(Figure
1025 dyn-cm)
(Mo
while smal
Ridge
however, appear
an inverse correlation between seismi
All
wtih
correlati
ignificant
The Indi an Ocean event
(Figure 5.9).
earthquakes
a,
5 .11).
have
le r events occur throu gh out
One
possible explanation
that slip during the la rger earthquakes occurs preferentially
the uppermost crust, so that the centro d depths of the larger
earthquakes are biased shallow relative
to the
mid-depth
of
fault rupture.
DEPENDENCE OF
We
local
SOURCE
next consider
spreading
Mid-Atlantic
rate as
rate.
Ridge
predicted
PARAMETERS ON
the
In
events
by
the
SPREADING
RATE
variation of
source parameters
Figure
the
is
RM2
5.12,
centroid depth
plotted against
velocities
the
with
of
half-spreading
of Minster
and
Jordan
203
[1978].
with
The
90%
adequate
A similar
the
azimuthal
wi th
less
Atlantic
and
to shoal
with
This
rate
in
given in
cal bars.
vert
rate for
half-spreading
Figure
we excl ude
If
5.13.
coverage , the
ing spreading rate (Figure
ombi ned
ng tends
5.14).
decre ase with
maximum thickness of the zo ne of brittle
the
ridges.
The maximum depth of
outlines the depth of the
p robably
earthquak e faulting
by
show that the earthquake faul
slow-spr eading
for
indicated
rect evidence for a general
di
of events
centroid depths
depth ve sus
is
d ata
increas
for
adequate azimuthal
than
provides
spreading
behavior
even ts
Indian
ranges
coverage are
of cen troid
plot
India n Ocean
events
confidence
It is
brittle-d
uctile trans
that the
lithosphere is colder and mechanically stro ng er beneath
the
slowest-spreading
for
the
lithos phere
magma
crustal
We
next
spreading
mined
consider
versus
Atlantic
and
(M o
events
> 9 x
largest
where
for
depth,
rates
all
less
than
spreading
spreading
we
rate
data
rate is
major
episodes
of
all
29
shows
that
1969)
about
the
range
occurred
6 mm/yr.
of
seismic
The
combined
the
larger
with
ridges
on
14 mm/yr, while
in
ridge-axis
RM2.
illustrate
local
deter-
a plot
given by
occurred
with
reliably
is more
5.15
as
moment
seismic
include all
sets
rates
(March 31,
of
moment
Figure
1024 dyn-cm)
earthquake
the
seismic
Ocean
because there is more time
between
variation
study.
Incian
likely
volcanism.
half-spreading
half-spreading
occurred
down
and
the
Since
of this
segments
cool
to
the centroid
earthquakes
moment
ridge
injection
rate.
than
ition beneath the ridge axis a
smaller events
1-25 mm/yr.
i n the
Th is
Red
plot
The
Sea
supports
204
the hypothesis
stronger
needed
that
beneath
the lithosphere
slow-spreading
to confirm this
is
colder and
mechanically
However, more data
ridges.
are
hypothesis.
NEAR-RIDGE EARTHQUAKES
In Chapter 4, we noted that a thrust faulting earthquake
1969) occurred at 3 km depth in very young oceanic
14,
(December
lithosphe re
near the northern Car 1sberg Ridge,
just south of a
stri king at about
small
fra
cture zone.
3180,
is
nearly parallel to the 1 ocal t rend of the ridge.
noted
i n
Chapter 4, another event with a similar setting and
occurred on March 29,
mechanism
Solomon
Bergman
a nd
several
n ear-ridge
thrust-fa ulting,
ri
Ocean,
in
occurred
1985x].
[
1984] have noted
erized by
charact
least two (November 25,
epicenters,
about
1965
and May 9
events occurre d in the Pac ific
From ba thymetry
30
the
km from
December 14,
t he median
1969,
val ley
in
and
eart hqua ke
lit hosp here
about
old.
the
evidence
and
Solomon
spreading
near-ridge
the
Stein
li thosphere older than 9 m.y.
Bergman
along
Indian Ridg e (18.880S 9
[1984] and Wi ens and
Both of these
dqe axis.
earthquake
l.y.
