2017, Attiya Darensburg - TDL

Global mapping of the 410 km upper mantle boundary using PP precursor waves

by

Attiya Darensburg, M.S.

A Thesis

In

Geoscience

Submitted to the Graduate Faculty

of Texas Tech University in

Partial Fulfillment of

the Requirements for

the Degree of

Master of Science

Approved

Dr. Harold Gurrola

Chair of Committee

Dr. Hal Karlsson

Dr. George Asquith

Mark Sheridan

Dean of the Graduate School

August, 2017

Texas Tech University, Attiya Darensburg, August 2017

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ACKNOWLEDGMENTS

I would like to extend my gratitude toward my advisor, Dr. Gurrola for his patience

and guidance throughout my academic endeavors at Texas Tech University. I would

also like to thank him especially for his immense guidance with MATLAB and

troubleshooting. I have enjoyed being his student a great deal and I have flourished

under his mentorship. I have also enjoyed the casual conversations about current

events with heaping doses of humorous commentary. I would also like to take the time

out to thank the other members of my committee, Dr. Asquith and Dr. Karlsson, for

graciously making themselves available to serve on my defense committee.

Last, but definitely not least, I would like to thank my mother, Barbara Darensburg,

for her support throughout my life and always being there for words of encouragement

and unconditional love. Through the sacrifices she made for me as a child and into

adulthood, I am able to be the woman and scholar that I am today and words really

cannot express how much she means to me. I would like to also thank friends and

family who have been a tremendous source of support throughout my life.


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TABLE OF CONTENTS

ACKNOWLEDGMENTS……………………………….......................................... ii

ABSTRACT ………………………………………………………............................ v

LIST OF TABLES …………………………………………………………………. vi

LIST OF FIGURES ……………………………………………………………….. vii

I. INTRODUCTION ……………………………………………………………….. 1

II. GEOLOGICAL BACKGROUND …………………………………………….. 4

Pyrolite vs Piclogite Mantle ……………………………………........................... 4

Previous studies using seismic waveforms to image discontinuities

at 410, 520, and 660 km ………………………………………………………… 10

III. METHODS …………………………………………………….......................... 23

Data processing …………………………………………………………………. 23

Crustal tests ……………………………………………………………………... 30

IV. RESULTS ……………………………………………………………………… 37

PP precursor functions beneath Hawaii ………………………………………… 43

PP precursor functions beneath Alaska………………………….......................... 46

PP precursor functions beneath Eurasia…………………………………………. 49

PP precursor functions beneath northwestern Europe, Greenland,

and Iceland ……………………………………………………………………… 53

PP precursor functions beneath South America…………………......................... 57

V. DISCUSSION …………………………………………………………………... 61

Analysis of global variations in the 410 km discontinuity using

PP precursors …………………………………………………………………… 61

Hawaii ……………………………………………………………....................... 62

Alaska and the Aleutian Islands …...………………………..…........................... 65

Eurasia and N. India …………………………………………….......................... 68

Greenland and Iceland ……………………………………………...................... 73

VI. CONCLUSION ………………………………………………………………... 77

WORKS CITED …………………………………………………………………… 80


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APPENDICES

A. USER MANUALS …………………………………………………………..... 87

B. QC GUIDE ……………………………………………………………………. 91


v

ABSTRACT

Globally mapping the mineral phase discontinuities in the earth’s upper mantle

will enable us to develop detailed 1D models of the upper mantle discontinuity at 410

km depth using underside reflections from PP precursor waves. The 520 km

discontinuity has proven difficult to image with PP waves due to interference from

sidelobes from reflections off of the underside of discontinuity boundaries at 410 km

and 660 km depth. PP precursors cannot be used to map the 660 km discontinuity

either since the reflection from this boundary does not appear consistently. We have

stacked and binned 25 years-worth of seismic data collected between 1990 and 2015

from stations all over the globe available from the IRIS data management services

(DMC). Using the stacked PP functions, we will assess the behavior of the 410 km

discontinuity in regions where there are significant temperature anomalies based on

the mapped depth of the 410 and amplitude of the pulses reflected from this

discontinuity.


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LIST OF TABLES

2.1. Chemical composition results from three different studies.

Results from Jagoutz et al.[1979]: least depleted ultramafic

xenoliths. Results from Sun[1982]: Komatiite – dunite model.

Results from Green et al.[1979]: harzburgite-MORB model.

(Ringwood,1991) …………………………………………............................. 5

2.2. Estimates of the average compositions of subdivisions of

the mantle and the cosmic abundance of elements expressed

as wt % oxides. (Anderson and Bass,1986) ……………………………......... 6

2.3. P and S wave seismic velocity estimations at 400 km

and 650 km calculated by Kennett [1991]. (Ringwood,1991) ………...……... 8


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LIST OF FIGURES

1.1. PP precursor reflections off of discontinutiy depths.

The variable “d” represents the depth of underside reflections.

(Chambers et al.,2005)……........................................................................ 3

2.1. The illustration of the composition of the Earth’s

upper mantle and lithosphere. (Ringwood, 1991) ……………………...... 6

2.2. Density and depth mineral assemblage plot for pyrolite

with respect to volume fraction. The geotherm near 410 km is

assumed to be around 1400°C, while the geotherm near 660 km

is assumed to be around 1600°C in accordance with the mantle

geotherm of Brown and Shankland [1981]. (Ringwood, 1991) …….......... 7

2.3. Phase diagram between the three major olivine mineral phases

(olivine (α) wasleyite (β), wasleyite (β)

magnesiowustite and perovskite (γ)) at 1600°C between

4 GPa and 22 GPa. The shaded rectangle represents the

approximate area where the estimated chemical composition

of olivine in the upper mantle is ((Mg0.89Fe0.11)2(SiO4)).

(Akaogi, 1989) ………………………………………………………....…. 9

2.4. Phase boundary diagram for a pyrolytic mantle with variations

in temperature relative to depth. (Akaogi et al., 1989) ………………...… 10

2.5. Maps displaying the global topography for the 410 km discontinuity.

(a) The result from using P410P waves. (b) The result from using

S410S waves. (Flanagan and Shear, 1999) ……………………………..... 11

2.6. This diagram displays the effect water has on the thickness

of the transition boundary between the α and β phases of

olivine with water content defined by weight percentage (wt % ).

Notice the broadening effect water has on the thickness

of the discontinuity versus the thickness of the same transition

boundary under dry conditions. (Frost, 2008) ………………………..….. 13

2.7. Stacked traces grouped into regions A, B, C and D for the PP

data set. (Chambers et al., 2005) …………………………...………......... 14

2.8. A) PP precursor stacks from all data and from regions

A, B, C and D. The dashed lines represent 95% confidence

limits for the stack determined through bootstrap resampling.

The red curves are stacks of the synthetic seismogram which

include the reference phase and reflections from 220 km,

410 km and 660 km. C) Display of the depth mapping for


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PP precursor of all data and regions A, B, C and D.

The black bar represents the mean amplitude, the gray and

white bars represent the 95% confidence limits. The width

of each bar represents the degree of uncertainty in

discontinuity depth. (Chambers et al., 2005) …………………………..... 16

2.9a. Map of cross sections A-a and B-b. (Chambers et al., 2005) ………….... 17

2.9b. Tomography of the cross sections labeled in figure 2.11a.

Notice the trend in the location of the P410P beneath

Asia and North America versus the Pacific Ocean.

Earlier arrivals of the P410P pulses beneath the Pacific Ocean

are indicative of a depression in the 410 km boundary and

vice versa for the areas beneath Asia and North America.

(Chambers et al., 2005) ………………………………………………..... 18

2.10. PP and SS precursor reflections off of the underside of the 410 km

discontinuity. Depressions in the 410 are clustered in the western

region of China with elevations located near the subducting

Indian plate boundary. Diamonds represent PP precursors and

circles represent the SS precursors. (Lessing et al., 2014) ……………… 22

2.11. SdS reflection coefficients relative to incidence angles for

the 410 km boundary. The black curve represents olivine

to wadsleyite transition zone for a pyrolytic mantle.

The red curve represents reflection coefficients generated from

a synthetic seismogram using the ak135 mantle model

(Kennet et al., 1995). The blue curve represents reflection

coefficients generated from a synthetic seismogram using

the PREM mantle model (Dziewonski and Anderson, 1981).

The angle of incidence is always measured relative to horizontal …........ 22

3.1. This is a map of the global data obtained from 1990 – 2015

from every available seismic recording station provided by the data

management center (DMC). The black dots represent the midpoints

(bouncepoints) for every seismic wave. The red dots indicate

the locations of each station, with noticeable coverage

throughout the United States. The blue dots represent the

epicenters of the earthquakes. Notice how the majority of these

seismic events are located on or near plate boundaries .………………….. 24

3.2. Illustration of beamforming seismic records by receiver location.

Each event is cross correlated with other PP functions based

on the location of the corresponding receiver location.

The black triangles represent the events that are within the


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search radius for cross correlation. The red triangles are

outside of the search radius and therefore will not be

considered for cross correlation …..………………………………….…. 25

3.3. Steps for simultaneous iterative deconvolution.

Step 1 (a): the raw receiver function is cross correlated with

the estimated source function. Step 2 (b): the largest peak is found

and normalized by autocorrelation of the source function.

Step 3 (c): the largest peak from the cross correlated records is

removed from the cross correlation and added to the computed

receiver function. Step 4 (d): the new computed receiver function

is used to estimate the original data by convolution with the receiver

and source function. Step 5 (e): the convolution is used to replace

data in step 1. Steps 6 through 9( (f) to (i)): the process is repeated

with another iteration starting again at step 1. New peaks are added

to the computed receiver function until the original earth response

is found with all relevant discontinuities from the raw data without

added noise (Rogers, 2013) …………………………………………........ 28

3.4. Illustration of stacking source-receiver pairs by bouncepoint location.

Seismic records are stacked by the location of their bouncepoints

relative to other events within the defined stacking radius

(0.5°,1°,2°,4°,8°, or 12°). The black triangles represent the events

that are within the stacking radius. The red triangles are outside of

the search radius and therefore will not be considered for stacking …….. 30

3.5. P410P amplitudes at 2, 1, 0.5, and 0.25 Hz filter frequencies

for simultaneous deconvolution of 8 synthetic seismograms with

each having a crust of random thickness (between 20 km and 60km).

Amplitudes were normalized by the expected amplitude of the

P410P phase. The horizontal axis is the trial number for the 12 tests …… 32

3.6. Synthetic seismogram using PREM model. Frequency is 2Hz (a)

and 0.25 Hz (b) with a sampling rate of 40sps. The reflection

corresponding to the P410P boundary arrives around 87.5s ……………... 34

3.7. Crustal test with 50 iterations at 2Hz (a), 1Hz (b), 0.5Hz (c),

0.25Hz (d) and 0.125Hz (e). The blue trend-line connects the

mean amplitude for 8 crustal layers at various thickness ranges

from 20km±10km, 30km±10km, 40km±10km, and 50km±10km.

The black asterisks represent individual P410P amplitudes ……………… 35

4.1a. Map of the 410 km discontinuity with respect to depth at 1Hz.

The depth range is between 370 km and 450 km below the Earth’s

surface …………………………………………………………………….. 38


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4.1b. Map of the 410 km discontinuity with respect to amplitude at 1Hz.

The amplitudes displayed are measured relative to the main pulse

of the direct PP arrival by normalization at a range between 0 and 0.06 …... 39

4.1c. Map of the 410 km discontinuity with respect to depth at 0.5Hz.

The depth range is between 370 km and 450 km below the

Earth’s surface …………………………………………………………...…. 39

4.1d. Map of the 410 km discontinuity with respect to amplitude at 0.5Hz.

The amplitudes displayed are measured relative to the main pulse

of the direct PP arrival by normalization at a range between 0 and 0.06 .…... 40

4.1e. Map of the 410 km discontinuity with respect to depth at 0.25Hz.

The depth range is between 370 km and 450 km below the Earth’s

surface ………………………………………………………………………. 40

4.1f. Map of the 410 km discontinuity with respect to amplitude at 0.25Hz.

The amplitudes displayed are measured relative to the main pulse of the

direct PP arrival by normalization at a range between 0 and 0.06 ………...... 41

4.2a. Global distribution of S wave velocity (Vs) perturbations

represented as the percentage of the expected Vs at

400 km depth (~ 4.77 to 4.93 km/s). Unlike Vp, Vs is more

sensitive to temperature variations, as shear wave reflections

vary more immediately when propagating through warmer mantle

environments ………………………………………………………………... 41

4.2b. Global distribution of P wave velocity (Vp) perturbations represented

as the percentage of the expected Vp at 400 km depth (~ 8.91 to 9.13

km/s) ……..………………………………………………………………….. 42

4.3. Depth of the 410 km discontinuity beneath Hawaii at 1Hz (a)

and 0.5Hz (b). The 410 km discontinuity is slightly elevated to a

depth of approximately 400km beneath the island of Hawaii.

The 410 then appears to deepen toward the northeast. The deepest

410 appears around 420 km depth beneath the island of Kaua’i …………… 44

4.4. Velocity perturbation expressed as the percent of the referenced

S wave velocity (Vs) around 400km (~4.77 to 4.93 km/s).

Hawaii is located toward the center of this figure, where the mantle

velocity in proximity to the 410 is between 1% and 1.5% slower

than the reference Vs ………………………………………………………... 45

4.5. Map of P410P amplitudes beneath Hawaii and the immediate


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surrounding area at 1Hz(a) and 0.5Hz(b). The average amplitdue

beneath Hawaii ranges between 0.025 and 0.03

(2.5% to 3% of the main PP pulse) for 1Hz and 0.5Hz ……………………. 46

4.6. Depth of the 410 km discontinuity beneath Alaska at 1Hz(a),

0.5Hz(b), and 0.25Hz(c). 410 depth ranges from 430 km – 440 km

beneath Alaska at 1Hz and 0.5 Hz. The 410 appears shallower at

0.25 Hz ……………………………………................................................... 47

4.7. Vs velocity perturbation beneath Alaska at 400 km depth.

There is a fast Vs anomaly observed beneath the Aleutian Islands

up toward the Alaskan mainland to the north/northeast.

Vs in this region is around 1% faster than expected ……………………….. 48

4.8. Map of P410P amplitudes beneath Alaska at 1Hz(a), 0.5Hz(b),

and 0.25Hz(c). For all images, the P410P amplitude observed

beneath Alaska ranges from 0.02 to 0.03. The highest amplitudes

are observed to the south and southwest of the Alaskan mainland ………… 49

4.9. Depth of the 410 km discontinuity beneath Eurasia at 1Hz(a),

0.5Hz(b), and 0.25Hz(c). At 1 Hz, the 410 appears depressed

beneath eastern Russia at around 420 km depth. The 410 is

more uniformly depressed throughout China and eastern Russia

at 420 km depth with the use of PP functions filtered at 0.5 Hz.

There are considerable depressions observed beneath Lake Bikal

and areas to the east at 1 Hz and 0.5 Hz ……………………….……….…… 51

4.10. Vs velocity perturbation beneath Eurasia at 400 km depth.

The fastest P wave velocities are observed beneath southern Japan

and the Sea of Japan at approximately 2% faster than average at depth ……. 52

4.11. Map of P410P amplitudes beneath Eurasia at 1 Hz(a), 0.5 Hz(b),

and 0.25 Hz(c). The amplitude of the P410P pulse beneath Eurasia

is relatively small for the majority of the area with normalized values

between 0.01 and 0.03. Some of the greatest amplitudes are observed

beneath the region north of Lake Baikal …..................................................... 53

4.12. Depth of the 410 km discontinuity beneath NW Europe, Greenland

and Iceland at 1Hz(a) and 0.5 Hz(b). The 410 appears to be right at

410 km below most of Sweden and Norway, with a slight elevation

to around 420 km depth toward the southeastern region of Sweden.

A similar pattern in the depth of the 410 km discontinuity is observed

beneath the United Kingdom and Ireland with a depth range between

410 km and 420 km. The 410 appears at approximately 450 km beneath

Iceland ………………………………………………………………………. 55


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4.13. Velocity perturbation trends beneath NW Europe, Iceland

and Greenland. Slower S wave velocities are observed

primarily around Iceland where Vs is ~1.5 to 2% slower than

expected at 400 km depth. The faster velocities are present beneath

most of Greenland and NW Europe where Vs is 1% faster than

expected at its fastest ……………………………………………..………… 56

4.14. Map of P410P amplitudes beneath NW Europe, Greenland and

Iceland at 1Hz(a) and 0.5 Hz(b). The P410P pulse beneath Iceland

has a magnitude of approximately 0.03 for both 1 Hz and 0.5 Hz.

The P410P pulse beneath Greenland ranges between 0.01 and 0.02

for both frequencies as well. The amplitude range observed beneath

Finland and Sweden is 0.01 – 0.02 at 0.5Hz, but the amplitude

appears to increase to approximately 0.035 beneath Sweden and parts of

Finland at 1 Hz ………………………………………………………...……. 57

4.15. Depth of the 410 km discontinuity beneath South America at

1 Hz (a) and 0.5 Hz (b). The 410 appears to be depressed

throughout the majority of this region with depth ranges between

420 km and 430 km depth. The cool mantle region to the north

appears to be a data processing error due to the abrupt nature of the

apparent elevation of the 410 km discontinuity. The 410 does appear

to gradually elevate to the east, beneath the south Pacific ………………….. 58

4.16. Velocity perturbation trends beneath South America. Slower S

wave velocities are observed beneath the majority of central

South America where the slower anomaly is ~0.5 to1% faster

than expected at 400 km depth …………………………………………........ 59

4.17. Map of P410P amplitudes beneath South America at 1Hz (a)

and 0.5Hz (b). The average magnitude correlating to the P410P

pulse range between 0.01 and 0.025 throughout most of the

mapped region. A concentration of higher magnitudes are located

along the western coast of this area with a particularly strong P410P

pulse observed in the southwestern region …………………………………. 60

5.1. P-wave tomographic map. The proposed hotspot is seen beneath

Hawaii where the Vp is 0.5% slower than average.

(Nolet et al., 2007) ………………………………………………………….. 63

5.2. Depth profile of Hawaiian islands at 0.5 Hz filter frequency.

Starting from left (southeast) to right (northwest), the 410 appears at

approximately 400 km beneath Hawaii and begins to depress toward

the northwest around Maui …………...…………………………………...... 64


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5.3. Transition zone thickness anomaly beneath south-central Alaska.

TZ thickness variations are relative to the average thickness of

250 km. (Yinshuang et al., 2005) ………………………………………….. 66

5.4. Depth profile beneath southern Alaska at 0.5 Hz filter frequency.

Starting from the east(left), moving to the west(right), the 410

is clearly depressed beneath the entire landform. There also appears

to be a consisnent negative velocity layer directly above the 410,

indicative of partial melt. The 410 is highlighted by the bold dashed

line …………….……………………………………………………………. 67

5.5. Depth profile beneath the western end of the Aleutian Island chain

at 0.5 Hz filter frequency. Starting from the island to the south(left)

and progressively moving to the north toward Russia(right),

a possible artifact from the subducting slab is observed by the

negative P wave discontinuity pulse ……………………………………….. 67

5.6. The detached Asian Lithospheric Mantle subducting beneath the

lithosphere of Tibet. (Kind et al., 2002) …………………………….……… 70

5.7. Depth profile beneath India and Tibet at 0.5 Hz filter frequency.

Starting from the south(left), moving roughly to the northeast(right)

across what would be part of the collisional arc stretching into

southern Tibet ………….………………………………………...…………. 71

5.8. Depth profile beneath northeast China at 0.5 Hz filter frequency.

Starting from the south(left), moving to the north/northwest(right).

The 410 is depressed to a depth of approximately 430 km as we

move across northeastern China. The P410P pulse widths begin to

narrow as the profile moves more inland to the northwest.

The 410 km discontinuity is outlined by the black dashed line ………..…… 73

5.9a. Depth profile beneath Iceland at 0.5 Hz filter frequency.

Starting from the west(left), moving to the east(right) ..…………….……… 74

5.9b. Depth profile beneath Iceland at 0.5 Hz filter frequency.

Starting from the southwest(left), moving to the northeast(right) ….………. 75

5.10. Depth profile beneath Greenland at 0.5Hz filter frequency.

Near the northeastern coast, there is a strong negative pulse is

observed below the P410P positive pulses by using the depth profile ……... 76


1

CHAPTER I

INTRODUCTION

The goal of this study is to gain a better understanding of the 410 km discontinuity

(410), which is the discontinuity that defines the upper mantle transition zone (TZ).

The 410 and the TZ in general is considered to be the result of mineral phase changes.

that is generally considered to be bound by mineral phase changes which occur at 410

km depth. The transition zone begins at 410km depth (~14GPa), as a result of the

transformation of olivine (Mg,Fe)2SiO4 into a denser olivine phase called wadsleyite,

also referred to as β-phase olivine or modified spinel (Frost, 2008). The 520km

(~17.5GPa) discontinuity is believed to be the result of wadsleyite (β-phase olivine)

transforming into ringwoodite, which is also referred to as γ-phase olivine or silicate

spinel (Frost, 2008). The final discontinuity defining the base of the transition zone is

observed at 660km depth (~24 GPa). The manifestation of the 660km discontinuity

occurs when ringwoodite breaks down into two separate chemical components,

perovskite (Mg,Fe)SiO3 and magnesiowustite (Mg,Fe)O (Frost, 2008).

The locations of the three major discontinuities were proposed through various

thermodynamic experiments by synthesizing olivine phases at their corresponding

pressure and temperature boundaries using a diamond cell anvil (see Akaogi et al.,

1989). Katsura and Ito (1989) found that the first olivine phase transition begins

around 400 km depth over a depth interval between 9 km to 17 km, where α-

olivine((Mg0.89,Fe0.11)Si2O4) transforms into β-olivine (wadsleyite) at temperatures

between 1400°C – 1600°C (1673K – 1873K). The Clayperon slope is defined by

dP/dT where the change in pressure is measured with respect to the change in


2

temperature. The estimated Clayperon slope for the 410km discontinuity was found to

be around 2.5 +/- 1 MPa/K (Katsura et.al, 1989). With a positive Clayperon slope, the

410 km transition boundary is believed to be exothermic. Consequently, the

discontinuity at 410 km tends to be deeper in regions where the mantle is hot (i.e.

where mantle plumes are present) and elevated in cold mantle regions (Flanagan and

Shearer, 1999). The discontinuity at 520 km depth is also thought to be exothermic, so

it has depth variations similar to those found for the 410. The phase transition at 660

km is thought to be endothermic, so depth variations as a function of temperature will

be opposite from the trend observed at the 410.

In an effort to study and understand the complex nature of the transition zone by

tomographic imaging, P (primary) waveforms and, in some instances, S (secondary)

waveforms are used to infer depth variations of discontinuity boundaries within the

upper mantle regions with no seismic stations. The PdP phase (a turning ray that

bounces of the bottom of a velocity boundary at depth “d” and then travels through the

mantle a second time to the recording station) are typically used to image the 410 km

discontinuity (P410P). The P660P phase are typically not observed due to long period

reflections at this depth and due to interference of other phases. The wave path of PP

precursors and direct PP waves are shown in figure 1.1. The following sections will

summarize the previous studies conducted to image the upper mantle at the top of the

transition zone (410 km) using PP precursors. As phase transition boundaries are

thought to be the result of mineral phase changes of olivine as a result of variations in

depth and temperature, the most basic information that we have as to the composition

of the upper mantle are through direct observations regarding the composition of the


3

upper mantle from xenoliths in kimberlites or alkali basalts extruded onto the Earth’s

surface (Ringwood, 1991). The known compositions are extrapolated to other mantle

depths by diamond anvil cells and other chemical experiments involving temperature

and pressure changes. As a result, mineral physicists have postulated that the upper

mantle is composed primarily of the mineral olivine with a pyrolytic composition

(Ringwood, 1991).

Figure 1.1. PP precursor reflections off of discontinutiy depths. The variable “d”

represents the depth of underside reflections. (Chambers et al., 2005)

Mineral physicsists found phase changes within a pyrolitic mantle are consistent

with the pressure and temperatures at 410 km, 520 km and 660 km depths (see Akaogi

et al., 1989) which further correlate well with depths found through seismic

observations. The reflection amplitudes of these discontinuities are a function of the

density and velocity contrast across the respective discontinuity (Bina and Helffrich,

1994).


4

CHAPTER II

GEOLOGICAL BACKGROUND

Pyrolite vs. Piclogite Mantle

It is important to understand the differences between hypothesized pyrolytic

and piclogitic mantle composition because these two possible mantle types will act

differently in response to pressure and temperature and have different seismic

properties. It is believed that the mantle is composed mostly of olivine, pyroxene(s),

and garnet. The principle rocks which contain these minerals are peridotite (olivine-

rich pyroxene) and piclogite (garnet-rich pyroxene) (Ringwood, 1991). The

composition of the upper mantle is still a topic of debate, but an overwhelming

majority of seismic and laboratory studies (see Ringwood, 1991, Frost, 2008, Akaogi

et al., 1989 and others) support the hypothesis that the mantle is primarily pyrolytic.

Pressure and temperature dependent experiments were conducted in order to

produce estimates for the chemical composition of the upper mantle relationships

between harzburgite and ancient MORB were investigated in a study conducted by

Green et al. [1971] in an effort to constrain the composition of the parental mantle.

The resulting chemical composition is provided in table 2.1.


5

Table 2.1. Chemical composition results from three different studies. Results from

Jagoutz et al.[1979]: least depleted ultramafic xenoliths. Results from Sun[1982]:

Komatiite – dunite model. Results from Green et al.[1979]: harzburgite-MORB model.

(Ringwood, 1991)

Pyrolite Model Compositions(as percentages)

Jagoutz et al., 1979

Sun, 1982 Green et al., 1979

SiO2 45.13 44.49 45.0

TiO2 0.22 0.22 0.17

Al2O3 3.96 4.3 4.4

Cr2O3 0.46 0.44 0.45

CaO 3.5 3.5 3.4

MgO 38.3 37.97 38.8

FeO 7.82 8.36 7.6

NiO 0.27 0.25 0.26

MnO 0.13 0.14 0.11

Na2O 0.33 0.39 0.4

100Mg/(Mg+Fe) 89.7 89.0 90.1

Experimentally derived density ranges for various mineral assemblages

provided additional evidence for a pyrolytic mantle composition. The mineral

assemblage plot of the various density ranges with respect to volume fraction is

available in figure 2.1. These densities were measured in a uniform zero-pressure

environment in order to establish unbiased density values for the various minerals.


6

Figure 2.1. The illustration of the composition of the Earth’s upper mantle and

lithosphere. (Ringwood, 1991)

One of the earliest studies by Anderson and Bass [1986] supports the theory that the

upper mantle has a piclogitic composition by comparing mantle composition to the

cosmic abundance of elements that likely contributed to chemical constituents of the

mantle.

Table 2.2. Estimates of the average compositions of subdivisions of the mantle and

the cosmic abundance of elements expressed as wt % oxides. (data taken directly from

Anderson and Bass, 1986)

Cosmic composition (%)

Shallow mantle composition (%)

Transition zone composition (%)

Lower mantle composition (%)

MgO 36.6 42.2 24.0 36.8

CaO 2.89 1.92 8.0 2.4

Al2O3 3.67 2.05 8.6 3.4

SiO2 50.8 44.2 47.0 53.2

FeO 6.08 8.92 10.8 4.8

Anderson and Bass [1986] found that the observed amplitude for seismic

reflections were smaller than those predicted for a pyrolite mantle model. In an

attempt to match the observed velocities at 400 km and 650 km, they assumed a

homogeneous mantle with a garnet to clinopyroxene ratio that matches observed


7

velocities in the depth range of 400 km to 550 km. For the lower boundary at 660 km,

they found that the P660P phase velocity spike at this depth could be explained by

piclogite transforming to Mg-perovskite. Ringwood [1991] contradicted Anderson and

Bass [1986] and found that the mantle was better modeled as peridotite with relatively

small localized sections of eclogite dispersed throughout the mantle. Figure 2.1

displays the layered composition of oceanic and continental crust based off of

Ringwood’s hypothesis.

The estimated mineral assemblage for the upper mantle were also derived with

the use of observed P and S wave velocities, which were estimated to an approximate

depth range through data inversion (Ringwood, 1991), and is displayed in table 2.3.

Figure 2.2. Density and depth mineral assemblage plot for pyrolite with respect to

volume fraction. The geotherm near 410 km is assumed to be around 1400°C, while

the geotherm near 660 km is assumed to be around 1600°C in accordance with the

mantle geotherm of Brown and Shankland [1981]. (Ringwood, 1991)


8

Table 2.3. P and S wave seismic velocity estimations at 400 km and 650 km

calculated by Kennett [1991]. (Ringwood, 1991)

Seismic parameter P-waves (km/s) S-waves (km/s)

Velocity changes at 400km discontinuity(%)

2.5 – 5.8 2.8 – 5.7

Velocity changes at 650km discontinuity(%)

3.6 – 7.3 3.0 – 7.5

Velocity gradients between 440 – 650 (km/sec)/km

2 × 10−3 − 5 × 10−3 1.8 × 10−3 − 2.9 × 10−3

Many studies have investigated the mantle properties with relation to the ratio of FE to

Mg in the olivine in the mantle (Akaogi et al. [1991], Katsura and Ito [1989] and Ito

and Takahashi [1989]). Brown and Shankland [1981] show the relationship between

mantle pressure and Mg/Fe ratio in the mantle olivine from Akaogi et al. [1989].

Temperature variations and its effects on the depth of various olivine phase boundaries

is shown in figure 2.4. Between figures 2.3 and 2.4 the estimation of the transition

between α-olivine and β-olivine occurring at ~410 km depth and 14 GPa is reasonable.


9

Figure 2.3. Phase diagram between the three major olivine mineral phases (olivine (α)

wasleyite (β), wasleyite (β) magnesiowustite and perovskite (γ)) at 1600°C

between 4 GPa and 22 GPa. The shaded rectangle represents the approximate area

where the estimated chemical composition of olivine in the upper mantle is

((Mg0.89Fe0.11)2(SiO4)) (Akaogi, 1989).