1969; a further thrust eve nt
earthquakes which are
at
As
which have P axes orient ed near ly perpe ndicular t o0 the
of
nearby
3
planes,
on November 6, 1984 [Dzi ewonski et al.,
67.350E)
1971)
nodal
) occurred near the Cent ral
6.2
(mb =
One of the
of
cooling
[1 984]
directi on
earthquakes
and
is
showed that a P axis
not
a common
suggested that
oriented
feature for
there is little
a "ridge-push"
compressive stress associated with
and
of oceanic
subsidence
lithosphere [Lister, 1975;
205
Dahlen,
1981 ] to
magnitude
negligible
1981].
to
most
is
shown
when
that,
hydrothermal
and
field
likely
is
large
in magnitude
predicted
Within
to
be
in
the
in
with
in
region
even ts
Bratt
region of
5 to
thrust
direction
of
[Dahlen,
oceanic
co oling
of
is
with P axes
e t al.
where a1
uppermost
push"
milli on years
very young
effect
the
"ridge
taken into accoun t,
the
the
Fu rther,
thrust
stress.
approximate
circulation
M.y.
of a few
cause for
thermoelastic
the
35
associated
ages
predi ct a near-ridge
lithosphere.
are
stress
least
spreading direction
lithosphere
models
of at
a t lithosphere
The
parallel
the
of
ages
is
10
[1985]
due
to
thermal
nearly
km
fau lting,
spreading.
of
have
stress
horizont
the
the
P axes
al
206
Table 5.1
Comparison
body
of
waveform
the Harvard
inversion
CMT
solutions
results
for
and
recent
ridge-crest earthquakes
This
Harvard CMT
solution
i
e i sm i
Date
Double
couple
Double
couple
Seismic
momenta
356/44/318
7 October 1981
study
4.7
Seismic
momenta
311/50/257
336/56/ 2 82 b
4.5
4.8 b
29 January 1982
353/61/246
9/44/264
5.4
25 October 1982
293/73/260
323/60/292
4.1
8 December 1982
288/51/272
290/56/254
3.4
12 May
354/50/259
341/48/251
7.1
287/62/249
307/45/268
2.9
1983
8 December 1983
a
x
1024
dyn
b Second-best
cm
solution
(see
Chapter 4)
207
Table
5.2.
Comparison of observed stri
predicted from plate veloci
Ridge earthquakes
Date
Nodal planes
Strike/Dip
angle with that
for Mid-Atlantic
Observed-predicted
strike angle
1
June 3, 1962
4/43
179/48
-6
-11
2
June 2, 1965
187/57*
335/38
-3
-35
3
Nov.
25/46
211/44
14
21
4
July 4, 1966
135/49
309/41
-30
-36
5
April
20,
1968
328/34
121/59
-11
-37
6
Sept.
20,
1969
23/42
215/49
7
May 31,
8
April
9
16,
1965
1971
209/40
51/52
-1
22
8/48*
212/44
2
26
June 6, 1972
20/52*
227/41
9
36
10
June 28,
1977
1/46
204/46
-9
14
11
June 28,
1977
1/44
201/48
-10
11
12
Jan.
1979
199/44
20/46
16
17
13
April
17/52
210/39
5
19
14
June
28,
1979
346/40*
153/51
15
Jan.
29,
1982
9/44
197/47
-1
7
16
May
188/45
341/48
-2
-29
3,
1972
28,
22,
12,
1979
1983
-9
-22
For each earthquake, the fault plane listed firs
one preferred on the basis of best agreement wit
kinematic model RM2 [Minster and Jordan, 19781.
is the
plate
* indicates fault planes chosen on the basis of agreement
with local ridge trend.
208
Table 5.3.
Comparison of observed strike angle with that
predicted from plate velocity for Indian Ocean
ridge-axis earthquakes
Date
1
March 19,
2
Nodal planes
Strike/Diip
1964
Observed-predicted
strike angle
282/38*
138/58
17
19
Dec. 3, 1964
151/48*"
304/46
6
-25
3
Aug. 15,
129/40
294-51*
0
-15
4
March 31,
1969
112/53
294/37
8
10
5
April 22,
1969
113/43*
350/63
-7
50
6
June 28,
1972
288/40
121/51
4
17
7
Sept. 7, 1972
143/36
333/55
10
20
8
Jan. 3, 1980
322/42
163/ 50*
12
32
9
Oct. 7, 1981
151/42
311/50
7
-14
10
Oct. 25,
323/60
104/36
16
-24
11
Dec. 8, 1982
290/56*
137/37
-13
14
12
April
305/48*
161/48
-4
32
13
Dec.
130/45
307/45
0
-3
1966
1982
11,
8,
1983
1q83
For each earthquake, the fault plane listed first
one prefe rred on the basis of best agreement with
kin ematic model RH12 [tiinster
and Jordan, 1978].
* indicates fault planes chosen on the basis
with local ridge trend.
is the
plate
of agreement
209
FIGURE
Figure 5.1.
CAPTIONS
The greatest azimuthal gap in P waveform coverage
versus the centroid depth uncertainty for all ridge-axis
earthq uakes studied.
Tables
3.4a and 4.4a.
Figure 5.2.
Average azimuthal
P waveform and
depth
Figure
Figure
for all
Seismic moment
all
ridge-axis
5.4.