10

Figure 2.4. Phase boundary diagram for a pyrolytic mantle with variations in

temperature relative to depth. (Akaogi et al., 1989)

Previous studies using seismic waveforms to image discontinuities at 410, 520 and 660 km

Flanagan and Shearer (1998, 1999) stacked P410P and S410S phases from a

global data set and found an average depth for the 410 km discontinuity of 418 – 419

km (see figure 2.5). Regional variations in depth range from 400 km to almost 440

km.


11

Figure 2.5. Maps displaying the global topography for the 410 km discontinuity

(Flanagan and Shear, 1999). (a) The result from using P410P waves. (b) The result

from using S410S waves.

These models found deeper than expected for the 410 beneath the Pacific Ocean and

South America in addition to a consistent elevation of the 410 discontinuity beneath

the Indian Ocean and parts of Africa and Antarctica between the P410P and S410S

topographic maps. Flanagan and Shearer found the 410 discontinuity to be absent in

regions with subduction zones and spreading ridges. Lee and Grand [1996] were also

unable to image the 410 beneath the East Pacific rise. Flanagan and Shearer (1998,

1999) found no correlation between depth variations to the 410 and ocean-continent

regions, ruling out the theory that variations in the depth of the 410 could be due to

differences in lithosphere types.

Chambers et al. [2005] found that variations in the P410P and S410S

reflection amplitudes versus S410S reflection amplitudes could be the result of

a)

b)


12

variations in water content, the presence of partial melting and chemical

heterogeneities within the mantle. The amplitudes of PP and SS precursors do not rely

solely on impedance contrasts between thermochemical boundaries and the density of

each heterogeneous layer, but they do depend on intrinsic attenuation and anisotropy.

Anisotropy will result in differences in the velocity of seismic waves depending on

which mineral crystalline axis the incoming wave travels along. Intrinsic attenuation

occurs as the result of the structure of the mantle and anisotropy of minerals within the

mantle (Chambers et al., 2005). In his study, the differences in amplitudes have been

interpreted to be a result of changes in discontinuity thickness when short period

waveforms are used. Variations in precursor amplitudes using long period data has

been interpreted to be the likely result of changes in impedance contrasts.

Observations by Vidale and Benz [1997] restricted the thickness of the 410 km

discontinuity to approximately 4 km using a high frequency reflection from that

boundary. This approximation was made citing reflected phases are only sensitive to

impedance gradients less than 1/4th of a wavelength. A study conducted by Rost and

Weber [2002] focused on imaging the transistion zone beneath the western region of

the Pacific Ocean. Their results showed that the 410 km discontinuity must be sharper

than 6 km with a minimum impedance contrast of 6.5%. The 410 has also been

observed to be much thicker, particularly beneath continents. Helffrich and Wood

[1996] found that a 5 km linear impedance gradient can result in reflection coefficients

that are similar to a 10 km nonlinear impedance gradient, which is the case for the

olivine transition between the α and β phases (Stixrude, 1997). van der Meijde et al.

[2003] used receiver functions from reflections beneath Europe to study the


13

topography of the 410, where the discontinuity was estimated to be between 25 and 35

km thick. As previously mentioned, variations in the thickness of the 410 km

boundary can be attributed to the amount of water present within the mantle

(Chambers et al. 2005 and Frost, 2008). Water partitions into wadsleyite (β-olivine)

over α-olivine by approximately 10:1 (Wood, 1995), where the β-olivine phase is

more stable at lower pressures. Chen et al. [2002] found that β-olivine is more soluble

in water and therefore corresponds to a phase transition at lower pressures when water

is present. As a result, regions where both α and β phases are stable is broadened and

correlates to an increased presence of water in the mantle. Figure 2.6 shows the effect

that water has on the 410 km discontinuity.

Figure 2.6. This diagram displays the effect water has on the thickness of the

transition boundary between the α and β phases of olivine with water content defined

by weight percentage (wt % ). Notice the broadening effect water has on the thickness

of the discontinuity versus the thickness of the same transition boundary under dry

conditions (Frost, 2008).


14

The broadband data used by Chambers et al. [2005] was sampled at a period of

1 second using IRIS, IDA and USGS seismic networks between 1990 and 1999.

Shallow seismic records, events that are 75 km in depth or less and epicentral

distances of 80° – 140° and 100° – 160° for PP and SS waveforms, respectively, were

used to minimize interference with multiple phases. A Butterworth bandpass filter was

applied to both the PP and SS data sets with intervals between 8s to 75s and 15s to 75s

respectively. A Hilbert transform was then applied to the PP data set and the SS data

set was rotated in order to get the transverse component. Data sets with SNR of 3 or

greater and cross correlation coefficients equal to or greater than 0.6 were kept. The

epicentral distance used for the PP data was 110° and 130° for the SS data. This range

was chosen in order to maximize the reflections from the 410 discontinuity while

minimizing interference from other P wave or S wave phases.

The stacks generated for this study were developed by grouping bouncepoints

by 4 different regions (A, B, C, D) around the world, where regions A and D contain

continental bouncepoints and regions B and C contain oceanic bouncepoints (see

figure 2.7).

Figure 2.7. Stacked traces grouped into regions A, B, C and D for the PP data set

(Chambers et al., 2005).


15

These regions were selected for the purposes of enhancing the quality of the

stacked midpoints without averaging out the lateral variability. From these stacked

results, the 410 km discontinuity arrival was the clearest and largest feature in the data

sets with an arrival time of approximately 80s. The sidelobes of the P410P arrived

around 100 seconds before the main PP pulse. This sidelobe was separated out with a

deconvolution algorithm. Within each of these regions, there was a noticeable

variation in the amplitudes generated from the reflections off of the underside of the

transition zone. These variations where shown using a frequency histogram where the

amplitudes were measured in 1000 bootstrap samples as shown in figure 2.8.


16

Figure 2.8. A) PP precursor stacks from all data and from regions A, B, C and D. The

dashed lines represent 95% confidence limits for the stack determined through

bootstrap resampling. The red curves are stacks of the synthetic seismogram which

include the reference phase and reflections from 220 km, 410 km and 660 km. C)

Display of the depth mapping for PP precursor of all data and regions A, B, C and D.

The width of each bar represents the degree of uncertainty in discontinuity depth.

(Chambers et al., 2005)

The stacked traces for the PP data set shows that the amplitude of the P410P is

higher beneath oceanic regions B and C and lower beneath continental regions A and

D when compared to the amplitudes of the global stack. Figures 2.9a and 2.9b displays

Am

plitu

de relativ

e to m

ain P

P

pulse

Am

plitu

de relativ

e to m

ain P

P

pulse


17

the results of stacks using PP precursors beneath continental and oceanic areas. As

previously mentioned, the 410 discontinuity appears to be shallower beneath

continental regions and deeper beneath oceanic regions as cross sections A-a and B-b.

However, results from the stacked SS precursors does not show much variation in the

location of the 410 discontinuity.

Figure 2.9a. Map of cross sections A-a and B-b. (Chambers et al., 2005)


18

Figure 2.9b. Tomography of the cross sections labeled in figure 2.11a. Notice

the trend in the location of the P410P beneath Asia and North America versus

the Pacific Ocean. Earlier arrivals of the P410P pulses beneath the Pacific

Ocean are indicative of a depression in the 410 km boundary and vice versa for

the areas beneath Asia and North America. (Chambers et al., 2005)

Observations were also made with the amplitudes of the P410P. These

variations were attributed to lateral heterogeneity within the mantle, which would

directly affect the impedance contrast and cause variations in reflection coefficients.

Since heterogeneity of the mantle has been attributed to the possible presence of water

in the mantle, the hydrated olivine phases would directly affect the bulk modulus and

elastic properties of the mantle. Yusa and Inoue [1997] has shown that the addition of

water to pure Mg-wadsleyite can reduce the bulk modulus of the transition zone by

5% to 11% while Jacobsen et al. [2004] showed that the presence of water at 0.1% wt

in γ-olivine can reduce the impedance of PP precursors by 5.8%.


19

Lessing et al. [2014] studied PdP and SdS (focusing on SdS phases) reflections

within the transition zone beneath western China and India. Seismic waveforms with a

maximum deviation of ±5° from the theoretical backazimuth were used in order to

avoid analyzing scattered energy that would arrive at the same time as PP precursors.

Each broadband seismogram was then filtered with a second order Butterworth

bandpass filters with the following corner frequencies: 2s and 20s, 3s and 10s, 5s and

25s, 6s and 50s, 8s and 75s, 15s and 75s, 10s and 100s. These ranges of corner

frequencies were used in order to investigate frequency dependent behavior of the

observed seismic records. Discontinuity depths were derived though a migration

technique used by Thomas and Billen [2009]; Schmerr and Thomas [2011] where

inverse projections from PP and SS precursor pulses to reflection points were

calculated. This process also reduced the size of the Fresnel zone, which improved the

lateral resolution of the stacked results. A 40° by 40° 3D grid was then placed around

theoretical PP/SS bouncepoints with a 1° spacing between each grid. The depth range

for each grid was 0 km to 900 km at a 5 km increment. Travel times for each event

were calculated by ray tracing through the ak135 model (Kennett et al., 1995) from the

source location to the grid point and from the grid point to the receiver location.

Further corrections were made in an effort to account for variations in crustal

structure and lateral heterogeneity of the mantle. SS wave tomography model,

S40RTS (Ritsema et al., 2011), PP wave tomography model, MITP08 (Lee et al.,

2008), and crustal model CRUST2.0 (Bassin et al., 2000) were all used to make

corrections to the travel time calculations. Each seismogram was then subsequently

shifted and stacked by their respective midpoints. The seismograms were stacked


20

through the application of linear and 4th root stacking methods. The integrity of these

stacked results were then tested through bootstrap resampling. The arrays used for this

particular study contained 20 seismograms or less and a simplified bootstrapping

method was performed where two seismic traces were randomly replaced with the

remaining two seismic traces within each vespagram.

The majority of the reflection points for this study were beneath western China

between the Tien Shan Mountains and Eastern Tibet. These seismic events had source

locations near subduction zones from Sumatra to the Banda Sea. There was a total of

36 events that had SS precursor pulses and were subsequently used to study the nature

of the SS/SdS phases in this region along with 68 events for PP precursors. The mantle

beneath western China appears to be warmer than expected as the 410 km

discontinuity appears deepest at 440 km. The observed depths of the 410 would

correlate with temperature differences between 100K – 200K warmer than the average

mantle geotherm (Lessing et al., 2014). This observed depression is also supported by

results from a study conducted by Kosarev et al. [1999], where an upwelling of mantle

material is observed beneath the Tibetan plateau. The proposed cause of this upwelling

was the subduction of the Indian lithosphere. The average resolved depth was

approximately 410 ± 18km with the shallowest boundary depth at 380 km located

beneath the Tien Shen Mountains and the Himalayas north of Bangladesh. This

observation was made with the use of 20 S410S events and 53 P410P events where the

410 boundary is clearly seen.

These reflections were also near plate boundaries and regions of active

subduction. The presence of cooler subducting lithosphere within proximity of these


21

elevations was proposed as the cause of this phenomena, as the subducting Indian

lithosphere reaches the upper mantle, the boundary at 410 km moves to a shallower

depth. These elevated bouncepoints correspond to a mantle temperature that is 200K –

400K cooler than the average mantle geotherm (Lessing et al., 2014). Although this

particular study did not account for the presence of water in the mantle, placed there

by a hydrated subducting slab, the storage capacity of the olivine to wadsleyite

reaction is not as significant as the water storage capacity of the wadsleyite to

ringwoodite at 520 km or the complex transition of ringwoodite to magnesiowuesite

and magnesium perovskite at 660 km. The presence of water near the 410 km

discontinuity can cause a reduction in the amplitude of either P410P or S410S

reflections in addition to causing increased travel times as a result of the presence of

slower mantle velocities. Water would also release at the 410 km boundary in order to

accommodate for the difference in storage capacities between α-olivine and β-olivine,

causing the occurrence of partial melt directly above the 410 discontinuity. Depending

on the thickness of the partial melt layer, negative pulse reflections can cause an

amplification of the sidelobes to the P410P and S410S precursors, resulting in the

appearance of one positive pulse followed by a negative pulse. Since two separate

pulses were not observed in the stacked images for Lessing’s study, it is unlikely that

there is the presence of a partial melt layer above the 410 km discontinuity. A map of

the observations made for the 410 km discontinuity is displayed in Figure 2.10, where

there is a clear depression of the 410 primarily located in western China and elevations

of the 410 are seen within the proximity of the Indian plate, although there is one

outlier displayed near the Tien Shen Mountains just south of Kazakhstan.


22

Figure 2.10. PP and SS precursor reflections off of the underside of the 410 km

discontinuity. Depressions in the 410 are clustered in the western region of China with

elevations located near the subducting Indian plate boundary. Diamonds represent PP

precursors and circles represent the SS precursors. (Lessing et al., 2014)

Figure 2.11. SdS reflection coefficients relative to incidence angles for the 410 km

boundary. The black curve represents olivine to wadsleyite transition zone for a

pyrolytic mantle. The red curve represents reflection coefficients generated from a

synthetic seismogram using the ak135 mantle model (Kennet et al., 1995). The blue

curve represents reflection coefficients generated from a synthetic seismogram using

the PREM mantle model (Dziewonski and Anderson, 1981).


23

CHAPTER III

METHODS

Data Processing

Our study started with data acquisition from the Global Seismic Digital

Network (GSDN) by using the IRIS data management service. After sorting through

the data for useable events we were left with over 300,000 PP waveforms to include

for data processing. These seismic records span a total of 25 years from 1990 to 2015

with earthquake magnitudes 6.0 or greater. An epicentral distance range of 60° – 180°

was used in order to minimize the interference between records from multiple local

events. The magnitude of the PP precursor pulses are measured relative to the

magnitude of the main PP pulse.


24

Figure 3.1. This is a map of the global data obtained from 1990 – 2015 from every

available seismic recording station provided by the data management center (DMC).

The black dots represent the midpoints (bouncepoints) for every seismic wave. The

red dots indicate the locations of each station, with noticeable coverage throughout the

United States. The blue dots represent the epicenters of the earthquakes. Notice how

the majority of these seismic events are located on or near plate boundaries.

Data processing started with sorting all earthquake data into folders (high, low,

and others). Records with signal to noise ratios (SNR) values of 3 or greater were

labeled as “high” SNR events, records with SNR events between 1.5 and 3 were

labeled as “others”, and records with SNR events of 1.5 or lower were considered to

be “low”. The next step in data processing involved hand picking seismic events with

clear PP arrivals. We selected all high SNR events and enough of the “others” to have

at least 30 reference phases for each earthquake. The raw data was selected for its

clear PP phase and a lack of obvious noise. The data was cut to include 200s before

the PP phase and 90s after and then resampled to a uniform 20 sps (samples per

second). Cross correlation was used to find more events categorized as “others” or


25

“low” SNR that would possibly be useable. This was accomplished by keeping any

event with a cross correlation coefficient of 0.6 or greater, with at least 3 reference PP

phases. Waveforms with focal depths greater than 60 km were also eliminated in an

effort to eliminate arrivals from other P wave phases that might interfere with the

P410P. To convert P410P arrival times for all of our selected waveforms to depth, we

raytraced all events by utilizing the 1D PREM velocity model (Dziewonski and

Anderson,1981).

Figure 3.2. Illustration of beamforming seismic records by receiver location. Each

event is cross correlated with other PP functions based on the location of the

corresponding receiver location. The black triangles represent the events that are

within the search radius for cross correlation. The red triangles are outside of the

search radius and therefore will not be considered for cross correlation.

Each waveform was deconvolved by filtered and unfiltered source functions. The

filtered source functions with “wet” bouncepoints are de-oceaned. The term de-

oceaned refers to the process of removing bouncepoint multiples or “ghosts” generated


26

by wave reflections off of the ocean bottom for events that have bouncepoints in

regions beneath water. These ocean bottom reflections result in the appearance of two

parallel horizons where an erroneous shallow reflection is generated from the seismic

wave reflecting off of the bottom of the seafloor. We are interested getting rid of this

shallow “ghost” reflection in an effort not to confuse this horizon with the real image

of the P410P boundary at the correct depth. The data was then beamformed (stacked)

into bins of various radii. All deconvolved waveforms were then stacked relative to

each bin size.

After these waveforms went through another round of quality checks (QC), the

beamformed records were stacked by a process called simultaneous iterative

deconvolution. The simultaneous iterative deconvolution process begins with the cross

correlation of the receiver function with the estimated source function. Receiver-

source pairs with the same bouncepoint locations are aligned and adjusted to the

correct time with moveout correction, which accounts for differences in the velocities

of various crustal layers. Once these source-receiver pairs are stacked, the largest or

peak amplitude is selected, representing the arrival of the PP pulse. Reflections from

the source function are then normalized by the peak amplitude generated from the PP

arrival. The peak amplitude is added to a synthetic seismogram and removed from the

source function. This process is repeated, finding the next peak, typically represented

by the arrival of the P410P and added to the synthetic seismogram. This process

continues ideally until the synthetic seismogram consists of at least P410P and PP

precursor arrivals from shallower discontinuities with other relevant arrivals without

the addition of noise. Differences between the horizontal component of the observed


27

seismogram and the predicted signal derived from the convolution of the iteratively

updated receiver function and the vertical component of the observed seismogram are

then calculated using the method of least squares minimization until differences are

negligible.


28

Figure 3.3. Steps for simultaneous iterative deconvolution. Step 1 (a): the raw receiver

function is cross correlated with the estimated source function. Step 2 (b): the largest

peak is found and normalized by autocorrelation of the source function. Step 3 (c): the

largest peak from the cross correlated records is removed from the cross correlation

and added to the computed receiver function. Step 4 (d): the new computed receiver

function is used to estimate the original data by convolution with the receiver and

source function. Step 5 (e): the convolution is used to replace data in step 1. Steps 6

through 9( (f) to (i)): the process is repeated with another iteration starting again at

step 1. New peaks are added to the computed receiver function until the original earth

response is found with all relevant discontinuities from the raw data without added

noise. (Rogers, 2013)


29

Since data was stacked during the simultaneous iterative deconvolution

procedure, we will call the parameters used for simultaneous iterative deconvolution

stacking parameters as well. After several stacking runs were completed, the best

results came from stacked receiver functions which had a minimum SNR of 3, low

pass filter frequencies of 0.5Hz or 0.25Hz, a minimum of 10 events relative to the

location of the station for each respective receiver function, a maximum

stacking/search radius of 5° relative to the location of each respective station, a 2°

radius spacing between each respective stack, a minimum of 50 events within the set

stacking radius relative to the bouncepoint of each event, and 80 stacking iterations. If

the criteria for minimum number of events stacked is not satisfied with having the

preset minimum search radius, the search radius is expanded to another stacking bin

spacing divided by √2.


30

Figure 3.4. Illustration of stacking source-receiver pairs by bouncepoint location.

Seismic records are stacked by the location of their bouncepoints relative to other

events within the defined stacking radius (0.5°,1°,2°,4°,8°, or 12°). The black triangles

represent the events that are within the stacking radius. The red triangles are outside of

the search radius and therefore will not be considered for stacking.

Crustal Tests

Due to the fact that our study did not implement a crustal model in our study,

we decided to test how the quality of our data would have changed and assessed

whether the addition of a crustal model would be necessary. A crustal model would

have accounted for changes in the amplitude and shape of the P410P pulse caused by

multiple crustal layers with different Vp values. However, since the P410P never

passes through the crust at its midpoint, the P410P phase itself is not altered by the

crust but in the simultaneous iterative deconvolution we use the PP phase, which does

pass through the crust, will be effected. Therefore, the deconvolved P410P phase

should be altered to some degree by the crust. To determine to what extent we will


31

have to compensate for the P410P phase response to crustal structure, we have made a

series of synthetics for different crustal structure and combinations of crustal structure.

We used PREM velocity model (Dziewonski and Anderson,1981) structure and change

the crust in each test. Each test was run with variations in the thicknesses of the crust

using simultaneous deconvolution of 8 different seismograms with different crustal

thicknesses. These tests were repeated at the following frequencies: 2, 1, 0.5, and 0.25

Hz. The first round of tests addressed the effects of differing crustal thicknesses as a

function of frequency on amplitudes of P410P, arrival times of P410P, and amplitude

ratio between P410P and PP (Figure 3.5).

0

0.5

1

1.5

0 2 4 6 8 10 12 14

2 Hz Crust Test

P410P arrival P410P amplitude P410P/PP

Test number

0

0.2

0.4

0.6

0.8

1

1.2

0 2 4 6 8 10 12 14

1 Hz Crust Test

P410P arrival P410P amplitude P410P/PPTest number


32

Figure 3.5. P410P amplitudes at 2, 1, 0.5, and 0.25 Hz filter frequencies for

simultaneous deconvolution of 8 synthetic seismograms with each having a crust of

random thickness (between 20 km and 60km). Amplitudes were normalized by the

expected amplitude of the P410P phase. The horizontal axis is the trial number for the

12 tests.

All numerical values are normalized to the mean for each category (i.e.: P410P time,

P410P amplitude, and P410P/PP ratio). The arrival time for the P410P appears to be

unaffected by the presence of a crustal layer at every frequency. This is particularly

important for our results since we can rule out crustal effects for variations in the

depth of the 410 km discontinuity in our results at any frequency. There is, however, a

0

0.2

0.4

0.6

0.8

1

1.2

0 2 4 6 8 10 12 14

0.5 Hz Crust Test

P410P arrival P410P amplitude P410P/PP

0

0.2

0.4

0.6

0.8

1

1.2

0 2 4 6 8 10 12 14

0.25 Hz Crust Test

P410P time P410P amplitude P410P/PP

Test number

Test number


33

clear effect on the amplitude of the P410P phase with variations in filter frequencies

and crustal thickness. However, we see that there were only errors in the expected

amplitude for 2 of the 12 trials. These results imply that, for random crustal

thicknesses between 40 and 60 km, we can expect the correct amplitude the majority

of the time. Pulses generated from this PP precursor had noticeably broader pulse

widths and were less robust at lower frequencies in comparison to the sharper PP

precursor spikes observed at 2Hz and 1Hz. A synthetic seismogram for eight

consecutive crustal models, each having a thickness of 30 km, is displayed in figure

3.6. Notice the variation in amplitude and pulse widths between synthetic

seismograms filtered at 2 Hz and 0.25 Hz. Sharper pulses are observed for seismic

records filtered at 2 Hz, while more diffuse and shorter pulses are observed in seismic

records filtered at 0.25 Hz.


34

Figure 3.6. Synthetic seismogram using PREM model. Frequency is 2Hz (a) and 0.25

Hz (b) with a sampling rate of 40sps. The reflection corresponding to the P410P

boundary arrives around 87.5 s.

The second round of testing was performed to determine the effect of various crustal

thicknesses on amplitude as a function of frequency. Results from this test are

displayed in figure 3.7 for four different crustal thicknesses of 20, 30, 40 and 50 km

with a 10 km variation for each thickness parameter. Fifty iterations for each average

crustal thickness were performed at filter frequencies of 2, 1, 0.5, 0.25 and 0.125 Hz.

For every iteration, each of the 8 synthetics were produced using 8 different source

functions that were extracted from the data. The precursor arrival time and PP arrival

time are consistent for every iteration regardless of the crustal thickness used or

frequency. So depth estimates should not be significantly affected by crustal thickness.

a)

b)


35

The mean amplitude for each P410P arrival is plotted as a single point with error bars.

The error bars are defined as the standard deviation of the 50 amplitude measurements

of the P410P phase from each crustal test. These crustal tests were conducted for

crustal thickness ranges of 20km ± 10km, 30km ± 10km, 40km ± 10km and 50km ±

10km. All of the precursor amplitudes were normalized by the maximum amplitude of

the main PP arrival.

Figure 3.7. Crustal test with 50 iterations at 2Hz (a), 1Hz (b), 0.5Hz (c), 0.25Hz (d)

and 0.125Hz (e). The blue trend-line connects the mean amplitude for 8 crustal layers

at various thickness ranges from 20km±10km, 30km±10km, 40km±10km, and

50km±10km. The black asterisks represent individual P410P amplitudes.

The P410P had errors in the amplitude at all frequencies between 2Hz and

0.5Hz. The amplitudes for P410P vary considerably for filter frequencies between 2Hz

a) b)

c) d)

e)


36

and 0.5Hz, with a maximum mean amplitude variation of 0.0064 at 2Hz, 0.0069 at

1Hz, and 0.0068 at 0.5Hz. The best results were observed using filter frequencies of

0.25Hz and 0.125Hz where the amplitudes were the most consistent with smaller

standard deviation values. Although the mean amplitude variation for 0.25 Hz is

comparable to the other mean amplitude corresponding to higher filter frequencies, the

consistency of the individual amplitudes is more pronounce, especially for crustal

thicknesses between 30 km and 50 km, lead us to believe even at these frequencies

P410P amplitudes have value if the crust at the midpoint is normal continental crust

(30 to 40 km). These results could indicate that imaging the 410 could be problematic

at 0.125Hz for events with bouncepoints located in thinner crustal regions, like

oceanic regions or rift zones.


37

CHAPTER IV

RESULTS

Analysis on the P410P phase as a function of frequency will help determine the

depth interval of the velocity gradient associated with the 410 km discontinuity, so we

picked the depths and amplitudes for the P410P phase at frequencies of 1, 0.5, and

0.25 Hz (see figures 4.1a through 4.1f). PP-functions were also produced at a low pass

filter frequency of 0.125 Hz. While the resolution of the 0.125 Hz PP-functions was

noticeably cleaner than those produced at 0.5 Hz and 0.25 Hz, the P410P was not

consistently observed and the data density was too sparse for global interpretation

when sampled at a frequency of 0.125 Hz, so these results were not included. Where

data density was high and the quality of the P410P phase was satisfactory, the depths

to the P410P boundary were mapped. There are larger variations in the amplitudes of

the P410P phase, resulting in the appearance of abrupt spikes in pulse amplitudes.

These spikes in amplitude could be attributed to changes in crustal thickness and the

detection of localized variations in the topography of the upper mantle at 410 km

depth, and since the frequency of the data used for our study was higher than those

used in many older studies (which are typically filtered at 0.01Hz), we expect better

resolution.

Maps of P and S wave velocity perturbation at 400 km are displayed in figure

4.2, where Vp and Vs are expressed as the percentages of the expected velocity of the

mantle at 400 km depth, which is approximately between 8.91 to 9.13 km/s and 4.77

to 4.93 km/s, respectively. These maps will be used to determine if there is any

correlation between velocity, amplitdude and depth of the 410 km discontinuity which


38

presumably would be related to variations in temperature. Negative percentage values

are indicative of a slower mantle region and positive percentage values are indicative

of a faster mantle region. Overall, Vp appears to be faster beneath continents and

slower beneath areas that are within close proximity to active rift zones, hot mantle

plumes, and subduction zones. However, since shear seismic waves are more sensitive

to temperature variations than P-waves, we will rely more heavily on the Vs velocity

perturbation maps to better understand the significance of the observed depths and the

amplitude variations inferred from P410P phases. The fastest S wave velocities are

observed beneath parts of western Europe, central Africa, southern Japan and

Australia. The slowest S wave velocities are observed beneath areas where high

thermal gradients are found like the East African rift, Iceland, Hawaii and the South

Pacific.

Figure 4.1a. Map of depths to the 410 km discontinuity at 1Hz. The depth range is

between 370 km and 450 km below the Earth’s surface.

Dep

th (k

m)


39

Figure 4.1b. Map of the amplitudes of the P410P phase at 1Hz. The amplitudes

displayed are measured relative to the main pulse of the direct PP arrival by

normalization at a range between 0 and 0.06.

Figure 4.1c. Map of depths to the 410 km discontinuity at 0.5Hz. The depth range is


Dep

th (k

m)

P410P

/PP


40

Figure 4.1d. Map of the amplitudes of the P410P phase at 0.5Hz. The amplitudes



Figure 4.1e. Map of depths to the 410 km discontinuity at 0.25Hz. The depth range is


Dep

th (k

m)

P410P

/PP


41

Figure 4.1f. Map of the amplitudes of the P410P phase at 0.25Hz. The amplitudes



Figure 4.2a. Global distribution of S wave velocity (Vs) perturbations represented as

the percentage of the expected Vs at 400 km depth (~ 4.77 to 4.93 km/s). Unlike Vp,

Vs is more sensitive to temperature variations.

% o

f expected

Vs at 4

00 k

m

P410P

/PP


42

Figure 4.2b. Global distribution of P wave velocity (Vp) perturbations represented as

the percentage of the expected Vp at 400 km depth (~ 8.91 to 9.13 km/s).

By mapping stacked PP-precursor functions at filter frequencies of 1 Hz, 0.5

Hz, and 0.25 Hz, we are able to observe many trends regarding the nature of the 410

km discontinuity worldwide. The deepest depressions of the 410 km discontinuity

tends to be concentrated along the major oceanic-continental subduction zones around

the world. These depressions are most apparent along the western coast of North

America and Eurasia. Deeper 410 km boundaries are also observed beneath inland

regions of Eurasia around Tibet where the Indian plate is currently colliding with the

Eurasian plate. Other notable areas where there is a deepening of the 410 are observed

beneath central South America, Greenland and Iceland where the 410 appears deepest.

We also analyzed the geographic distribution of variations in the amplitudes of the

P410P phase in relation to depth variations of the 410 km discontinuity and the global

velocity perturbation model (Vp and Vs). The lowest amplitudes are observed beneath

% o

f expected

Vp at 4

00 k

m


43

regions of deeper 410 km discontinuity depths, which are all indicative of the presence

of a warmer mantle region. Cooler mantle regions correspond to shallower 410 km

discontinuity depths and are located primarily beneath the Pacific Ocean and parts of

the Atlantic Ocean. The higher relative amplitudes observed beneath these regions

also support the theoretical presence of a cooler mantle at 410 km depth.