(Top)
observed
1978]
predicted
Figure
for
5.5.
the
and
(Top)
observed
angles
and
of
dip
strike
strike
togram
ngle
on
the
mi nimi zi ng the
fa
It
the dif ference
by model RM2
d fference
between
to
the
and
[Minster
Ridge ea rthquakes.
of the
the
fault
planes
difference
between
angles.
the
di fference
between
the perpendi cular to the
predicted by model
planes
studied.
depth un ertainty
perpendic ular
showing
and
rthquake
studied.
angles
ri dge-axi s earthquakes.
of
rsus the centroid
ea
centroid
of minimizing
observed
Hi
the
predicted
basis
spreadin g direct
Ocean
versus
Mid-Atlantic
Histogram
selected on
ridge-axis
Histogram showing
strike angle
(Bottom)
(average of the gaps in
earthquakes
spreading direction
Jordan,
gap
SH waveform coverage) ve
uncertainty
5.3.
for
Depth uncertainties are listed in
RM2
for the
India
(Rottom ) Histogram of dip
selected
between
on
the basis
pred icted
and
of
observed
strike angles.
Figure
5.6.
Stress drop versus centroid depth for ridge-axis
earthqu akes.
Error bars for stress drop are calcualted
210
from the errors
3.2
and
4.2)
Only events
for both
Figure
5.7.
in
and centroid
with
P and
drop
of 2
km.
Figure
5.8.
fault
source
Figure
greatest
Ridge
azimuthal
centroid
3.4b).
than
less
1800
is
centroid
(Tables
is
constant
depth unde
the
than
r the
length.
fa ult
on of the
and 4.2).
centroid
earthquakes.
km.
of faulting
extent
half
2.8
at
from the durati
3.2
versus
less
depth unde r the
a vertical
calculated
gap
gap
a nd
depth
for
Only
events w ith
1800
for
both
a
P and SH
are plotted.
waves
Figure
is
function
Mid-Atlantic
azimuthal
to
versus
Seismic moment
5.9.
(Tables
the fault width
length
time
3.4a
versus
drop
assumption that
The
depth
(Tables
are used.
corresponds
Stress
durati on
the fault width
assumption that
a width
time function
a greatest
SH waves
Stress
Such
source
5.10.
Seismi c moment versus centroid depth for Indian
rid ge-axi
Ocean
azimuthal
s earthquakes.
Only
events with
a greatest
gap 1ess than 1800 f or both P and SH waves are
used.
Figure
5.11.
combined
Seismic moment
versus centroid depth for t he
data sets from Atlant ic and Indian Ocean ridges.
5.12.
Centroid depth versus
half spreading rate
Mid-Atlan
tic ridge earthquakes
The solid circles
Figure
represent
earthquakes with a g reatest azimuthal
than
for
1800
both
P and
SH
represents the uncertainty
confidence.
waves.
for
gap
less
The vertical bar
in the centroid depth
at 90%
211
Figure
5.13.
Centroid
Indian Ocean
5.14.
combined
Only
29
Centroid
depth
data
from
set
events with
for both
Figure
ridge-axis
versus
half
spreading
earthquakes.
rate
See Figure
for
5.12
for
explanation.
further
Figure
depth
5.15.
P and
ridge-axis
Atlantic
a greatest
SH waves
Seismic
versus
moment
are
and
rate for
Indian
azimuthal
gap
Ocean
less
the
ridges.
than
1800
included.
versus
earthquakes
spreading
half
of this
spreading
study
are
rate.
included.
All
212
300r
I
-
i
i
.
I
I
I
250
20d0
0-
1500
0
0
0
o100
50
i
i
5
DEPTH UNCERTAINTY, km
Figure 5.1
i
10
213
300'
250"-
200"
150
a.
< 10050
*
5Q
0
0
5
DEPTH UNCERTAINTY, km
Figure 5.2
10
214
106
15
1
-c
z
o
0
10
m
O5
0
0
*
*
5
10
DEPTH UNCERTAINTY, km
Figure 5.3
215
ATLANTIC OCEAN
-20
-30 0
-10
0
0
00
20 0
0 10
10
30
0
STRIKE - ( RM2+90
4
2
i
15
1
" i
I
25
0
•
•i
35
I
55
•
0o
45
DIP
Figure 5.4
I
I
I
i0
65
0
I
I
75
0
216
INDIAN OCEAN
8r
41
CI
II
I- I I
l
w
-30
-2 0
0
ill
-10
0
10
STRIKE-
LL
I,.