PP-precursor functions beneath Hawaii

A moderately depressed 410 is observed beneath Hawaii, appearing between

410 km and 420 km, with the deepest 410 observed beneath the northern most parts of

the island chain beneath the islands of Kauai and Ni’ihau (see figure 4.3). The mapped

P410P depth using stacks filtered at 0.25Hz was excluded here due to the lack of

coverage.


44

Figure 4.3. Depth of the 410 km discontinuity beneath Hawaii at 1Hz (a) and 0.5Hz

(b). The 410 km discontinuity is slightly elevated to a depth of approximately 400km

beneath the island of Hawaii. The 410 then appears to deepen toward the northeast.

The deepest 410 appears around 420 km depth beneath the island of Kaua’i.

Results from our depth analysis agree with previous studies conducted near Hawaii,

one in particular used only half of the currently available transportable array (TA) data

(see Ainiwaer, 2014). Ainiwaer [2014] mapped the 410 beneath Hawaii and found the

boundary to be slightly depressed beneath the northwestern part of the island chain at

approximately 425 km depth for 0.25Hz and 0.5Hz, although the depression is less

pronounced on the 0.5Hz map. Due to the exothermic nature of the 410, this

discontinuity is expected to be deeper beneath warm regions which would suggest a

deeper 410 near Hawaii is related to the hotspot in this region. Using the S wave

velocity perturbation model (Simmons et al., 2011), the mantle at 400 km depth is

a)

b)

Dep

th (k

m)

Dep

th (k

m)


45

warmest toward the oldest islands in the northwest where Vs is 1% and 1.5% slower

than the reference Vs.

Figure 4.4. Velocity perturbation expressed as the percent of the referenced S wave

velocity (Vs) around 400km (~4.77 to 4.93 km/s). Hawaii is located toward the center

of this figure, where the mantle velocity in proximity to the 410 is between 1% and

1.5% slower than the reference Vs.

The presence of a warmer mantle is also supported by observations made

through the amplitudes of the stacked P410P pulses. The amplitude of the P410P

precursor arrivals are noticably small with respect to the amplitude of the main PP

arrival, being approximately 2% to 3% of the amplitude of the main PP pulse. The

expected amplitude of the P410P pulse, which is ~0.035, is derived from the velocity

and density contrasts in PREM (Dziewonski and Anderson, 1981). The depth interval

of the velocity gradient at the 410 km discontinuity has also been shown to have an

effect on the amplitude and pulse width of the P410P phase (see Ainiwaer, 2014).

Ainiwaer (2014) found that P410P pulse amplitudes were lower and the pulse was

wider for a simulated 30 km thick velocity gradient at 410 km depth as compared to

the P410P phase for the expected 10 km gradient. Due to the presence of a warmer

mantle beneath Hawaii, the chemical phase chnge from olivine to wadsleyite is

expected to be relatively thin (see Akaogi et al., 1989). Figure 4.5 displays the

%V

s at 40

0 k

m


46

amplitdue patterns observed beneath Hawaii using stacked PP functions filtered at 1Hz

and 0.5Hz, respectively. The average amplitdue beneath Hawaii ranges between 0.025

and 0.03 for 1Hz and 0.5Hz, which are lower than the 0.035 value predicted by the

model of the velocity gradient at 410 km depth using PREM (Dziewonski and

Anderson, 1981).

Figure 4.5. Map of P410P amplitudes beneath Hawaii and the immediate surrounding

area at 1Hz(a) and 0.5Hz(b). The average amplitdue beneath Hawaii ranges between

0.025 and 0.03(2.5% to 3% of the main PP pulse) for 1Hz and 0.5Hz.

PP-precursor Functions beneath Alaska

The average depth to the 410 beneath the majority of Alaska is 430 km which

would be indicative of a mantle that is hotter than usual. Even though there are no

known hotspots near Alaska, there is a subduction zone along the western coast where

the Pacific plate is subducting beneath Alaska from the northwest at a rate of 54

mm/yr (Yinshuang et al., 2005). The 410 is particularly deep beneath Alaska, but the

b)

a)

P4

10

P/P

P

P410P

/PP


47

410 is shallower (405 km to 410km) to the immediate southeast of the subduction

zone (see figure 4.6). There is also a slight deepening of the 410 to the northwest

between Russia and Alaska. While the depth results appear consistent between the 1

Hz and 0.5 Hz maps, the depth of the 410 km discontinuity appears to vary the most

with stacked PP functions filtered at 0.25 Hz. The velocity perturbation model of the

mantle at 400 km beneath Alaska and the surrounding areas show a fast velocity ano-

maly stretching from the Aleutian island chain in the southwest to the north/northeast,

ending a little more than half way beneath the Alaskan mainland where the Vs

velocity is approximately 1% faster than expected (see figure 4.7).

Figure 4.6. Depth of the 410 km discontinuity beneath Alaska at 1Hz(a), 0.5Hz(b),

and 0.25Hz(c). 410 depth ranges from 430 km – 440 km beneath Alaska at 1Hz and

0.5 Hz. The 410 appears shallower at 0.25 Hz.

a) b)

c)

Dep

th (k

m)

Dep

th (k

m)

Dep

th (k

m)


48

Figure 4.7. Vs velocity perturbation beneath Alaska at 400 km depth. There is a fast

Vs anomaly observed beneath the Aleutian islands up toward the Alaskan mainland to

the north/northeast. Vs in this region is around 1% faster than expected.

P410P amplitudes appear to be relatively low (around 0.02) beneath the majority of

Alaska at 1 Hz, with higher amplitudes observed to the southeast of Alaska (around

0.04) and as high as 0.06 to the south (see figure 4.8). The amplitudes vary

considerably for the stacked PP functions filtered at 0.5 Hz, as the average P410P

pulse amplitude is around 0.02 to 0.03 with an anomaly in amplitudes of 0.04 to 0.06

south of Alaska. The results from stacked PP functions filtered at 0.25 Hz has a

considerable variation in amplitude as well. P410P pulse amplitudes are observed to

be around 0.02 beneath most of Alaska, but can be as high as 0.04. Higher amplitudes

are also observed to the south/southwest.

%V

s at 40

0 k

m


49

Figure 4.8. Map of P410P amplitudes beneath Alaska at 1Hz(a), 0.5Hz(b), and

0.25Hz(c). For all images, the P410P amplitude observed beneath Alaska ranges from

0.02 to 0.03. The highest amplitudes are observed to the south and southwest of the

Alaskan mainland.

PP-precursor functions beneath Eurasia

The Eurasian plate, particularly beneath eastern most Russia and China, have

significant temperature anomalies in the upper mantle because it is also part of a

subduction zone. This area has two subduction events. To the north, the Pacific plate is

subducting beneath the North American and Eurasian plates. To the south, the

Philippine plate is subducting beneath the Eurasian plate. The depth of the 410 km

discontinuity beneath this region indicate the presence of a warmer mantle beneath the

Eurasian continent and the Sea of Okhotsk, with a progressively cooler mantle to the

east beneath the Pacific Ocean. The depth of the 410 appears to be shallower beneath

a) b)

c)

P4

10

P/P

P

P4

10

P/P

P

P410P

/PP


50

Russia with an average depth of ~420 km. The depth observed beneath China is ~430

km for PP functions filtered at 1 Hz (see figure 4.9). The depth of the 410 was more

consistent beneath China and Russia with an average depth of 420 km when PP

functions were filtered at 0.5 Hz. The 410 appears to deepen beneath southwest Russia

and Mongolia to approximately 430 km. The 410 appears deepest beneath Lake Baikal

and the region approximately 15° to the east when PP functions filtered at 0.25 Hz are

used. Although coverage is weakest for PP functions filtered at 0.25Hz beneath

eastern Russia, the 410 is clearly observed to be approximately 420 km deep. The 410

appears to gradually shallow beneath Russia around 135°E and continues to shallow to

depths between 400 km and 410 km at ~105°E to the west. The 410 also appears

elevated beneath southern Kamchatka when PP functions filtered at 1 Hz and 0.5 Hz

were used, but for 0.25 Hz PP functions, the 410 is shallow beneath the entire

peninsula of Kamchatka.


51

Figure 4.9. Depth of the 410 km discontinuity beneath Eurasia at 1Hz(a), 0.5Hz(b),

and 0.25Hz(c). At 1 Hz, the 410 appears depressed beneath eastern Russia at around

420 km depth. The 410 is more uniformly depressed throughout China and eastern

Russia at 420 km depth with the use of PP functions filtered at 0.5 Hz. There are

considerable depressions observed beneath Lake Bikal and areas to the east at 1 Hz

and 0.5 Hz.

Velocity perturbations at a depth of 400 km beneath Eurasia are generally fast. The

fastest P and S wave velocity is observed beneath southern Japan and the Sea of Japan

where it is estimated to be around 2% greater than expected. S wave velocities

a)

b)

c)

Dep

th (k

m)

Dep

th (k

m)

Dep

th (k

m)


52

observed beneath nearly all of the northeastern region of China show velocity

anomalies that are about 1% to 1.5% greater than average. Vs appears to be close to

average beneath Lake Baikal (see figure 4.10).

Figure 4.10. Vs velocity perturbation beneath Eurasia at 400 km depth. The fastest P

wave velocities are observed beneath southern Japan and the Sea of Japan at

approximately 2% faster than average at depth.

The amplitude of the P410P beneath Eurasia is relatively small for the majority

of the area with amplitudes between 0.01 and 0.03. The highest amplitudes are

observed beneath South Korea and Russia, especially around Lake Baikal. The highest

P410P amplitude is observed on the 0.25 Hz plot, where the amplitude is as high as

0.06 in South Korea and just north of Lake Baikal. The magnitude averages around

0.04 on the 1 Hz and 0.5 Hz plot for the majority of the region beneath Lake Baikal,

with particularly high amplitudes to the north. The amplitude of the P410P reflection

(see figure 4.11) is also greater beneath most of the Kamchatka peninsula, which is in

agreement with the trend observed with the depth of the 410 in figure 4.9.

%V

s at 40

0 k

m


53

Figure 4.11. Map of P410P amplitudes beneath Eurasia at 1 Hz(a), 0.5 Hz(b), and

0.25 Hz(c). The amplitude of the P410P pulse beneath Eurasia is relatively small for

the majority of the area with normalized values between 0.01 and 0.03. Some of the

greatest amplitudes are observed beneath the region north of Lake Baikal.

PP-precursor functions beneath northwestern Europe, Greenland, and Iceland

The 410 km discontinuity is deep beneath the north Atlantic from Greenland

through western Europe. The 410 appears to be deepest (~450 km) beneath Iceland,

which is to be expected given the presence of the hot spot beneath Iceland. There is a

a)

b)

c)

P4

10

P/P

P

P410P

/PP

P

410P

/PP


54

small area of gradual shallowing of the 410 observed beneath the Norwegian Sea and

appears to be right at 410 km below most of Sweden and Norway, but is slightly

deeper (~420 km) beneath the southeastern region of Sweden. The depth of the 410

km discontinuity beneath the United Kingdom and Ireland is between 410 km and 420

km.

In the 1 Hz map, the 410 shallows beneath most of Greenland, to ~420 km but

is deepest beneath the eastern and northern coasts of Greenland at depth of 430 km.

The 410 also deepens beneath the eastern coast of Greenland. A slightly elevated 410

is observed beneath the northeastern coast of Greenland (~400km) at 0.5 Hz. Depths

to the 410 for the 0.25 Hz results generally agree with the 0.5 Hz and 1 Hz results

except beneath Iceland. Instead, the 410 appears to be right at 410 km depth in the

0.25 Hz image beneath Iceland. The rest of the region was not data rich enough at

0.25Hz. Due to this discrepancy, results from 1 Hz and 0.5 Hz will be displayed for

the depth plots since these plots had higher data density in this area, and is in

agreement with the known thermal anomalies inferred beneath Iceland from velocity

anomalies. The temperature of the mantle at 410 km depth appears to decrease toward

the south at approximately 60°N latitude.


55

Figure 4.12. Depth of the 410 km discontinuity beneath NW Europe, Greenland and

Iceland at 1Hz(a) and 0.5 Hz(b). The 410 appears to be right at 410 km below most of

Sweden and Norway, with a slight elevation to around 420 km depth toward the

southeastern region of Sweden. The 410 appears at approximately 450 km beneath

Iceland.

S wave velocities are shown in figure 4.13 and as expected, the mantle

temperature beneath Iceland is relatively warm, where Vs is ~2% slower than

expected at 400 km. The S wave velocity gradually increases to the south until Vs is

~1% faster than expected beneath the north Atlantic Ocean. S wave velocity anomalies

compliment the pattern observed in the depth plot beneath Greenland as well, as Vs is

slightly faster than the expected velocity along the eastern coast, and becomes

progressively faster inland where Vs is at most 1% faster than expected. There appears

to be no real velocity anomaly beneath Finland, but the Vs beneath Sweden appears to

a)

b) D

epth

(km

) D

epth

(km

)


56

be approximately 1% faster than expected. There appears to be no real velocity

anomaly beneath Finland, but the Vs beneath Sweden appears to be approximately 1%

faster than expected.

Figure 4.13. Velocity perturbation trends beneath NW Europe, Iceland and

Greenland. Slower S wave velocities are observed primarily around Iceland where Vs

is ~1.5 to 2% slower than expected at 400 km depth. The faster velocities are present

beneath most of Greenland and NW Europe where Vs is 1% faster than expected at its

fastest.

Amplitude trends for P410P reflections correlate with respect to mantle

temperatures inferred from regional depths of the 410 km discontinuity. The relative

amplitude of the P410P beneath Iceland was around 0.03 at 1 Hz and 0.5 Hz. The

lowest observed amplitudes were found beneath Greenland and parts of Finland. The

amplitudes beneath Greenland and Finland are lower than observed beneath Iceland

where the mantle is much warmer, the P410P amplitudes correlate with the

tomography of the upper mantle at the 410 km discontinuity. The average amplitude

beneath the coast of Finland is around 0.01 to 0.02. Greenland has the same

approximate range as Finland, although, the lower amplitude values at 0.5 Hz extend

%V

s at 40

0 k

m


57

farther inland into Sweden. The average amplitude observed beneath Sweden is

approximately 0.035 at 1 Hz.

Figure 4.14. Map of P410P amplitudes beneath NW Europe, Greenland and Iceland at

1Hz(a) and 0.5 Hz(b). The P410P pulse beneath Iceland has a magnitude of

approximately 0.03 for both 1 Hz and 0.5 Hz. The P410P pulse beneath Greenland

ranges between 0.01 and 0.02 for both frequencies as well. The amplitude range

observed beneath Finland and Sweden is 0.01 – 0.02 at 0.5Hz, but the amplitude

appears to increase to approximately 0.035 beneath Sweden and parts of Finland at 1

Hz.

PP-functions beneath South America

The 410 km discontinuity appears to be deep beneath the center of South

America, 420 km and 430 km, from Peru to Brazil at 1 Hz and 0.5 Hz. Depths picked

from PP functions filtered at 0.25 Hz are shallower, between 380 km and 400 km for

P4

10

P/P

P

P410P

/PP

a)

b)


58

the same area. However, the data density of this region at 0.25 Hz is also sparse, so

these results are not likely to be reliable.

Figure 4.15. Depth of the 410 km discontinuity beneath South America at 1 Hz (a)

and 0.5 Hz (b). The 410 appears to be depressed throughout the majority of this region

with depth ranges between 420 km and 430 km depth. The cool mantle region to the

north appears to be a data processing error due to the abrupt nature of the apparent

elevation of the 410 km discontinuity. The 410 does appear to gradually elevate to the

east, beneath the south Pacific.

S wave velocity anomalies beneath South America do not correlate as well with the

mapped depth results for the 410 for the most part. The Vs beneath this region are as

much as 1% faster than expected. Small areas on the map show no apparent velocity

anomalies (see figure 4.16).

Dep

th (k

m)

Dep

th (k

m)

a

)

b

)


59

Figure 4.16. Velocity perturbation trends beneath South America. Slower S wave

velocities are observed beneath the majority of central South America where the

slower anomaly is ~0.5 to1% faster than expected at 400 km depth.

The amplitude of the P410P reflection beneath this region is relatively small in

the region where velocity anomalies infer a warmer mantle (see figure 4.17). The

amplitude from the 1 Hz PP functions are between 0.01 and 0.025 beneath the

majority of the region. The average amplitude is 0.04 beneath the western coast of

South America. The largest P410P reflection, with an amplitude of 0.05, is observed in

the southwestern portion of this region. The mapped amplitude using 0.5 Hz PP

functions shows the same general pattern observed in the 1 Hz map with a few

exceptions. The amplitude of the P410P reflection observed beneath the western coast

at 1 Hz was not only greater in amplitude, but extended further inland beneath what

would be the western region of Brazil, where the range of amplitude is between 0.02

and 0.03. The peak P410P amplitudes are still located around the southwestern portion

of the mapped region at 0.5 Hz, but the pulse amplitudes can be as high as 0.06. P410P

pulse amplitudes remain small beneath the south Pacific.

%V

s at 40

0 k

m


60

Figure 4.17. Map of P410P amplitudes beneath South America at 1Hz (a) and 0.5Hz

(b). The average magnitude correlating to the P410P pulse range between 0.01 and

0.025 throughout most of the mapped region. A concentration of higher magnitudes

are located along the western coast of this area with a particularly strong P410P pulse

observed in the southwestern region.

P410P

/PP

P

410P

/PP

a

)

b

)


61

CHAPTER V

DISCUSSION

Analysis of global variations in the 410 km discontinuity using PP precursors

The PP functions that were computed from seismic events mapped in figure

3.1 provided a means for studying the nature of the 410 km discontinuity on a global

scale. We focus here on patterns in depth and amplitude variations estimated from

analysis of P410P phases. We see a significant depression of the 410 (as much as 450

km) beneath all of Iceland, where a known hotspot exists (see Vink, 1984). There are

also some zones that seem contradictory to the hypothesized exothermic behavior of

the 410 km discontinuity, most notably beneath much of northeastern China. The 410

km discontinuity appears to be deeper (~430 km) beneath northeastern China, despite

the fact that the mantle region is thought to be cooler due to the subduction of the

Phillipine plate beneath the Eurasian plate (Li and Yuan, 2003). A possible cause of

this region being warm despite the presence of a subducting slab is the accumulation

of mélange from the Philippine plate at the 660 (Suetsugu et al., 2010), which could

drive a local convection cell that would warm the 410 (see Zhao et al. [2012] and

Duan et al. [2009]). Another possibility for the warming of the 410 is the presence of

partial melt due to mantle hydration from the subducted Philippine plate.

PP precursor amplitudes in many areas appear to correlate with depth

anomalies but there are some notable exceptions. The 410 is a boundary where density

and seismic velocity of the mantle increases quickly over a small enough interval to

cause seismic reflections. This velocity change, however, occurs as a gradient over a


62

small interval whose depth ranges can change depending on the temperature of the

mantle. For example, the P410P phase amplitudes are expected to be larger and

narrowed due to a narrow velocity gradient in the mantle in warm regions (see

Ainiwaer, 2014), which is observed to some degree beneath regions like Iceland, NE

China and Hawaii. Otherwise, patterns in the amplitude of P410P reflections vary

somewhat haphazardly throughout the globe with some patterns of high P410P

amplitudes being observed roughly near the mid-Atlantic Oceanic Ridge,

west/southwest of Hawaii, and the southwestern Pacific. These observations roughly

agree with P410P and S410S amplitude trends in the study conducted by Chambers et

al. [2005], where higher S410S amplitudes were observed beneath the south Pacific,

north/northwest of Hawaii and along the mid-Atlantic Oceanic Ridge. Although

Chambers et al. [2005] amplitude results from P410P reflections vary noticeably,

higher P410P amplitudes are observed along the Nazca Ridge, along the southeastern

African coast, east of Hawaii and west of Hawaii for our study. Low P410P reflections

appear beneath the majority of most continents like Russia, northeastern China, and

the middle of South America. Our low P410P reflection trends are in agreement, for

the most part, with the P410P and S410S analysis by Chambers et al. [2005] with the

exception of South America. The following sections will focus on specific regions

were the degree of data coverage is relatively high and are in proximity to major

tectonic events and geological features.

Hawaii

We found that the depth to the 410 beneath the area surrounding Hawaii is

greater than 415 km which is generally deeper than expected (see figure 4.3). The


63

deepest observed 410 is in the northwest direction along the island chain (~420 km).

This is consistent with a P410P bouncepoint study by Duncan (2012) using only half

the data that is currently available from the transportable array (TA). The work

presented here includes all the TA stations and any other global seismic stations with

data available through the IRIS DMC. Nolet et al. [2007] observed a low velocity

anomoly to the north of the Hawaiian island chain, tilting to the south from the surface

down to the core-mantle boundary (CMB). The P-wave tomographic map is shown in

figure 5.1, where Vp beneath Hawaii is 0.05% slower than average. But the S-wave

velocities imaged beneath the region at 400 km depth in the GYPSUM model

(Simmons et al, 2011) are ~0.5 to 1% lower then those of PREM which would be

consitent with only a small regional increase in temperature and the small degree of

deepening in the 410 that we observe.

Figure 5.1. P-wave tomographic map. The proposed hotspot is seen beneath Hawaii

where the Vp is 0.5% slower than average. (Nolet et al., 2007)

Figure 5.2 is a depth profile that starts in the Pacific Ocean, southwest of the

main Hawaiian island, and crosses beneath the island chain ending to the northwest of

Kaua’i. The P410P phase appears at ~400 km depth through a series of larger sharp

spikes with narrow pulse widths. As the profile moves across the islands a slight


64

depression of the 410 appears to deepen to ~420km beneath Maui and older islands to

the northwest. Approximately halfway between O’ahu and Kaua’i, the P410P pulse

narrows.

Figure 5.2. Depth profile of Hawaiian islands at 0.5 Hz filter frequency. Starting from

left (southeast) to right (northwest), the 410 appears at approximately 400 km beneath

Hawaii and begins to depress toward the northwest around Maui. The 410 is

highlighted by the bold dashed line.

If the current position of the mantle plume is expected to be beneath the island

of Hawaii, we would expect the 410 to be at its deepest beneath this region, but

instead, we observe a slight depression of 410 beneath the older islands to the

west/northwest. A study conducted by Cao et al. [2011] also found evidence for a

thicker than average TZ beneath Hawaii and thinner to the east and west of the island

chain, which support the hypothesis that the hot spot passes through the TZ around

Hawaii, but not directly beneath the island. Another study utilized a 3 dimensional

simulation of the interaction between the mantle plume beneath Hawaii and the local

lithosphere where Moore et al. [1998] found that the cause of variations in heat flow

beneath these islands may be a result of convective cells or small convective rolls,

which are thought to be aligned with the motion of the Pacific plate. According to

Depth Profile beneath Hawaiian Islands

Dep

th (

km

)

410


65

Moore et al. [1998] the convective cells can cause localized thinning of the

lithosphere, which effectively leads to temperature variations within the

asthenosphere, correlating with the observed trend of a slightly deeper 410 km

discontinuity beneath Hawaii. Results for our study, however, more consistently aligns

with the findings of Cao et al. [2005], where the hotspot is not directly beneath

Hawaii, but is instead located closer to the older islands to the northwest.

Alaska and the Aleutian Islands

The Aleutian islands were formed as a result of the Pacifc plate subducting

beneath the North American plate. Subduction zones tend to cause the surrounding

mantle to become cooler due to the presence of a relatively cooler subducting slab.

Water trapped within hydrated minerals and pore spaces of a subducting oceanic plate

can cause partial melting in the surrounding mantle (Hyndman and Peacock, 2003)

which leads to the formation of a volcanic arc. Although the Aleutian Island chain is

known to be the product of forearc volcanism, the warmest mantle region is observed

beneath Alaska, where the highest depression of the 410 km discontinuity is located

beneath the eastern and northern areas of this region. Alternatively, the coolest mantle

region is observed beneath the Aleutian islands where the 410 is elevated to a depth of

approximately 400 to 405 km (see figure 4.6).

The depression of the 410 beneath Alaska could be correlated to local thinning

of the transition zone (see figure 5.3). Yinshuang et al. [2005] mapped variations in the

thickness of the transition zone beneath the south-central region of Alaska and

generally found thinning of the TZ, by as much as 30 km, beneath this region. Thinner

transition zones correlate to the presence of warmer manle regions due to the


66

exothermic and endothermic nature of the 410 and 660, respectively, and could

indicate that there is a warm mantle beneath Alaska. In comparison to our results

beneath south-central Alaska, the 410 is deepened to a depth range between 420 km

and 430 km for 1 Hz and 0.5 Hz data. The slightly thinner TZ observed in the upper

right corner of figure 5.3 is also observed to some degree in our results as well. The

410 shallows to 410 km in the same area but occurs in a smaller region.

Figure 5.3. Transition zone thickness anomaly beneath south-central Alaska. TZ

thickness variations are relative to the average thickness of 250 km. (Yinshuang et al.,

2005)


67

Figure 5.4. Depth profile beneath southern Alaska at 0.5 Hz filter frequency. Starting

from the east(left), moving to the west(right), the 410 is clearly depressed beneath the

entire landform. There also appears to be a consisnent negative velocity layer directly

above the 410, indicative of partial melt. The 410 is highlighted by the bold dashed

line.

Figure 5.5. Depth profile beneath the western end of the Aleutian Island chain at 0.5

Hz filter frequency. Starting from the island to the south(left) and progressively

moving to the north toward Russia(right), a possible artifact from the subducting slab

is observed by the negative P wave discontinuity pulse.

The presence of a thinner transition zone could be caused by warming due to partial

melt as a result of slab dehydration or localized convection from the subducting

Pacific plate. Low P410P amplitudes beneath Alaska also correlates to the presence of

Dep

th (

km

) D

epth

(km

)

Depth Profile beneath Aleutian Islands

410

Possible partial melt artifact from subducting slab

410

Depth Profile beneath Alaska


68

a slow velocity layer found on the subducting slab beneath Alaska in a study

conducted by Yinshuang et al. [2005]. The persistent negative velocity anomoly

observed directly on top of the positive P410P pulse in figure 5.5 is indicative of the

presence of partial melt which supports the hypothesis that the mantle is warm due to

local hydration of the mantle from the subducting slab.

Eurasia and N. India

The next region of interest was mapped in a data rich area of the eastern

Eurasia, from northern Russia to eastern China. The 410 appears to be depressed

beneath northeastern Russia, where the Pacific plate subducts beneath NE Russia and

China, and inland towards Mongolia and the Lake Baikal rift zone (see figure 4.9).

The 410 km discontinuity shallows back up to 410 km beneath the southern tip of

Kamchatka. Since there is nothing more than a transform fault in the Kamchatka

pennisula, there is no current tectonic activity that would not contribute to a change in

the temperature the upper mantle at all (see Chapman and Solomon [1976]).

An increase in the depth of the 410 to ~440 km beneath the Lake Baikal rift

zone (see Gao et al., 2003) suggests that this rift could be driven by magma rising

from a mantle plume, making this an active rift zone. Gao et al. [2003] found the

Baikal rift, created by the rifting of the Siberian and Amurian microplates, to be one of

the most active rift zones in the world, where the thickness of the continental crust is

estimated to be reduced by no more than 5 km. However, other researchers believe

that the Baikal Rift is a passive rift caused by the crustal extension related to the

collision between Eurasia and India (see Petit et al. [1997]). Our results, however,

support the hypothesis that there is a warmer mantle beneath Lake Baikal (see figure


69

4.9). To cause a significant warming of the upper mantle, additional geologic

processes must be occurring as vertical mantle upwelling was found beneath Lake

Baikal in a study by Gao et al. [2003]. Evidence supporting a deep source for the

Baikal rift include: observations from bouguer anomaly modeling that the lithosphere

above the rift zone has been uplifted due to vertical mantle convection (Gao et al.

[2003] and Zorin et al. [2003]); and observations of seismic anisotropy which inferred

that the observed vertical alignment of the olivine crystals along the mineral’s fast axis

is the result of mantle upwelling (Gao et al. [2003]). North of Lake Baikal, the 410 is

elevated to the expected nominal depth of 410 km between approximately 85°E to

135°E. This gradual cooling of the mantle can be attributed to the fact that this region

is composed of relatively old lithosphere with no significant temperature anomalies.

A vast area beneath Mongolia and Tibet also appears to have a warmer than

usual mantle, as the 410 appears as deep as 440 km beneath some areas. Seismic and

orogenic activity in the Indian subcontinent and Tibet, are the products of the Indian

plate colliding into the Eurasian plate. Kind et al. [2002] mapped the discontinuity of

the 410 and 660 beneath the Himalayan plateau using data from two profiles across

Tibet. Kind et al. [2002] found evidence supporting the presence of the Tethyan plate,

an older oceanic slab, which would have detached from the India plate. This slab

subduction may cause the surrounding mantle to become progressively warmer over

millions of years due to partial melt caused by dehydration of the slab. This is a

reasonable hypothesis since Kind et al. [2002] mapped a subducting slab beneath Tibet

in their results (see figure 5.6), where they refer to this slab as the detached Asian

Lithospheric Mantle (ALM).


70

Figure 5.6. The detached Asian Lithospheric Mantle subducting beneath the

lithosphere of Tibet. (Kind et al., 2002)

Another study conducted by Li et al. [2006] imaged a subducting slab as well, but

theorized that part of the Indian slab is subducting along with the older Tethyan plate.

Figure 5.7 shows our depth profile beneath northern India and southern Tibet from

P410P observations. Although the detached ALM is not seen in this profile, the

negative amplitude anomaly directly above and below the 410 could be the result of a

reflection off of partial melt due to the migration of water to the 410 from the Tethyan

plate.


71

Figure 5.7. Depth profile beneath India and Tibet at 0.5 Hz filter frequency. Starting

from the south(left), moving roughly to the northeast(right) across what would be part

of the collisional arc stretching into southern Tibet. The P410P is not only deep, but it

appears to be flanked by low velocity layers above and below. This could be indicative

of the presence of continent crust below and partial melt above.