!|
0
20
30
65
750
( RM2+90
0
rz
l.I
15
25
35
450
DIP
Figure
5.5
550
4
STRESS DROP, bars
0
I
150
100
50
,
I
i
.' i
E
4
w
0
cr
z
w
6
8
Figure 5.6
200
300
250
I
I
uI
I
STRESS DROP, bars
O,
50
100
150
200
E
H
Fa.
o
a
4
z
4
O
w
F--j
O
FAULT WIDTH -
figure 5.7
2.8km
STRESS DROP, bars
0
I
..I I
50
100
I
I
150
I
I
200
T
T
E
2
03
0
a-
w
o
I-
z
w
4
FAULT LENGTH / FAULT WIDTH -
6
Figure 5.
2
I
ATLANTIC OCEAN
SEISMIC MOMENT, X10
10
Figure 5.9
24
dyn-cm
15
20
INDIAN OCEAN
24
SEISMIC MOMENT,
dyne-cm
x10
10
I
I
I
I
I
I
I
I
I
I
I
E
I
W
H
w
ICIIITTIIIIIIIII111
=
U
U
13
0
W
6)
~I I
I I
I
~
I
lIIIIIl
I
I
20
15
I
I
Figure 5.10
I
IIU
SEISMIC MOMENT, x10
10
E
0
w
O
w1
Figure 5.11
24
dyne-cm
15
20
ATLANTIC OCEAN
HALF SPREADING RATE, mm/yr
5
w
O-
10
hFigure55.12
O
10
Figure 5.12
15
20
25
INDIAN OCEAN
HALF SPREADING RATE, mm/yr
5
0
10
,
15
20
25
,
40.
E
Iaa
0
F-
z
w
100
10
'
a
I
.
.
a
a I
a
a
a
a
Figure 5.13
I
a
a
a
I
a
a
a
a
&
b
HALF SPREADING RATE, mm/yr
0
5
15
i0
*
O
Iw
L-
Figure 5.14
20
*
25
106
15
E
O
_0
o
-
10
z
0
0
o
-
0
•
5
w
-
0
5
15
10
HALF SPREADING RATE,
Figure 5.15
mm/yr
20
25
227
CHAPTER 6.
THESIS SUMMARY AND
SUGGESTIONS
In this
thesis.
chapter we summarize the major
We have
earthquakes
along
Atlantic and
conducted
Indian
earthquakes.
For
coverage,
1800
best
A major
for
for
both
(2)
In
P and
P waves,
largest
depths
is
of
our
in
large
the
study has
been
ridge-axis
very
gap
good
important
for shallow
azimuthal
azimuthal
in
earthquakes.
should
be
less
than
SH waves.
general,
larger
earthquakes
uncertai nties in source parameters,
probably
for
for 29
centers
waveform inversion,
centroid
the
studies
of this
that
long-period body
results
goal
centroid depths
found
reliable
source
results
ridge spreading
Oceans.
particularly
determining
For
We
detailed
mid-ocean
to determine reliable
(1)
FOR FURTHER WORK
have smaller
centroid
particularly
depth,
because more stations are used and the mechanism is
better constrained than for smaller events.
(3) All
mid-ocean
ridge earthqu akes ha e mechanisms
characte rized predominantly by norma 1 fault
ng on fault planes
dipping at angles of about 450.
are all
The
T axes
parallel to
the direction of spreading.
(4)
The water depths estimated from the predominant period
of water -column reverberations in the P wavetrains
the ri dge earthqu akes
inner fl
(5)
0 or
in this
of th e me dian
The cent roid
coverage are all
study
all
occurred
indicate that
beneath
the
valley.
depths of eart hquakes with
very shallow,
good
azimuthal
rangi ng from 1 to 4.2 km.
If
the
228
centroid
depth marks
earthquake faulting
these
be
events.
the mean
depth
extended to
A limited
2-8
amount
of
fault
km
beneath
slip
at
slip,
then
the seafloor
greater depth
for
cannot
excluded.
(6)
The
spreading
greatest
rate.
decrease with
zone of
centroid
This
An
depths
provides direct
spreading
rate
brittle behavior
(7)
example was
in
for
found
of thrust
in
the
greatest
compression
is
parallel
(8)
Red
that
likely
Sea
The
centroid
support
continental
faulting
the
along the
of
earthquakes
for
two
faulting
greater depths
thesis
in
very young
the axis
thermoelastic
earthquakes
seismic
oceanic
we
simulations.
depth.
We mainly
depth
in
the
of
stress
norther
accompanying
than
does
seismic
spreading centers.
lead
to
several
each
of
the very shallow
To
waveform
simulations.
tests.
in this
investigating
carried
uncertainties
precision.
For
have studied
considerable effort
body
faulting
by
depths
this
ridges.
spreading direction.
caused
axes of
of the
to the
was
to
for a general
Indian Ocean;
event
extends
increasing
suggestions
further work:
Numerical
the
northern
view that
rifting
The findings
for
this
with
evidence
slow-spreading
lithosphere
is
shoal
the maximum thickness
(2 m.y.)
It
of
infer
we
the
res
hypothesis
have
obtained
the accuracy obtaina
inverison,
Nabelek
out
thesi
we
[1984]
should
has
carry
carried
A more detailed follow-on study,
s, we have spent
olution of centroid
testing, however, and
are measures of
ble from long-period
out numerical
out some such numeri cal
perh aps using a Monte
229
Carlo
method,
is
necessary to
the method.
There are
the details
of a Monte
provide
document
formalized
the true capability
procedures
Carlo experiment,
better estimates,
suggest
for
of
learning
from
details that may
new measures
of misfit,
or
both.
Extension to
other
regions.
This
study
raises
several
issues concerning the tectonic
nterpretati on of ridge crest
earthquakes,
ionship of microearthqu akes and
incl uding the rel
large earthquakes along a given ridge-axis segment, the
relationship of ea rthquake activity within the inner median
valley to fault st ructures in the adjacent rift
mountains,
and
variations of som e source properties with spreading rate.
To
address these iss
additional
band.
events
es will require extensions of this study to
other regions, and an expanded frequency
Spreading c enters suitable for studies paralleling those
of this thesis in
lude the southern Mid-Atlantic Ridge, the
Southwest Indian 0cean Ridge, the American-Antarctic Ridge, the
Gorda
Ridge,
and
the
Galapagos
Spreading Center.
230
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I
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Broad-b and
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1976.
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250
APPENDIX A.
TESTING HYPOTHESES
INTRODUCTION
a statement or claim about the state
A hypothesi s is
nature.
Scientists often have hypotheses about thei r
particular studies, hypotheses that need substa ntiat io n or
To
verification--or reje ction--for one purpose or anoth er
this end, they gather data and let
doubt on their hypoth esis.
The process
term
is
or cast
sup port
fI t
the
esting of
To acce pt a hypothesis does not m ean t he same
hypotheses."
investig ators or in
thing to all
statistical
the data
experimen t,
accepting
that the hypothesis is "proved"
the data in a sample
all
give
situations
rigorous
any
in
population and can easily be misleading.
to "believe," in the
weak
sense
To
because
e,
n about a
may mean
"a ccept
being willin
of
sens
inforimatio
incomplete
only
not mean
does
a hypothesis
a
In
act,
g to
or to
proceed in some way, as though the hypothesis w ere t rue.
The hypothesis to be tested is
and the set of other
states of
nature
possible for a given experiment
hypothesis.
A null
called
is
hypothesis is
or
models
that
hy'pothesis,
tt ed as
admi
a Itemrn
called the
often
null
the
of
"Ino
at ive
di fference"
in the effects of two treatments, an applicatio n tha t
to the adjective "nul I."
denoted
by Ho
and
the
The
null
alternative
hypothesis
either by
will
HO or
gave rise
u sually
H1 .
be
251
TYPE
I AND TYPE
II ERRORS
The application of a statist ical test may lead to the
wrong conclusion.
Types
I and
II [Folks,
Rejecting Ho when Ho is
II error:
Accepting Ho when Ho is false.
Introducing a little
terminology from hypothesis testing,
=
but this is
Instead, we are forced to settle for controlling
the probabil ity of such errors.
a
of
true.
Naturally, we would like to eliminate both errors,
impossible.
errors
1981]:
Type I error:
Type
in ferences are called
These wrong
let
probability of Type I error = significance level
5 = probability of Type II error
1-8
= power of the hypothesis test
but as the a risk is
We would lik e both a and B to be small,
reduced, the B risk becomes larger; when we make B smaller, a
19811.
tends to bec ome larger [Folks,
In application, it is
good idea to determine not merely whether the given hypothesi
is accepted or rejected at the specified significance level,
but also to determine the smallest significance level
at which the
[Choi,
1978]
based on the
(P valu
hypothesis can be rejected for the given evidenc
This number
enables
significance level
others to make a decision
of their choice.
PROCEDURE
The
components
(1)
Null
assumed
and
alternative
probability
respectively,
in
our
test
are
as follows:
hypotheses
for
a parameter of an
of a hypothesis
distribution
case).
(i.e.,
Ud
=
0 and Pd > 0,
252
or a
(2) a and B,
n (number of samples).
(3)
A test
(4)
A decision
(5)
A random
(6)
Calculation
(7)
A decision either to
statistic.
rule.
sample
hypothesis obtained
statistic
TESTS
and
and
from the
of the
from
statistic.
accept
or
the calculated
the decision
CONCERNING THE
test
population.
reject
the
null
value of the
test
rule.