The next area of interest is northeast China where the Philippine plate is

subducting beneath the Eurasian plate to the south. The P410P pulse widths appear to

narrow toward the northwest. The presence of water has been hypothesized to narrow

the width of the 410 km discontinuity (see Chen et al., 2002). Frost et al. [2007],

however, hypothesized that the presence of water will broaden the mineral phase

change interval at 410 km since water will preferentially partition into wadsleyite. The

abundance of iron could be a possible cause for a narrowed transition boundary at the

410 km as well, leading to a sharper P410P reflection. Although the bulk modulus of

β-olivine is thought to increase with an increase in iron content, an increase in the bulk

modulus would result in only a slight increase in the P-wave velocity gradient between

α-olivine and wadsleyite (Sinogeikin et al., 1998). Iron content in olivine is also

thought to decrease under higher pressures due to the formation of relatively iron rich

Depth Profile beneath northern India and southern Tibet

Dep

th (

km

)


72

majorite garnet (Irifune and Isshiki, 1998), effectively reducing the pressure interval

between α-olivine and β-olivine (wadsleyite), and narrowing the phase change interval

between α-olivine and wadsleyite.

Whether the subducting slab is penetrating the base of the 660 or stagnant,

both processes can result in the cooling of the surrounding mantle, which would result

in the elevation of the 410. An argument could also be made for the possibility of a

warmer mantle region closer to 410 km since the dehydration of the Philippine plate

would cause partial melting. Depth anomalies associated with the α – β olivine

transition zone could also be attributed to remnant chemical heterogeneities from the

subducting slab (Li and Yuan, 2003) or mantle plume generation through re-heated

peridotite layers that have accumulated from stagnant slabs (Ringwood, 1991).

Although the Vp and Vs anomolies observed beneath northeastern China at 400 km

depth was faster by approximately 1% of the expected P wave and S wave velocity,

respectively, around 400 km according to Chambers et al. [2005], our results show

that there is a warmer mantle at the 410 km boundary as the 410 appears around 430

km at the deepest along the northeastern coast of China. This apparent contradiction

will require further investigation.


73

Figure 5.8. Depth profile beneath northeast China at 0.5 Hz filter frequency. Starting

from the south(left), moving to the north/northwest(right). The 410 is depressed to a

depth of approximately 430 km as we move across northeastern China. The P410P

pulse widths begin to narrow as the profile moves more inland to the northwest.

Greenland and Iceland

Iceland is split between the North American and Eurasian continental plates,

laying directly on top of the mid-Atlantic Ridge where these two plates are drifting

apart from each other. The formation of Iceland is theorized to be the result of the

presence of a magmatic plume in this part of the mid-Atlantic Ridge, where Vp and Vs

are estimated to be up to 0.5% and 2.4% slower than expected respectively between 50

km and 200 km depth (Foulger et al., 2000). The thinning of the lithospheric crust in

this region also supports the theory of magma upwelling.

The 410 is deeper throughout Iceland (see figure 4.12). The amplitudes of the

P410P phase beneath Iceland decreases from west to east (see figure 5.9a). Another

profile taken along Iceland from the southwest to the northeast shows a depression of

the 410 km discontinuity as well, where the P410P pulses are sharp with narrow pulse

widths (see figure 5.9b). Sharper P410P reflections could also be the result of a

Depth Profile beneath northeast China

410

Dep

th (

km

)


74

smaller velocity gradient between α-olivine and wadsleyite as this boundary tends to

thin at higher temperatures (see figure 2.4).

Figure 5.9a. Depth profile beneath Iceland at 0.5 Hz filter frequency. Starting from

the west(left), moving to the east(right). The 410 remains depressed throughout this

depth profile. As the profile moves to the east, the P410P pulse amplitude

progressively decreases. A persistent low velocity layer is also present directly above

the 410 discontinuity.

Depth Profile beneath Iceland (West to East)

Dep

th (

km

)

410


75

Figure 5.9b. Depth profile beneath Iceland at 0.5 Hz filter frequency. Starting from

the southwest(left), to the northeast(right). This profile taken along Iceland was closer

to the western end of the landform. The P410P pulses are sharp and depressed to ~450

km depth. There also appears to be a persistent low velocity layer above and below the

discontinuity, possibly caused by mantle upwelling.

The 410 is depressed beneath parts of Greenland as well, where the deepest

410 is mapped along the southeastern coast (see figure 4.12). Near the northern coast

of Greenland, a strong negative pulse is observed above and below the P410P positive

pulses in figure 5.10. The P410P phase also appears to reduce in amplitude in the same

area. The remnant warming of the mantle observed beneath Greenland in our study

aligns with the theorized hotspot track traced by Lawver and Muller [1994], where the

residual warming of the upper mantle via Iceland’s hotspot can be seen beneath central

Greenland and the eastern coastline (see figure 4.12). The region closer to Iceland is

likely to have a higher than normal mantle temperature since the mid-Atlantic Ridge

and the hotspot are in the vicinity.

Depth Profile beneath Iceland (Southwest to Northeast)

410 Dep

th (

km

)


76

Figure 5.10. Depth profile beneath Greenland at 0.5Hz filter frequency. Near the

northeastern coast, there is a strong negative pulse is observed below the P410P

positive pulses by using the depth profile. The P410P also appears to reduce in

magnitude in the same area. The region closer to Iceland could reasonably expect to

have a higher than normal mantle temperature.

Depth Profile beneath Greenland

Dep

th (

km

)

410


77

CHAPTER VI

CONCLUSION

The 410 km discontinuity cannot be described as a simple exothermic reaction.

We have found some unexpected and complex variations in the depth and amplitude

of the 410 with respet to the proximity of the discontinuity to various styles of tectonic

activity around the world. We have seen the hypothosized transistional boundary

between α-olivine and wadsleyite to occur at unexpected depths to the 410 and P410P

amplitudes relative to thermal anomalies (inferred from maps of Vs). For instance, the

southern tip of the Kamchatka pennisula has P410P amplitudes that are higher than

expected, even though the 410 appears to be shallow. Otherwise, large and sharp

P410P reflections were observed in most areas where the 410 was deep. Higher P410P

amplitudes with narrow pulse widths are expected when the precursor P wave

encounters a narrow velocity gradient between α-olivine and β-olivine/wadsleyite,

where the narrowest velocity gradient yielded the largest and sharpest P410P

amplitudes. As previously mentioned in the discussion section, a deep 410 and a

narrow depth range for the phase change to occur are expected in warmer mantle

environments. A narrow depth range for the phase change should result in a narrower

P410P with a larger amplitude. Large P410P amplitudes were mapped beneath areas

where warm mantle environments are observed like Hawaii, Lake Baikal, NE China,

Iceland, and Alaska.

While most of our results agree with the results from previous studies

conducted in various regions of Earth. While the warm mantle beneath Hawaii and

Iceland are caused by the presence of a hotspot, the source of upper mantle thermal


78

anomalies in other regions like Lake Baikal and NE China are poorly understood. For

instance, we found that the mantle is abnormally warm beneath China based on the

depth estimates of the 410, but other studies found the mantle to be cooler beneath

China. We have proposed that despite the presence of a cooler subducting slab

penetrating the upper mantle, the 410 could be warmer due to partial melt caused by

the dehydration of the Philippine plate. Another proposed cause for a warm mantle at

the 410 was the presence of a localized convection cell driven by chemical

replenishment in the peridotite depleted mantle through mélange accumulation at the

base of the TZ due to slab buckling or stagnation (see Ringwood, 1991). Another

possible cause for P410P depth and amplitude variations could be due to a chemically

heterogeneous upper mantle, where the mantle may not be pyrolytic in composition

everywhere (Zindler et al., 1984). If the mantle is indeed warmer at the 410, but cooler

at the 660 beneath China, then it would be reasonable to propose that the transition

between olivine and wadsleyite is more than a simple exothermic reaction as the

inclusion of other chemical compounds complicates the velocity gradient pattern of

this boundary even more. The mantle beneath Lake Baikal is another area where there

is much debate as to the behavior of the mantle beneath this rift zone. Some propose

that the Baikal rift zone is passive (Nilsen and Thybo, 2009) while others believe they

have found proof that it is indeed active where the rift is possibly driven by local

mantle convection (Gao et al., 2003 and Zorin et al., 2003). Our study found the 410 to

be deeper than expected beneath Lake Baikal, which supports the hypothesis that the

Baikal rift is active with a convection cell beneath it.


79

Suggestions for future studies

Produce PP and SS precursor maps using 3D velocity modeling

Consider the possibility of iron partitioning in the upper mantle and how it

affects various discontinuities within the upper mantle

Note differences between the region;al depth trends of the 410, 520 and 660,

do they follow the tradition exothermic/endothermic reaction behavior or is the

velocity gradient more complex

Consider a test using a piclogitic mantle model in areas were anomalies occur

Test the velocity gradient of the 410, 520, and 660 on a global scale with a

wider range of frequencies to find the cleanest response without too much

interference from local anomalies


80

WORKS CITED

A. E. Ringwood. Composition and Petrology of the Earth’s Mantle. McGraw-Hill,

New York. 1975.

A. E. Ringwood. Phase transformations and their bearing on the constitution and

dynamics of the mantle. Geochimica et Cosmochimica Acta, vol. 55, pages 2083 –

2110. 1991.

Ainiwaer, Ailiyasi. “Synthetic study in amplitude and frequency content of the P410P

phases from common upper mantle discontinuity.” M.S. Thesis, Texas Tech

University, Lubbock, TX, 2014

Akaogi et al. Olivine-Modified Spinel-Spinel Transitions in the System Mg2SiO4 –

Fe2SiO4: Calorimetric Measurements, Thermochemical Calculation, and Geophysical

Application. Journal of Geophysical Research, vol. 94, pages 15671 – 15685.

November 1989.

Anderson and Bass. Transition region of the Earth’s Upper Mantle. Nature, vol. 320,

pages 321 – 328. March 1986.

Artemieva and Thybo. Deep Norden: Highlights of the lithospheric structure of

Northern Europe, Iceland and Greenland. University of Copenhagen, vol. 31, pages

98 – 106. March 2008.

Arwen Deuss et al. The Nature of the 660-Kilometer Discontinuity in Earth’s Mantle

from Global Seismic Observations of PP Precursors. Science, vol. 311, pages 198 –

201. January 2006.

Arwen Deuss. Global Observations of Mantle Discontinuities Using SS and PP

Precursors. Surv Geophys., vol. 30, pages 301 – 326. May 2009.


81

Bassin et al. The current limits of resolution for surface wave tomography in North

America. EOS, Trans. Am. Geophys. Un., vol. 81. 2000.

Bina and Helffrich. Phase transition Clapeyron slopes and transition zone seismic

discontinuity topography. Journal of Geophysical Research, vol. 99, pages 15853 –

15860. August 1994.

Chambers et al. Reflectivity of the 410-km discontinuity from PP and SS precursors.

Journal of Geophysical Research, vol. 110, pages 1 – 13. February 2005.

Chen et al. Effect of water on olivine-wadsleyite phase boundary in the (Mg, Fe)2SiO4

system. Geophysical Research Letters, vol. 29. September 2002.

Deschamps et al. Anomalies of temperature and iron in the uppermost mantle inferred

from gravity data and tomographic models. Physics of the Earth and Planetary

Interiors, vol. 129, pages 245 – 264. 2002.

Duan et al. Seismic structure and origin of active intraplate volcanoes in Northeast

Asia. Tectonophysics, vol. 470, pages 257 – 266. January 2009.

Duncan, Gregory. “Processing innovations to improve PP precursor analysis and

increase frequency content of studies in the Mid-Pacific.” M.S. Thesis, Texas Tech

University, Lubbock, TX, 2012

Dziewonski and Anderson, Preliminary reference Earth model, Phys. Earth Planet

Inter., vol. 25, 297 – 356. 1981.

Flanagan and Shearer. Global mapping of topography on transition zone velocity

discontinuities by stacking SS precursors. Journal of Geophysical Research, vol. 103,

pages 2673 – 2692. February 1998.


82

Flanagan and Shearer. Topography on the 410-km seismic velocity discontinuity near

subduction zones from stacking of sS, sP, and pP precursors. Journal of Geophysical

Research, vol. 103, pages 21165 – 21182. September 1998.

Flanagan and Shearer. A map of topography on the 410-km discontinuity from PP

precursors. Geophysical Research Letters, vol. 26, pages 549 – 552. March 1999.

Frost and Dolejs. Experimental determination of the effect of H2O on the 410-km

seismic discontinuity. Earth and Planetary Science Letters, vol. 256, pages 182 – 195.

January 2007.

Frost. The Upper Mantle and Transition Zone. Elements, vol. 4, pages 171 – 176. June

2008.

Gao et al. Evidence for small-scale mantle convection in the upper mantle beneath the

Baikal rift zone. Journal of Geophysical Research, vol. 108, pages 1 – 12. April 2003.

Helffrich and Wood. 410-km discontinuity sharpness and the form of the olivine α-β

phase diagram: Resolution of apparent seismic contradictions. Geophysical Journal

International, vol. 126, pages 7 – 12. 1996.

Hirschmann et al. Dehydration melting of nominally anhydrous mantle: The primacy

of partitioning. Physics of the Earth and Planetary Interiors, vol. 176, pages 54 – 68.

April 2009.

HuaiWei et al. Water and partial melting of Earth’s mantle. Science China Earth

Sciences, vol 59, pages 720 – 730. April 2016.

Irifune and Isshiki. Iron partitioning in a pyrolite mantle and the nature of the 410-km

discontinuity. Nature, vol. 392, pages 702 – 704. April 1998.


83

Jacobsen et al. Sound velocities of iron-bearing hydrous ringwoodite. Physics of the

Earth and Planetary Interiors, vol. 143 – 144, pages 47 – 56. 2004.

Katsura et al. Adiabatic temperature profile in the mantle. Physics of the Earth and

Planetary Interiors, vol. 183, pages 212 – 218. July 2010.

Kennett, B.L.N. (Ed.), IASPEI 1991 Seismological Tables, Res. Sch. Of Earth Sci.,

Aust. Natl. Univ., Canberra, 1991.

Kind et al. Seismic Images of Crust and Upper Mantle Beneath Tibet: Evidence for

Eurasian Plate Subduction. Science, vol. 298, pages 1219 – 1221. November 2002.

Lawver and Muller. Iceland Hotspot Track. Geology, vol. 22, pages 311 – 314. April

1994.

Lessing et al. Mantle transition zone structure beneath India and Western China from

migration of PP and SS precursors. Geophysical Journal International, vol. 194, pages

396 – 413. January 2014.

Lessing et al. On the difficulties of detecting PP precursors. Geophysical Journal

International, vol. 201, pages 1666 – 1681. 2015.

Nilsen and Thybo. No Moho uplift below the Baikal Rift Zone: Evidence from a

refraction profile across southern Lake Baikal. Journal of Geophysical Research, vol.

114, pages 1 – 22. August 2009.

Nolet et al. Mantle plume tomography. Chemical Geology, vol. 241, pages 248 – 263.

January 2007.

P. M. Shearer. Global mapping of upper mantle reflectors from long-period SS

precursors. Geophysical Journal International, vol. 115, pages 878 – 904. May 1993.


84

P. M. Shearer. Transition zone velocity gradients and the 520-km discontinuity.

Journal of Geophysical Research, vol. 101, pages 3053 – 3066. February 1996.