MEAN OF
DISTRIBUTI ON WITH
A NORMAL
UNKNOWN
VARIANCE
Let
a
X denote
population with
an
random
unknown
from
a normall y-d
variance a2,
whose me an
sample
po or greater than
to be either
The problem
is
known
is to test
o.
=
Ho
po.
stributed
against
HI
Let
x and
statistic
S 2 be the
t
=
X-P
mean
1
> Po
and
variance
follows
then
the
test
a t-distribution with
n-1
degrees
of X;
of
SI/ N
fre
om.
t i
larg er than
cri
cal
reg ion of such a test would he t > ta.
of
e er
ror s for this test
For an intuitively reasonable test, we choose
some
constant;
is
as
a = P (rejecting Ho
= P (t
= P
> ta
(x-Po
1
> t
let
Ho
choose
An
Ho .
follows:
is
true)
)
us assume that
us
assume
that
if
The
asse ssment
o)
Scalculate
, let
In order to calculate B,
otherwise we
H1
where
9 = 91 where
253
u1
>o
H 1 is tru
when
e.
is
B = P (rejecting H1
=
P (t
< ta
(x -
=P
true)
I P
I
< t
P
=
11)
S//N
=P (
< ta)
S/ VN
90-91
S//N
S/JN
=
or1
slJ
1
(t
P
+ ta)
+ ta)
S//N
Note that
(
1-1o)
is a consta nt, making it possible to compute
S//N
8 from the t-table.
Let us now elaborate on
Figure A.1 shows the relation
errors.
null
In this figure,
hypothesis,
Ho,
the
with a t
error, a, is shown on the rig
lower distribution represents
r
distrib
ide
he
increase
of
I
and
is
o10.
distribu
point
with
ses
gr eater
8
but
at
The
the
Type
tion.
than
which
r isk.
II
repre sents
1 of
th is
B.
Type
alternat e hypothesi s,
vertical dashed line represen
the line to the left
ution
value for
ich
line moves to the right, a de
Type
between
with a true value for p of 111
rejected with a risk and acce
between a and
trade-off
Ho
When
The
H1,
110o
The
can be
the
dashed
B increases; moving
and decre ases
B.
I
254
FIGURE
Figure A.1.
Relation between the probability of type
a, and that of type
is
CAPTION
chosen a priori
II error, 8 [Folks,
and 8 is
1981].
I error,
Usually a
calculated after the test.
8
is a function of the difference ul-po and is important in
assessing the power of the test
(see text).
255
Distribution under the
null hypothesis;
H0
= 40
100a%
JAo
(for example)
Distribution under the
alternative hypothesis;
H, : p > po
(for example)
Acceptance
Rejection
region
region
for null hypothesis
Figure A.1
256
APPENDIX
RESOLUTION
In
R2
this
versus
centroid
Appendix
centroid
depth
information
for
The quantities R 2
respectively.
mean-squared
horizontal
indicates
depth
line
in
confidence
shown
plots
the
test
earthquakes
t are
residual
of
is
in
the
in
value
the lower
versus
of the
line
this
at
depth
each
R2
indicates the
and
normalized
about the
depth
when two
85%
and
4.1.
(14),
by the
The
centroid depth
inversion.
centroid
3.1
(12)
time window.
interval
Epicentral
Tables
is
residual
versus
study.
equations
versus
statisic;
t
in
in
body-waveform
of t
statistic
defined
of
confidence
the
plot
plot
of mean-squared
given
waveform over the same
the 90%
the
and
earthquakes
and
The
bar
found
all
the
CENTROID DEPTHS
we compile
depth
for
OF
B.
best-fitting
The horizontal
is
the 90%
horizontal
confidence
lines
value.
are
257
ATLANTIC OCEAN
RIDGE EVENTS
258
JUNE 3, 1962
4
t
2
0
I
I
I
I
i
I
I
i
.4
2
R
.2 -
5
Centroid Depth, km
10
259
JUNE
1965
4
t
2
0
I
I
I
I
I
I
I
I
I
I
.4
2
R
.2
m
S
-
I
5
Centroid Depth, km
10
260
NOVEMBER 16, 1965
4
t
2
0
.5
2
R
.3
II
I
I
5
Centroid Depth, km
1
I
10
4
4
JULY 4, 1966
4-
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
.4
2
R
t
.2
I
I
I....l........I.I.
10
Centroid Depth, km
15
20
N
,,..i
APRIL 20, 1968
2
t
1
0
.4
t
.2
II
I
I
10
Centroid Depth, km
15
20
263
SEPTEMBER 20, 1969
4
t
2
0
I
I
I
I
I
I
I
I
I
I
.4
2
R
.2
-Ht
S
)
I
5
Centroid Depth, km
I
1
10
264
MAY 31, 1971
I
I
i
I
i
I
I
I
!
I
I
I
I
I
I
I
I
I
I
I
I
I
I
4
t
2
0
I
I
I
I
.3
R
2
t
.1
[-1
S1
)
15
5
Centroid Depth, km
I
10
265
APRIL 3, 1972
I
I
I
I
I
I
I
I
I
I
.5 -
i-
t
.3 -
5
Centroid Depth, km
10
266
JUNE 6, 1972
4
t
2
0
I
i
I
I
I
I
I
I
I
I
.4
2
R
.2
1
I
SI
I
I
5
Centroid Depth, km
1
10
267
JUNE 28, 1977
4
t
2
0
1
I
I
I
I
I
I
I
.4
I
2
R
.2
5
Centroid Depth, km
10
JUNE 28, 1977
4
t
2
0
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
A
15
.4
I
I
I
I
I
I
I
I
I
1
5
I
10
I
I
15
Centroid Depth, km
co
269
JANUARY 28, 1979
I
4
t
I
i
I
I
I
I
I
I
I
-
-/
2
0
I
I
I
I
I
II
I
I
I
I
I
I
I
I
I
I
I
.6
.4
)
5
Centroid Depth, km
10
270
APRIL 22
I
I
I
I
I
1979
I
I
I
I
1
I
I
.4
2
R
.2
10
Centroid Depth, km
271
JUNE 28, 1979
4
t
2
0
I
I
I
I
I
.6
.4
SI
I
I
I
5
Centroid Depth, km
10
272
JANUARY 29, 1982
4
t
2
I
I
I
I
I
0
.6
R2
.4
)
5
Centroid Depth, km
10
273
MAY 12 1983
4
t
2
I
I
I
I
I
I
I
I
I
I
0
I
I
I
I
I
I
I
.5
R2
I
.3
-
t
-
I
5
I
I
I
Centroid Depth, km
I
I
10
274
INDIAN
OCEAN
RIDGE EVENTS
MARCH 19, 1964
2
t
1
0
I
.3
I
I
I
I
I
I
i
I
I
I
I
- I
1
I
2
R
m
,
iI
i
I
I
ai
I
10
Centroid Depth, km
i
i
I
a
a
15
276
DECEMBER 3, 1964
4
t
2
0
SI
.5
I
I
I
I
1
I
-
I
2
R
5
5
Centroid Depth, km
I
I
I
10
277
AUGUST 15, 1966
4
t
2
0
I
I
.5
2
R
.3
0
5
Centroid Depth, km
10
ar
4
4
4
MARCH 31, 1969
5
t
3
1
I
I........
I
I
i
I
a
..
I
I
a
a
I
i
I
i
I
I
.6
.4
i
a
a
a
10
Centroid Depth, km
I
I
I
I
I
I
.1
I
20
279
APRIL 22, 1969
4
t
2
0
I---
I
I
I
I
I
I
I
.4
2
R
.2 F
II
p
I
5
Centroid Depth, km
10
10
280
DECEMBER 14, 1969
4
t
2
0
.3
2
R
I
I
.1
S5
I
l
l5
l
i
l
Centroid Depth, km
l
10
JUNE 28, 1972
1
I
_
I
1
I
_
I
I
I
I
I
_
_
I
_
I
_
_
I
_
I
_
I
_
I
_
I
_
I
_
I
I
I
I
I
I
I
I
_
I
\
.4
.2 1
I
I
I
I
15
Centroid Depth, km
I
20
282
SEPTEMBER 7 1972
4
ii
-
2
,
I
I
I
I
I
I
!
I
I
I
I
I
I
0
I
I
II
I
.6
2
0
R
I
51I
i
I
I
I
.4
)
5
Centroid Depth, km
10
-
283
JANUARY 3 ,1980
0
.4
2
R
.21
)
5
Centroid Depth, km
10
OCTOBER 7, 1981
4
S
0
S
S
S
p
£
I
I
I
I
I
I
I
I
I
I
I
I
I
I
I
0
.4
2
R
1
.2
0
5
10
Centroid Depth, km
15
I
20
285
OCTOBER 25, 1982
0
.3
2
R
I
I
5
Centroid Depth, km
10
286
DECEMBER 8, 1982
i
I
I
I
I
I
I
I
I
2
R
I
)
I
5
Centroid Depth, km
10
4P
APRIL 11, 1983
4
2
0
I
.4
I
I
I
I
I
I
I
I
I
I
I
2
R
.2
SI
I
I
I
I
I
I
I
10
Centroid Depth, km
I
20
4
288
DECEMBER 3 1983
SI
I
I
I
I
I
4
t
2
/1
0
SI
I
I
I
I
I
I
I
I
I
I
I
I
I
.4F
2
R
.2
i
!-
~-----f
5
Centroid Depth, km
10
289
APPENDIX C.