Q. Cao et al. Seismic Imaging of Transition Zone Discontinuities Suggests Hot Mantle

West of Hawaii. Science, vol. 332, pages 1068 – 1071. May 2011.

Ritsema et al. S40RTS: a degree-40 shear-velocity model for the mantle from new

Rayleigh wave dispersion, teleseismic travel time and normal-mode splitting function

measurements. Geophysical Journal International, vol. 184, pages 1223 – 1236. 2011.

Rogers III, Kenneth De Witt. “Investigation of the upper mantle beneath the Hawaiian

Island chain using PP precursors.” M.S. Thesis, Texas Tech University, Lubbock, TX,

2013

Rost and Thomas. Array seismology: methods and applications. Rev. Geophys., vol.

40. 2002.

Rost and Weber. The upper mantle transition zone discontinuities in the Pacific as

determined by short-period array data. Earth Planet. Sci. Lett., vol. 204, pages 347 –

361. 2002.

Schafer et al. Depth variations of the 410 and 520 km-discontinuities beneath Asia and

the Pacific from PP precursors. Geophysical Research Letters, vol. 36. June 2009.

Schmerr and Thomas. Subducted lithosphere beneath the Kuriles from migration of

PP precursors. Earth Planetary Science Letters, vol. 311, pages 101 – 111. 2011.

Shearer and Masters. Global mapping of topography on the 660-km discontinuity.

Nature, vol. 355, pages 791 – 795. February 1992.


85

Shearer et al. Experiments in migration processing of SS precursor data to image

upper mantle discontinuity structure. Journal of Geophysical Research, vol. 104,

pages 7229 – 7242. April 1999.

Song et al. P-wave velocity structure under the Changbaishan volcanic region, NE

China, from wide-angle reflection and refraction data. Tectonophysics, vol. 433,

pages 127 – 139. February 2007.

Stixrude. Structure and sharpness of phase transitions mantle discontinuities. J.

Geophys. Res., vol. 102, pages 14835 – 14852. 1997.

Suetsugu et al. Depths of the 410-km and 660-km discontinuities in and around the

stagnant slab beneath the Philippine Sea: Is water stored in the stagnant slab?

Physics of the Earth and Planetary Interiors, vol. 183, pages 270 – 279. September

2010.

Thomas and Billen. Upper mantle structure along a profile in the southwest Pacific.

Geophysical Journal International, vol. 176, pages 113 – 125. 2009.

van der Meijde et al. Seismic evidence for water deep in the Earth’s upper mantle.

Science, vol. 300, pages 1556 – 1558. 2003.

Vidale and Benz. Upper-mantle seismic discontinuities and the thermal structure of

subduction zones. Nature, vol. 356, pages 678 – 683. 1992.

Vink. A hotspot model for Iceland and the Voring Plateau. Journal of Geophysical

Research, vol. 89, pages 9949 – 9959. 1984.

Wood. The effect of water on the 410-kilometer seismic discontinuity. Science, vol.

268, pages 74 – 76. 1995.


86

Xueqing Li and Xiaohui Yuan. Receiver functions in northeast China – implications

for slab penetration into the lower mantle in northwest Pacific subduction zone. Earth

and Planetary Science Letters, vol. 216, pages 679 – 691. September 2003.

Yinshuang et al. The crust and upper mantle discontinuity structure beneath Alaska

inferred from receiver functions. Physics of the Earth and Planetary Interiors, vol. 150,

pages 339 – 350. January 2005.

Yusa and Inoue. Compressibility of hydrous wadsleyite (β-phase) in Mg2SiO4 by high

pressure x-ray diffraction. Geophysical Research Letters, vol. 24, pages 1831 – 1834.

1997.

Zhao et al. Imaging the subducting slabs and mantle upwelling under the Japan

Islands. Geophysical Journal International, vol. 190, pages 816 – 828. May 2012.

Zindler et al. Isotope and trace element geochemistry of young Pacific seamounts:

implications for the scale of upper mantle heterogeneity. Earth and Planetary Science

Letters, vol 70, pages 175 – 195. June 1984.

Zorin et al. The Baikal rift zone: the effect of mantle plumes on older structure.

Tectonophysics, vol. 371, pages 153 – 173. May 2003.


87

APPENDIX A

USER MANUALS

Run_addPPtimes Manual

This program is the macro for the 3D ray tracing. It runs Nathan's program and outputs

the data to be read in the following programming step. It also references nearly all

other programs that are affilitated with the raytracing process. global zz zz=[0 2 4 6 8 10:5:50 60:20:180 190:10:250 270:20:370 380:10:450

470:40:630 650:20:700 ]; %generates a pre-determined list of depths

we're interested in DIRS=dir([drive_letter ':\CUT_PP\' year]) %places us in the correct

directoyy for i=3:length(DIRS) if exist([drive_letter ':\CUT_PP\' year '\' DIRS(i).name

'\Good_files.mat'])~=0 load([drive_letter ':\CUT_PP\' year '\' DIRS(i).name

'\Good_files.mat']) %if Good_files exists in the directory, then we

want to load it

fid2=fopen('PP_raypath.txt','wt'); % seting up command line

for event to station with all raypath info for jmp=1:length(Good_files) % this loop sets up a bunch of

files to run nathan's program for midpoints RFNT=Good_files(jmp).name; RFNT=[drive_letter RFNT(2:end)]; dashes=find(RFNT=='/'); %correcting the slash issue that

arrises when QCing is done on a Mac RFNT(dashes)='\'; RFN(jmp).name=RFNT; load(RFN(jmp).name) mpname(jmp).name=set_up_midpoints(data,jmp); clear data fprintf(fid2,['-raytrace LLNL-G3D-JPS.e3d.binary '

mpname(jmp).name ' -p']); %initiates the 3D raytracing program. The

binary file must be in the path. fprintf(fid2,['\n']); end fprintf(fid2,['quit']); fclose(fid2); system('java -jar LLNL-Earth3D.5.3.jar < PP_raypath.txt') %

runs nathan's program for midpoints. The system command is needed

because it is written in java.

for jmp=1:length(Good_files) % reads files and gets midpoints if exist([mpname(jmp).name '.0.ascii.path']) > 0.5

[mlat(jmp),mlon(jmp),mz(jmp)]=get_midpoints(mpname(jmp).name); %calls

the get_midpoints program end end


88

fid5=fopen('PP_time_midpoint_to_station.txt','wt');% doing

the raytracing for each depth of the midpoint - here we are opening

the times for the first leg of the PP wave for jmp=1:length(Good_files) load(RFN(jmp).name)

[zzname(jmp).name,]=make_ZZ_vs_TT(data,mlon(jmp),mlat(jmp),mz(jmp),jm

p); %makes zzname file, times to the depths fprintf(fid5,['-raytrace LLNL-G3D-JPS.e3d.binary '

zzname(jmp).name]); %running Nathan's program fprintf(fid5,['\n']); end fprintf(fid5,['quit']); fclose(fid5) % done setting up command to trace from

midpoint to station P wave system('java -jar LLNL-Earth3D.5.3.jar <

PP_time_midpoint_to_station.txt') % initiating Nathan's program to

trace from midpoint to station

for jmp=1:length(Good_files) hi=RFN(jmp).name if exist([zzname(jmp).name '.TT']) > 0.5 % if the zzname

file exists, then load it to read data load(RFN(jmp).name) [ZZ,dtt]=get_ZZ_vs_TT(zzname(jmp).name,mz(jmp)); %

run the program to make the list of ZZ vs dtt data(3).zz=ZZ; % want the data in the 3rd channel data(3).dtt=dtt; D=data(3); clear data data=D; clear D dashes=find(RFN(jmp).name=='\'); RFN_out=[RFN(jmp).name(1:dashes(1)) 'CUT_PP_TT'

RFN(jmp).name(dashes(2):end)]; RFN_out(1)=outdrive; dashes=find(RFN_out=='\'); % fixes the slash issue on

Macs mkdir(RFN_out(1:dashes(end))) % makes a new directory

for the new data save(RFN_out,'data') %saves it end end end % each time this function is ran, the used data needs to be

deleted so that it isn't overwritten !del mid_pt*txt* !del Z_vs_t*txt* !del *txt !del *TT !del *TT end end


89

PPnSSbeamer Manual

In order to run this program you must open run_PPnSSbeamer to access enter the

input variables, ‘PP’ for phase is the first input, the second input is the drive letter you

want the program to write all of your outputs onto.

The time at which the PP precursor arrives is ideally around 200 seconds after the

initial impulse, so time_total or the total length of the time window for PP waves is set

at 289 seconds

D_dirs(i).name is used to categorize the highsn, others and lowsn folders as i=1,2,3

respectively for filtering data and obtaining amplitudes for convolution and stacking

later on

Empty matricies for Df (filtered deoceaned source functions) and MIDPT(midpoint

information) is created for future storage

If any of the event folders exists (highsn, lowsn, or others) the program changes the

working directory form the home directory(home1, the home directory always starts

from the year of interest) to the event folder with highsn, lowsn or others.

After the program locates all BHZ networks, it enters the “for” loop processing one

file for the event at a time. If the depth of the earthquake is greater than 60 km, it is

classified as an event that is too deep, the “for” loop ends and that event is skipped,

moving on to the next.

The variable iFs was set to zero before the start of the “for” loop. This is a counter,

keeping track of the number of events that are shallow enough to be used for the rest

of the data processing procedure.

The variable dt is assigned to store the value of the current sample rate of the data file

The variable Dt is assigned to store amplitude values for the data file

Using GFILT_2015: Dt is high pass filtered at 0.02 Hz using the raw data, the now

filtered data file is filtered again, but instead it is low pass filtered at 0.5Hz. These two

filtering steps effective apply a taper to the upper and lower limits of the raw data

The filtered data is interpolated to a sample rate of 0.05s/sample. The impulses for the

new filtered data is saved as variable Dft

The newly interpolated data is low pass filtered again at 0.5 Hz

Amplitudes/impulses are saved as within each file as data.data

The sample rate is saved within each file as data.dt

Variables w1 and w2 are the time windows for the PP precursor wave (10 s prior and

30 s after)

Impulses from the filtered waveform is saved as variable Df, noise from the Dft

version is minimized via standard deviation of the Dft wave, the same applies for Do,

except the Dt wave is used

Event names, locations and station locations are all saved with the “data” array as

NAME, EVENT and STA respectively

Once the information for the data file is saved within itself as the structured array

“data”, the directory is changed to the “Nathan_7_22_2016_Beamform” folder (this

enables raytracing and beamforming to run concurrently without overwriting midpoint

data being created in via the 3D raytracing program “run_addPPtimes”)


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Midpoint location, depth and elevation relative to sea level are all obtained via

“run_trace_PPmidptWB” with event and station information used as input variables.

EVENT is transposed in order to store the information in rows instead of columns

Midpoint data is saved within the empty vector, MIDPT, created at the beginning of

the program

All waveforms with dry or wet bouncepoints are convolved with themselves. The

maximum amplitudes are stored as variable “Afm”. Dry bouncepoints are defined by a

waveform with a midpoint beneath a landform. Wet bouncepoints are defined by

waveforms with midpoints in the middle of the ocean *note: the flipud command is

flipping the impulses/amplitudes for the waveform left/right instead of up/down due

to the impulse data being stored in a row instead of a column

Elevation of the bouncepoints for each data set are sorted from shallowest to deepest

elevation. All data information within each file is saved within itself as the structured

array “data(1)”

All wet bouncepoints are found via traced raypaths with the use of the

PP_RM_OCEAN program. The raw and filtered source functions are saved as Do_S

and Df_S respectively

Maximum amplitudes are saved as MAX_AMPs for events with highsn and others

Data is stacked through different degree bin sizes and saved as channels 11, 12, 13, 14,

15 and 16 (0.5°,1°, 2°, 4°, 8° and 16° respectively). These channels are reserved for

the raw data

Channels 35 and 36 are deoceaned source functions

AMPjunk is reserved for amplitudes and IMA is reserved at the location for each

impulse

BEAM180ALL.mat file is saved within every event folder

All deoceaned filtered data is stacked within 180° bins and stored as channel 40

(GCarc)

BEAM180ALL and BEAM180DRY are saved as mat files in every event folder

All event information is saved as the structured array “EVENTS_INFO”


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APPENDIX B

QC GUIDE

Attiya Darensburg

QC guide

For the PP data we are focusing on the 410 km discontinuity since it appears the

clearest for the PP phase. In particular, we want to select seismic wave forms where

the greatest impulse happens exactly at or close to 200 seconds (highlighted by a

vertical red line). Some seismograms can be tricky. If an impulse is observed toward

the beginning of the event 0 seconds or close to this point, it can be assumed that this

is the direct P wave arrival, an entirely separate event that will not interfere with other

impulses that may be observed later on in time for the same event. Some events that

have impulses which do not occur exactly when we would prefer might be acceptable

under certain circumstances, for instance, if there are a low number of wave forms or

seismograms for a particular event, forcing one to be more lenient in regards to which

seismogram should be kept. The following screenshots are examples of events that

should be accept or rejected as well as under which circumstances these decisions may

change.

When you run ALL_PPnSSCLEAN2_2015, the first input is the channel followed by

the phase type (ALL_PPnSSCLEAN2_2015(‘3’,’PP’)). The channel for all PP phases

is channel 3; channel 2 is reserved for SS phases.

The following screenshots are examples of seismograms that should be kept:


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The greatest amplitudes for these 3 seismograms appears to occur around 210 seconds,

but more importantly, the initial drastic increase in amplitude is observed exactly at

200 seconds

The next series of screenshots are examples of seismograms that should fall in the “in

between” category of whether they should be kept or rejected. This decision typically

depends on how many seismograms are left in a given event and how many have

already been accepted:


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There’s an initial impulse at 200 seconds here, but the noise that follows is

troublesome. In this particular instance there were over 300 events left with some

events already selected, so it would be ok if this seismogram was rejected since there

are so many others of good quality to choose from. If there were a total of 20 to 40

seismograms for this entire event with only a few functions left to choose from, this

event should be kept, especially if there are only a few other or no accepted

seismograms for this particular event.


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The largest impulse is observed at 220 seconds, but this should still be kept as the

largest amplitude does not arrive too late.


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Here’s another example of a seismogram that should be accepted only if your options

are limited. Notice how the largest observed impulse is after 260 seconds, long after

the PP arrival, but the initial impulse is around 210 seconds. The second arrival is

possibly far enough away to not interfere with the P410P arrival.


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This impulse arrives a little early, but once again, it would be a good seismogram to

keep if your options are limited. For seismograms in the “others” category it would be

a good idea to keep this seismogram since it is not likely that you will find nice and

clean signals as often as you would with “highsn” files.

*One thing to keep in mind is that the reason why we QC is to find waveforms that

can be successfully cross correlated with high quality waveforms. It is better to error

on the side of caution and accept “good enough” seismograms when you don’t have a

lot of options. When these individual selected events are processed through the cross

correlation program, the seismograms that could not cross correlate with a value equal

to or greater than 0.6 or 0.7 are rejected anyway.

The following screenshots are examples of bad seismograms that should be rejected:


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The signal to noise ratio is preferably high, but the initial impulse occurs way too late

at 240 seconds


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Noisy


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2017, Attiya Darensburg - TDL

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