COMPARISON
OF
ALTERNATIVE
In
this
seismograms
March
31,
SOURCE VELOCITY
Appendix we compare
for
the July 4,
1969 Red
structures,
INVERSION SOLUTIONS WITH
Sea
earthquake
as described
respectively.
1966
in
STRUCTURES
the observed
Azores
using
the text
and synthetic
earthquake
alternative
of Chapters
and
the
source
3 and
4,
290
FIGURE
CAPTIONS
Comparison of observed (solid line) long-period P
Figure C.1.
and SH waves
for the July 4, 1966, earthquake with
synthetic wavefo rms (dash ed line) from the secondary
solution (10.8
km centroi d d epth) for a source structure
consisting of a uniform c rus t
over
a mantle
halfspace.
The focal mechan ism solut ion obtained from the inversion
is plotted on th e lower f oca 1 hemisphere (equal-area
projection).
Al 1 amplitu des
instrument magni fication
of
are
normalized to
3000;
the
an
amplitude scales
correspond to the wavefor ms that would be observed on an
original seismog ram from su c h an instrument.
vertical
two
lines delimit th e p ortion of each time series
used in the inve rsion.
Symb ols
are open circle, dilatati
cross,
The
emergent
on;
arrival.
for
both
solid circle,
Fo r SH
waves,
types
of waves
compression;
compression
corresponds to positive moti on as defined by Aki
Richards [1980].
The
sou
and
rce time function obtained from
the inversion is also sho wn.
The best-fitting
solution
for this source structure is shown in Figure 3.9.
Figure C.2.
A,
1966,
Comparis on of ohs erv ed P and SH waves for the July
earthqu ake with
best-fitting sol ution
syn thetic waveforms
(6. 3 km centroid
depth)
from the
when
the
source structure consists of a a 2-layer crust over a
mantle halfspace [Searle, 1976].
further explanat ion.
See Figure
C.1 for
291
Figure C.3.
4,
Comparison of observed P and SH waves for the July
1966, earthquake with synthetic wave forms from the
secondary solution
(10.8
km centroid de pth)
structure consisting of a 2-layer crust
halfspace.
Figure C.4.
4,
See Figure C.1
Comparison
1966,
over a mantle
explanation.
of observed P and SH waves for the July
earthquake with synthetic wave forms from the
best-fitting solution
crustal
for further
a source
for
(3.7
halfspace structure
km centroid depth)
model.
See
Figure
for a
C.1
for
further explanation.
Figure C.5.
Comparison
4, 1966,
of observed
P and SH waves
for
the July
earthquake with synthetic wave forms from the
secondary solution
(7.9 km centroid dep th)
halfspace structure model.
See Figure C.1
for a crustal
for
further
explanation.
Figure C.6.
Comparison of observed P and SH waves for
July 4, 1966,
earthquake with synthetic
best-fitting solution (6.8 km centroid
halfspace structure model.
the
wavefo rms
from the
depth) for
a mantle
See Figure C.1 for fu rther
explanation.
Figure C.7.
Comparison of observed P and SH waves for the
July 4, 1966, earthquake with
secondary
solution
synthetic
waveforms from the
(11.2 km centroid de pth)
halfspace structure model.
See Figure C.1
for a mantle
for further
explanation.
Figure C.8.
Comparison of observed long-per iod
for the March
31,
1969, earthquake with
P and SH waves
synthetic
292
waveforms
from the best-fitting solutio n for a source
structure consisting of a 15-km-thick c rust
halfspace.
Figure C.9.
See Figure C.1 for further
over a mantle
explanation.
Comparison of the observed P an d SH waves for the
March 31,
1969 earthquake with
syntheti c waveforms
from
the best-fitting solution for an oceani c structure model.
See Figure C.1 for further
explanation.
JULY 4, 1966 ( 1-layer crust, secondary minimum )
S H WAVES
P WAVES
3
NOR
Source Time Function
KI
KON
C M:
[107
BOZ
/
(sec)
NU
SCP
COP
IST
[FC0
ATH
BEC
NAI
SDB
BOG
G
6 SDBJ
mm
60
i
I
Sec
Figure
I
I
I
sec
C.1
I
I
Figure
C.2
a
46
P WAVES
\k NOR
VVG/
JULY 4, 1966 ( 2-layer crust, secondary minimum )
S
KEV
Source Time Function
KON
-
C
(sec)
NURn
WES
-
-
IST
o
BEC
ATH
EC
SDB
SJG
BOG
'-
SCP
6
mm
60
sec
sec
L
Figure C.3
<0
4
&
Figure C.4
alll
4
A*4
40
a
Figure C.5
I-4
Figure
C.6
Figure C.7
°k
&
*
Figure C.8
Figure C.9